World Meteorological Organization

Global Ozone Research and Monitoring Project—Report No. 50

SCIENTIFIC ASSESSMENT OF

OZONE DEPLETION: 2006

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

Final Release: February 2007

From Scientific Assessment of Ozone Depletion: 2006

This report is dedicated to the memory of our friend and colleague, Prof. Gérard Mégie of the

Centre National de la Recherche Scientifique of Paris, France.

Gérard worked as a member of the Scientific Assessment Panel of the Montreal Protocol since the

very beginning; he was a participant in the Les Diablerets Panel Review Meeting in 1989. Later, he

became a Cochair of the Scientific Assessment Panel, cochairing the 1998 and 2002 assessment reports.

To that role, he brought his rare combination of scientific excellence and leadership talent, built upon

many years as a teacher, researcher, internationally acclaimed scientist, and leader in the atmospheric science community.

One mark of Gérard’s insightfulness was his work to guide the Panel toward its increasing

emphasis on the connections between the ozone layer and the climate system. The Panel’s assessment

report for 2006 does contain that emphasis, and indeed it is captured in separate chapters on these issues.

Gérard was dedicated to communicating scientific understanding at many levels, ranging from

international decisionmaking to the realm of the general public and schools. This is exemplified by his

thoughtful work associated with the Panel’s efforts to provide a set of “frequently asked” questions and

answers about the ozone layer.

Gérard’s many contributions to atmospheric science, as well as his role in forging interactions of the

scientific community at its interface with society, have had a lasting impact. He leaves a legacy of scientific leadership that has helped shape the world’s actions on the ozone layer. This legacy will endure.

Dedicated to the Memory of

Gérard Mégie

1946-2004

v

Ayité-Lô Nohende Ajavon Togo

Hideharu Akiyoshi Japan

Daniel L. Albritton USA

Stephen O. Andersen USA

John Anderson USA

Gustavo A. Argüello Argentina

Antti Arola Finland

Paul Ashford UK

Elliot Atlas USA

Pieter J. Aucamp South Africa

John Austin USA

Gnon Baba Togo

Alkiviadis F. Bais Greece

Mark P. Baldwin USA

Leonard A. Barrie Switzerland

David Battisti USA

LIST OF INTERNATIONAL AUTHORS,

CONTRIBUTORS, AND REVIEWERS

Assessment Cochairs

Ayité-Lô Nohende Ajavon

Daniel L. Albritton

Robert T. Watson

Assessment Scientific Steering Committee

Marie-Lise Chanin

Susana B. Diaz

John A. Pyle

A.R. Ravishankara

Theodore G. Shepherd

Chapters and Lead Authors

I. Source Gases

Chapter 1. Long-Lived Compounds (Cathy Clerbaux and Derek Cunnold)

Chapter 2. Halogenated Very Short-Lived Substances (Katherine S. Law and William T. Sturges)

II. Ozone Changes

Chapter 3. Global Ozone: Past and Present (Martyn P. Chipperfield and Vitali E. Fioletov)

Chapter 4. Polar Ozone: Past and Present (Paul A. Newman and Markus Rex)

III. Future Expectations for Ozone, Ozone-Depleting Substances, and UV

Chapter 5. Climate-Ozone Connections (Mark P. Baldwin and Martin Dameris)

Chapter 6. The Ozone Layer in the 21st Century (Greg Bodeker and Darryn W. Waugh)

Chapter 7. Surface Ultraviolet Radiation: Past, Present, and Future (Alkiviadis F. Bais and Dan Lubin)

Chapter 8. Halocarbon Scenarios, Ozone Depletion Potentials, and Global Warming Potentials

(John S. Daniel and Guus J.M. Velders)

Twenty Questions and Answers About the Ozone Layer: 2006 Update (David W. Fahey)

Coordinating Editor

Christine A. Ennis

Authors, Contributors, and Reviewers
vi

AUTHORS, CONTRIBUTORS, AND REVIEWERS

Gufran Beig India

Slimane Bekki France

Peter Bernath Canada

Germar Bernhard USA

Lars Olof Björn Sweden

Donald R. Blake USA

Nicola J. Blake USA

Thomas Blumenstock Germany

Mario Blumthaler Austria

Greg Bodeker New Zealand

Rumen D. Bojkov Germany

Janet F. Bornman New Zealand

Geir O. Braathen Switzerland

Peter Braesicke UK

Bram Bregman The Netherlands

Christoph Brühl Germany

Claus Brüning Belgium

James B. Burkholder USA

John P. Burrows Germany

Neal Butchart UK

James H. Butler USA

André Butz Germany

James M. Calm USA

Pablo O. Canziani Argentina

Daniel Cariolle France

Lucy Carpenter UK

Kenneth S. Carslaw UK

Marie-Lise Chanin France

Martyn P. Chipperfield UK

Natalia E. Chubarova Russia

Hans Claude Germany

Cathy Clerbaux France

Pierre-Francois Coheur Belgium

William J. Collins UK

Brian J. Connor New Zealand

David B. Considine USA

Eugene Cordero USA

R. Anthony Cox UK

John Crowley Germany

Paul Crutzen Germany

Derek M. Cunnold USA

Martin Dameris Germany

John S. Daniel USA

Michael Danilin USA

Philippe Demoulin Belgium

Susana B. Diaz Argentina

Marcel Dorf Germany

Anne R. Douglass USA

Katja Drdla USA

Ellsworth G. Dutton USA

Geoffrey S. Dutton USA

Kalju Eerme Estonia

Tatiana Egorova Switzerland

Andreas Engel Germany

Christine A. Ennis USA

Carynelisa Erlick Israel

David Etheridge Australia

Wayne Evans USA

Veronika Eyring Germany

David W. Fahey USA

Uwe Feister Germany

Wuhu Feng UK

Vitali E. Fioletov Canada

Horst Fischer Germany

Lawrence E. Flynn USA

Ian Folkins Canada

Piers M.D. Forster UK

Paul J. Fraser Australia

John Frederick USA

Katja Frieler Germany

Stacey Hollandsworth Frith USA

Lucien Froidevaux USA

Jan Fuglestvedt Norway

Rolando R. Garcia USA

Marvin A. Geller USA

Andrew Gettelman USA

H. Peter Gies Australia

Nathan P. Gillett UK

Marco A. Giorgetta Germany

Sophie Godin-Beekmann France

David M. Golden USA

Marco González Kenya

Florence Goutail France

Hans-F. Graf UK

Claire Granier USA/France

Julian Gröbner Switzerland

Serge Guillas USA

Joanna D. Haigh UK

Neil R.P. Harris UK

Dennis Hartmann USA

Alain Hauchecorne France

Didier A. Hauglustaine France

Sachiko Hayashida Japan

Peter H. Haynes UK

Jay Herman USA

David J. Hofmann USA

Lon Hood USA

Dale Hurst USA

Abdelmoneim A.R. Ibrahim Egypt

Mohammad Ilyas Malaysia

Takashi Imamura Japan

Ivar Isaksen Norway

Charles H. Jackman USA

Elena Jimenez Spain
vii

AUTHORS, CONTRIBUTORS, AND REVIEWERS

Paul Johnston New Zealand

Roderic L. Jones UK

Hiroshi Kanzawa Japan

David Karoly USA

Jussi Kaurola Finland

S. Randy Kawa USA

Jack A. Kaye USA

M.A.K. Khalil USA

Douglas E. Kinnison USA

Bjørn Knudsen Denmark

Malcolm K.W. Ko USA

Chester Koblinsky USA

Peter Koepke Germany

Laurie J. Kovalenko USA

Karin Kreher New Zealand

Nickolay A. Krotkov USA

Paul B. Krummel Australia

Janusz W. Krzy ́scin Poland

Lambert Kuijpers The Netherlands

Michael J. Kurylo USA

Esko Kyrö Finland

Shyam Lal India

Ulrike Langematz Germany

Kathleen O. Lantz USA

Katherine S. Law France

Mark G. Lawrence Germany

J. Ben Liley New Zealand

Zenobia Litynska Poland

Nathaniel J. Livesey USA

Jennifer Logan USA

Manuel López-Puertas Spain

Dan Lubin USA

Emmanuel Mahieu Belgium

Eva Mancini Italy

Gloria L. Manney USA

Alistair Manning UK

Elisa Manzini Italy

Marion Marchand France

Céline Mari France

Katja Matthes USA

Bernhard Mayer Germany

Archie McCulloch UK

Mack McFarland USA

Richard L. McKenzie New Zealand

Richard D. McPeters USA

Davit Melkonyan Armenia

Pauline Midgley Germany

Alvin J. Miller USA

John Miller USA

Kenneth R. Minschwaner USA

Stephen A. Montzka USA

Jens Mühle USA

Rolf Müller Germany

Nzioka John Muthama Kenya

Hideaki Nakane Japan

Eric R. Nash USA

Ray Nassar Canada

Michael Newchurch USA

Paul A. Newman USA

Ole John Nielsen Denmark

Simon O’Doherty UK

Alan O’Neill UK

Samuel Oltmans USA

David E. Oram UK

Yvan J. Orsolini Norway

Eduardo Palenque Bolivia

Panos Papagiannakopoulos Greece

Steven Pawson USA

Thomas Peter Switzerland

Irina Petropavlovskikh USA

Klaus Pfeilsticker Germany

Rubén D. Piacentini Argentina

Giovanni Pitari Italy

John M.C. Plane UK

Ulrich Platt Germany

Ian Plumb Australia

R. Alan Plumb USA

David Plummer Canada

Lorenzo M. Polvani USA

Lamont R. Poole USA

Robert W. Portmann USA

Gilles Poulet France

Michael Prather USA

Ronald Prinn USA

John A. Pyle UK

S. Ramachandran India

V. Ramaswamy USA

Cora E. Randall USA

William Randel USA

A.R. Ravishankara USA

Thomas Reddmann Germany

Claire Reeves UK

Stefan Reimann Switzerland

Ellis E. Remsberg USA

Markus Rex Germany

Robert Rhew USA

Curtis P. Rinsland USA

Alan Robock USA

José M. Rodríguez USA

Geert-Jan Roelofs The Netherlands

Bjørg Rognerud Norway

Howard Roscoe UK

Karen Rosenlof USA

Eugene Rozanov Switzerland
viii

AUTHORS, CONTRIBUTORS, AND REVIEWERS

James M. Russell III USA

Ross J. Salawitch USA

Michelle L. Santee USA

Benjamin D. Santer USA

Toru Sasaki Japan

Yasuhiro Sasano Japan

Robert Sausen Germany

Cornelius Schiller Germany

Ulrich Schmidt Germany

Robyn Schofield USA

Ulrich Schumann Germany

John F. Scinocca Canada

Robert B. Scott USA

Gunther Seckmeyer Germany

Megumi Seki Kenya

Theodore G. Shepherd Canada

Kiyotaka Shibata Japan

Drew Shindell USA

Keith P. Shine UK

Emily Shuckburgh UK

Peter Simmonds UK

Isobel Simpson USA

Howard J. Singer USA

Cynthia Shaw Singleton USA

Björn-Martin Sinnhuber Germany

Miriam Sinnhuber Germany

Christopher E. Sioris USA

Pavla Skrivankova Czech Republic

Harry Slaper The Netherlands

James R. Slusser USA

Sergei P. Smyshlyaev Russia

Susan Solomon USA

Johannes Staehelin Switzerland

Knut Stamnes USA

Wolfgang Steinbrecht Germany

Andrea Stenke Germany

Richard M. Stimpfle USA

Richard S. Stolarski USA

Frode Stordal Norway

Susan Strahan USA

Hamish Struthers New Zealand

William T. Sturges UK

Azadeh Tabazadeh USA

Masaaki Takahashi Japan

David W. Tarasick Canada

Yukio Terao USA

Larry W. Thomason USA

David W.J. Thompson USA

Wenshou Tian UK

Simone Tilmes USA

Claudisa Timmreck Germany

Darin Toohey USA

Cathy Trudinger Australia

Jan C. van der Leun The Netherlands

Costas Varotsos Greece

Guus J.M. Velders The Netherlands

Daniel P. Verdonik USA

Mario Visca Italy

Roland von Glasow Germany

Marc von Hobe Germany

John Michael Wallace USA

Timothy J. Wallington USA

Nicola J. Warwick UK

Robert T. Watson USA

Darryn W. Waugh USA

Elizabeth C. Weatherhead USA

Ann Webb UK

Mark Weber Germany

Philipp Weihs Austria

Debra K. Weisenstein USA

Ray F. Weiss USA

Paul Wennberg USA

Paul Wine USA

Stephen Wood New Zealand

Donald J. Wuebbles USA

Eun-Su Yang USA

Margarita Yela Spain

Shigeo Yoden Japan

Yoko Yokouchi Japan

Vladimir Yushkov Russia

Shari A. Yvon-Lewis USA

Rudolphe Zander Belgium

Francesco Zaratti Bolivia

Joseph M. Zawodny USA

Christos S. Zerefos Greece

Tong Zhu China
ix

CONTENTS

SCIENTIFIC ASSESSMENT OF OZONE DEPLETION: 2006

PREFACE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xiii

EXECUTIVE SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xvii

I. SOURCE GASES

CHAPTER 1: LONG-LIVED COMPOUNDS

Lead Authors: Cathy Clerbaux and Derek Cunnold

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

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

1.2 Halogenated Ozone-Depleting Gases in the Atmosphere. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.7

1.3 Emissions Estimated from the Halocarbon Measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.38

1.4 Other Trace Gases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.46

References. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.52

CHAPTER 2: HALOGENATED VERY SHORT-LIVED SUBSTANCES

Lead Authors: Katherine S. Law and William T. Sturges

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

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

2.2 Sources, Distributions, and Trends of VSLS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.9

2.3 Atmospheric Chemistry of VSLS. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.19

2.4 Dynamics and Transport in the Tropopause Region and Implications for VSLS. . . . . . . . . . . . . . . . . . . . . . 2.24

2.5 Contribution of Halogenated VSLS to Stratospheric Halogen Loading . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.31

2.6 Potential Impact of VSLS on Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.43

2.7 Potential Future Considerations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.45

References. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.46

II. OZONE CHANGES

CHAPTER 3: GLOBAL OZONE: PAST AND PRESENT

Lead Authors: Martyn P. Chipperfield and Vitali E. Fioletov

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

3.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.5

3.2 Update of Ozone Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.5

3.3 Summary of Other Observations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.19

3.4 Understanding and Interpretation of Ozone Changes. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.21

References. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.41

Appendix 3A: Description of the Ozone Datasets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.55
CHAPTER 4: POLAR OZONE: PAST AND PRESENT

Lead Authors: Paul A. Newman and Markus Rex

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

4.0 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.3

4.1 Polar Stratospheric Observations Update. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.4

4.2 Progress in Our Understanding of the Physical and Chemical Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.20

4.3 Recent Polar Ozone Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.30

4.4 The Influence of Precipitating Charged Particles on Polar Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.35

References. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.37

III. FUTURE EXPECTATIONS FOR OZONE, OZONE-DEPLETING SUBSTANCES, AND UV

CHAPTER 5: CLIMATE-OZONE CONNECTIONS

Lead Authors: Mark P. Baldwin and Martin Dameris

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

5.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3

5.2 Coupling of the Stratosphere and Troposphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3

5.3 Effects of Anthropogenic Climate Change and of Emissions on Stratospheric Ozone . . . . . . . . . . . . . . . . . 5.18

References. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.38

Appendix 5A: AOGCMs Used in Chapter 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.49

CHAPTER 6: THE OZONE LAYER IN THE 21ST CENTURY

Lead Authors: Greg Bodeker and Darryn W. Waugh

Scientific Summary. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.1

6.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.5

6.2 A Framework for Evaluating Changes in Ozone Abundances . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.5

6.3 Factors Affecting the Detection, Attribution, and Timing of Milestones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.8

6.4 Statistical Methods for Detection of Milestones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.11

6.5 Attribution of the Recent Behavior of Ozone. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.13

6.6 Projections of the Future Behavior of Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.20

References. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.37

Appendix 6A: Model Scenarios. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.43

CHAPTER 7: SURFACE ULTRAVIOLET RADIATION: PAST, PRESENT, AND FUTURE

Lead Authors: Alkiviadis F. Bais and Dan Lubin

Scientific Summary. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.1

7.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.3

7.2 Factors Affecting UV Radiation: New Findings . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.3

7.3 Resources for Studying Surface UV Radiation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.13

7.4 UV Climatology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.21

7.5 Expectations for the Future . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.29

7.6 Remaining Questions and Uncertainties . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.32

References. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.32

Appendix 7A: Spectral Data Available from Databases. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.49

Appendix 7B: Internet Addresses for UV Sites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.53

x

CONTENTS
CHAPTER 8: HALOCARBON SCENARIOS, OZONE DEPLETION POTENTIALS,

AND GLOBAL WARMING POTENTIALS

Lead Authors: John S. Daniel and Guus J.M. Velders

Scientific Summary. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.1

8.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.5

8.2 Halocarbon Lifetimes, Ozone Depletion Potentials, and Global Warming Potentials . . . . . . . . . . . . . . . . . . . 8.5

8.3 Future Halocarbon Source Gas Concentrations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.13

8.4 Other Processes Relevant to Future Ozone Evolution . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.34

8.5 Indirect GWPs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.35

References. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.36

TWENTY QUESTIONS AND ANSWERS ABOUT THE OZONE LAYER: 2006 UPDATE

Lead Author: David W. Fahey

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

II. The Ozone Depletion Process . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.9

III. Stratospheric Ozone Depletion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.22

IV. Controlling Ozone-Depleting Gases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.32

V. Implications of Ozone Depletion. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.37

VI. Stratospheric Ozone in the Future . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.43

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

xi

CONTENTS

xiii

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 Earth’s 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 Assessment Panel, and the

Technology and Economic Assessment Panel have each periodically prepared major assessment reports that updated the

state of understanding in their purviews. These reports have been scheduled so as 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 present 2006 report is the latest in a series of ten 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

the United Nations Environment Programme (UNEP). This report is the sixth 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 Adjustment 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 Adjustment

and Amendment

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

PREFACE

1995 Vienna Adjustment

1997 Montreal Adjustment

and Amendment

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

1999 Beijing Adjustment 

and Amendment

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

2006 Scientific Assessment of Ozone Depletion: 2006. WMO No. 50

2007 19th Meeting of the Parties

The Current Information Needs of the Parties

The genesis of Scientific Assessment of Ozone Depletion: 2006 occurred at the 15th Meeting of the Parties to the

Montreal Protocol in Nairobi, Kenya, at which the scope of the scientific needs of the Parties was defined in their

Decision XV/53 (4), which stated that “...for the 2006 report, the Scientific Assessment Panel should consider issues

including:

(a) Assessment of the state of the ozone layer and its expected recovery;

(b) Evaluation of specific aspects of recent annual Antarctic ozone holes, in particular the hole that occurred in

2002;

(c) Evaluation of the trends in the concentration of ozone-depleting substances in the atmosphere and their consistency with reported production and consumption of ozone-depleting substances;

(d) Assessment of the impacts of climate change on ozone-layer recovery;

(e) Analysis of atmospheric concentrations of bromine and the likely quantitative implications of the results on

the state of the ozone layer;

(f) 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 also into account the expected

impacts of climate change...”

The Assessment Process

The formal planning of the current assessment was started early in 2005. A Scientific Steering Committee consisting of five scientists from various countries assisted the Scientific Assessment Panel’s Cochairs throughout the planning, preparation, review, and finalization of the 2006 assessment. The Cochairs and Scientific Steering Committee considered suggestions from the Parties regarding experts from their countries who could participate in the process.

Furthermore, an ad hoc international scientific group also suggested participants from the world scientific community. In

addition, this group contributed to crafting the outline of the assessment report. As in previous assessments, the participants represented experts from the developed and developing world. In addition to the scientific expertise, the developing country experts bring a special perspective to the process, and their involvement in the process has also contributed

to capacity building.

The information of the 2006 assessment is contained in eight chapters associated with ozone-layer topics: 

I. SOURCE GASES

Chapter 1. Long-Lived Compounds

Chapter 2. Halogenated Very Short-Lived Substances
II. OZONE CHANGES

Chapter 3. Global Ozone: Past and Present

Chapter 4. Polar Ozone: Past and Present

III. FUTURE EXPECTATIONS FOR OZONE, OZONE-DEPLETING SUBSTANCES, AND UV

Chapter 5. Climate-Ozone Connections

Chapter 6. The Ozone Layer in the 21st Century

Chapter 7. Surface Ultraviolet Radiation: Past, Present, and Future

Chapter 8. Halocarbon Scenarios, Ozone Depletion Potentials, and Global Warming Potentials

A special resource for the Panel’s work was the 2005 Special Report of the Intergovernmental Panel on Climate

Change (IPCC) and the Technology and Economics Assessment Panel (TEAP) of the Montreal Protocol, Safeguarding

the Ozone Layer and the Global Climate System: Issues Related to Hydrofluorocarbons and Perfluorocarbons. This

report was used as a basis for many scenarios for modeling runs and hypothetical cases that are included here.

The initial plans for the chapters of the 2006 Scientific Assessment Panel’s report were examined at a meeting that

occurred on 26-27 July 2005 in Paris, France. The Lead Authors, Steering Committee, and Cochairs focused on the content of the draft chapters and on the need for coordination among the chapters.

The first drafts of the chapters were examined at a meeting that occurred on 30 November-2 December 2005 in

Herndon, Virginia, United States, at which the Lead Authors, Steering Committee, 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 125 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 19-23 June

2006, the responses to these mail review comments were proposed by the authors and discussed by the 77 participants.

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

on the UNEP web site on 18 August 2006) was prepared and completed by the attendees of the Les Diablerets meeting.

The 2006 State-of-Understanding Report

In addition to the scientific chapters and the Executive Summary, the assessment also updates the 2002 assessment

report’s answers to 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 updated 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: 2006 is the product of 310 scientists from 34 countries who contributed to its

preparation and review1 (205 scientists prepared the report and 183 scientists participated in the peer review process).

What follows is a summary of their current understanding of the stratospheric ozone layer and its relation to

humankind.

xv

PREFACE

1 Participants were from Argentina, Armenia, Australia, Austria, Belgium, Bolivia, Canada, Czech Republic, Denmark, Egypt, Estonia, Finland,

France, Germany, Greece, India, Israel, Italy, Japan, Kenya, Malaysia, New Zealand, Norway, Poland, Russia, South Africa, Spain, Sweden,

Switzerland, The Netherlands, The People’s Republic of China, Togo, United Kingdom, and United States of America.

xvii

EXECUTIVE SUMMARY

INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xix

Figure 1: Ozone-Depleting Substances, the Ozone Layer, and UV Radiation: Past, Present, and Future . . . . xix

RECENT MAJOR FINDINGS AND CURRENT SCIENTIFIC UNDERSTANDING . . . . . . . . . . . . . . . . . . . . . . . xx

Ozone-Depleting Substances . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xx

Ozone, Ultraviolet Radiation, and Other Related Changes to Date . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xxi

The Future of the Ozone Layer and Surface UV Radiation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xxiv

ADDITIONAL SCIENTIFIC EVIDENCE AND RELATED INFORMATION . . . . . . . . . . . . . . . . . . . . . . . . . . . xxv

Long-Lived Ozone-Depleting Substances . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xxv

Halogenated Very Short-Lived Substances . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xxvii

Global Ozone (60°S-60°N) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xxviii

Polar Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xxix

Impact of Climate Change . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xxxi

The Ozone Layer in the 21st Century . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xxxii

UV Radiation and Its Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xxxiii

Ozone Depletion Potentials and Global Warming Potentials . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xxxiv

IMPLICATIONS FOR POLICY FORMULATION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xxxiv

Contents

Final Release: February 2007

From Scientific Assessment of Ozone Depletion: 2006

xix

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 updated 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 scientific understanding were assessed by the Scientific Assessment Panel in

1989, 1991, 1994, 1998, and 2002. This information helped support discussions among the Parties that led to the subsequent Amendments and Adjustments of the 1987 Protocol. The 2006 Scientific Assessment summarized here is the sixth

in that series. The information contained in this Executive Summary is based on the eight detailed chapters of the full

report.

The previous Assessment presented evidence that the tropospheric abundances of most ozone-depleting substances,

as well as of stratospheric chlorine, were stable or decreasing due to actions taken under the Montreal Protocol (see

schematic Figure 1a, b), with the stratospheric abundances showing a time lag due to the time for surface emissions to

reach the stratosphere. Based on these facts, it was stated that “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.”
EXECUTIVE SUMMARY

xx

An important next step is to ask whether stratospheric ozone and surface ultraviolet (UV) radiation are responding

as expected to the controls on ozone-depleting substances imposed by the Protocol (see schematic Figure 1c, d). In

addressing this question, it is necessary to consider factors other than ozone-depleting substances that also influence

ozone and UV radiation. These factors include natural dynamical variability, volcanic eruptions, solar variations,

aerosols (airborne fine particles), and climate change.

Pre-1980 values are often used as a benchmark for ozone and UV recovery. However, because of the abovementioned factors, if and when ozone and UV radiation return to their pre-1980 values would not be associated solely

with the return of ozone-depleting substances to their pre-1980 values (see schematic Figure 1c, d).

We address the behavior of ozone-depleting substances, the response of stratospheric ozone and UV radiation to

ozone-depleting substances and other factors to date, and the future evolution of ozone and UV radiation. We first present

the main findings, then the detailed supporting evidence, and finally the implications for policy formulation.

RECENT MAJOR FINDINGS AND CURRENT SCIENTIFIC UNDERSTANDING

Since the Scientific Assessment of Ozone Depletion: 2002, numerous laboratory investigations, atmospheric

observations, and theoretical and modeling studies have produced new key findings and have strengthened the overall

understanding of the ozone layer and its effect on 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, as well as

the coupling between climate change and stratospheric ozone depletion. 

Ozone-Depleting Substances

The previous Assessment noted that the tropospheric abundances of ozone-depleting gases were declining, whereas

the stratospheric abundances were at or near their peak. Now we ask: Are the tropospheric and stratospheric abundances

of ozone-depleting gases showing further changes that can be linked to policy actions, and how are they expected to

evolve in the future?

• The total combined abundances of anthropogenic ozone-depleting gases in the troposphere continue to

decline from the peak values reached in the 1992-1994 time period.

– The shorter-lived gases (e.g., methyl chloroform and methyl bromide) continue to provide much of the decline

in total combined effective abundances of anthropogenic chlorine-containing and bromine-containing ozonedepleting gases in the troposphere. The early removal of the shorter-lived gases means that later decreases in

ozone-depleting substances will likely be dominated by the atmospheric removal of the longer-lived gases.

– By 2005, the total combined abundance of anthropogenic ozone-depleting gases in the troposphere had

decreased by 8-9% from the peak value observed in the 1992-1994 time period. The overall magnitude of

this decrease is attributable to the estimated changes in emissions and is consistent with the known atmospheric lifetimes and our understanding of transport processes.

– The tropospheric abundances of hydrochlorofluorocarbons HCFC-22, -141b, and -142b increased much less

than projected in the previous Assessment. The implication is that emissions have been lower than anticipated. 

– The sum of bromine from halons and methyl bromide in the troposphere has decreased by 3-5% since 1998

because of decreases in methyl bromide abundance. Methyl bromide abundance has decreased by 14%

between 1997 and 2004. This decrease was larger than expected and suggests that when anthropogenic emissions of methyl bromide are reduced, its atmospheric abundance decreases more than previously thought. 
EXECUTIVE SUMMARY

xxi

• The combined stratospheric abundances of the ozone-depleting gases show a downward trend from their

peak values of the late 1990s, which is consistent with surface observations of these gases and a time lag for

transport to the stratosphere.

– Measurements of hydrogen chloride give a clear signature in the stratosphere of the decrease in stratospheric

ozone-depleting gases since reaching their peak in the late 1990s. 

– Stratospheric bromine has increased in line with its tropospheric trends in preceding years. A decrease in

stratospheric bromine has not yet been identified. Bromine continues to play a major role in stratospheric

ozone depletion.

• Our quantitative understanding of how halogenated very short-lived substances contribute to halogen levels

in the stratosphere has improved significantly since the 2002 Assessment, with brominated very short-lived

substances believed to make a significant contribution to total stratospheric bromine and its effect on stratospheric ozone.

– Brominated very short-lived substances (which are mainly of natural origin) contribute ~5 parts per trillion

(ppt) of bromine (estimates range from 3 to 8 ppt) to current levels of total stratospheric bromine (which are

about 18 to 25 ppt). These values are derived from observations of tropospheric bromine source gases and

bromine monoxide in the stratosphere, based on our current understanding of chemical transformations and

transport of very short-lived substances to the stratosphere. 

– Chlorinated very short-lived source gases, which are mainly of anthropogenic origin, have now been observed

at levels of about 50 ppt in the tropical upper troposphere and currently may be a small source of chlorine to

the stratosphere.

– Evidence suggests that iodine is currently much less important than bromine and chlorine for stratospheric

ozone destruction. 

Ozone, Ultraviolet Radiation, and Other Related Changes to Date

The previous Assessment noted that stratospheric ozone depletion was expected to begin to reverse within the next

decade or so. Now we ask: Are there any signs of this beginning? Have other factors influenced the behavior of ozone

during this period? Has surface UV radiation responded as anticipated?

• Our basic understanding that anthropogenic ozone-depleting substances have been the principal cause of the

ozone depletion over the past few decades has been strengthened. During the recent period of near-constant

abundances of ozone-depleting gases, variations in meteorology have been particularly important in influencing the behavior of ozone over much of the polar and extrapolar (60°S-60°N) regions. 

POLAR OZONE

• Springtime polar ozone depletion continues to be severe in cold stratospheric winters. Meteorological variability has played a larger role in the observed variability in ozone, over both poles, in the past few years. 

– Large Antarctic ozone holes continue to occur. The severity of Antarctic ozone depletion has not continued

to increase since the late 1990s and, since 2000, ozone levels have been higher than in some preceding

years. These recent changes, seen to different degrees in different diagnostic analyses, result from increased

dynamical activity and not from decreases in ozone-depleting substances.
EXECUTIVE SUMMARY

xxii

– The Antarctic ozone hole now is not as strongly influenced by moderate decreases in ozone-depleting gases,

and the unusually small ozone holes in some recent years (e.g., 2002 and 2004) are attributable to dynamical

changes in the Antarctic vortex. 

– The anomalous Antarctic ozone hole of 2002 was manifested in a smaller ozone-hole area and much smaller

ozone depletion than in the previous decade. This anomaly was due to an unusual major stratospheric sudden

warming.

– Arctic ozone depletion exhibits large year-to-year variability, driven by meteorological conditions. Over the

past four decades, these conditions became more conducive to severe ozone depletion because of increasingly widespread conditions for the formation of polar stratospheric clouds during the coldest Arctic winters.

This change is much larger than can be expected from the direct radiative effect of increasing greenhouse gas

concentrations. The reason for the change is unclear, and it could be because of long-term natural variability

or an unknown dynamical mechanism. The change in temperature conditions has contributed to large Arctic

ozone losses during some winters since the mid-1990s.

– The Arctic winter 2004/2005 was exceptionally cold, and chemical ozone loss was among the largest ever

diagnosed. The Arctic remains susceptible to large chemical ozone loss, and a lack of understanding of the

long-term changes in the occurrence of polar stratospheric clouds limits our ability to predict the future

evolution of Arctic ozone abundance and to detect the early signs of recovery.

– The role of chemical reactions of chlorine and bromine in the polar stratosphere is better quantified. Inclusion

of these advances results in improved agreement between theory and observation of the timing of both Arctic

and Antarctic polar ozone loss.

GLOBAL OZONE (60°S-60°N)

• The decline in abundances of extrapolar stratospheric ozone seen in the 1990s has not continued.

– Ozone abundances in the extrapolar regions, 60°S-60°N, have not further declined in recent years. The midlatitude column (i.e., overhead) ozone values for the 2002-2005 period were approximately 3% below pre1980 values in the Northern Hemisphere and approximately 6% in the Southern Hemisphere, essentially the

same as in the 1998-2001 period. 

– The contribution of polar ozone depletion to midlatitude ozone depletion is substantial; the extent of the contribution is estimated to be about one-third in the Northern Hemisphere and one-half in the Southern

Hemisphere. The larger contribution in the Southern Hemisphere is to be expected, given the larger polar

ozone depletion in the Antarctic relative to the Arctic region, and may explain the differences in magnitude

and seasonality of the long-term ozone changes in midlatitudes between the two hemispheres.

– Changes in tropospheric and stratospheric meteorology are partially responsible for the observed Northern

Hemisphere midlatitude winter ozone decline from 1979 to the mid-1990s and the ozone increase thereafter.

Quantification of these dynamical effects on long-term trends ranges from ~20% to up to 50%, depending on

the study. Most of these dynamically induced ozone changes appear to occur in the lowermost stratosphere. 

– Upper stratospheric ozone declined between 1979 and the mid-1990s, with the largest decrease of about 10-

15% near 40 km over midlatitudes, but has been relatively constant during the last decade. Lower stratospheric (20-25 km) ozone over midlatitudes also showed a decrease of up to 10% from 1979 to the mid-1990s

and has been relatively constant thereafter.
EXECUTIVE SUMMARY

xxiii

– Ozone abundances in the Northern Hemisphere midlatitude lowermost stratosphere (12-15 km) show a strong

decrease between 1979 and the mid-1990s, followed by an overall increase from 1996 to 2004, giving no net

long-term decrease at this level. This lowermost stratosphere ozone change contributed significantly to the

column ozone change during the last decade. Southern Hemisphere observations do not show a similar

increase in the 12-15 km altitude range since the mid-1990s.

– Total column ozone over the tropics (25°S-25°N) remains essentially unchanged. This finding is consistent

with the findings of the previous Assessments.

• Observations together with model studies suggest that the essentially unchanged column ozone abundances

averaged over 60°S-60°N over roughly the past decade are related to the near constancy of stratospheric

ozone-depleting gases during this period. 

– The slowing of the decline and leveling off of midlatitude upper stratospheric (35-45 km) ozone over the past

decade have very likely been dominated by the corresponding changes in stratospheric ozone-depleting gases. 

– Over the past decade, changes in stratospheric ozone-depleting gases have likely contributed to the slowing of

the midlatitude total column ozone decline and the general leveling off of ozone abundances. Transport has

also played an important role in ozone changes, particularly in the lowermost stratosphere, making attribution

of column ozone changes to changes in ozone-depleting gases more difficult than in the upper stratosphere. 

STRATOSPHERIC TEMPERATURE AND SURFACE UV RADIATION

• The stratospheric cooling observed during the past two decades has slowed in recent years.

– Satellite and radiosonde measurements reveal an overall decrease in the global-mean lower stratospheric temperature of approximately 0.5 K/decade over the 1979-2004 period, with a slowdown in the temperature

decline since the late 1990s. The overall temperature decrease is punctuated by transient warmings of the

stratosphere associated with the major volcanic eruptions in 1982 and 1991. Model calculations suggest that

the observed ozone loss is the predominant cause of the global-mean cooling observed over this period. 

– The lower stratospheric cooling is evident at all latitudes, in particular in both the Arctic and Antarctic

winter/spring lower stratosphere, but with considerable interannual variability in those regions. 

– Satellite observations show larger cooling trends in the upper stratosphere, with values of 1 to 2 K/decade,

but little additional decline since the middle 1990s. Model calculations suggest that the upper stratosphere

trends are due, about equally, to decreases in ozone and increases in well-mixed greenhouse gases.

– The long-term cooling rate of the Antarctic stratosphere has likewise reduced compared with what was seen

by the time of the previous Assessment. There has been large variability in Antarctic temperatures in recent

years. 

• Measurements from some stations in unpolluted locations indicate that UV irradiance (radiation levels) has

been decreasing since the late 1990s, in accordance with observed ozone increases. However, at some Northern

Hemisphere stations UV irradiance is still increasing, as a consequence of long-term changes in other factors

that also affect UV radiation. Outside polar regions, ozone depletion has been relatively small, hence, in many

places, increases in UV due to this depletion are difficult to separate from the increases caused by other factors,

such as changes in cloud and aerosol. In some unpolluted locations, especially in the Southern Hemisphere, UV

irradiance has been decreasing in recent years, as expected from the observed increases in ozone at those sites.

Model calculations incorporating only ozone projections show that cloud-free UV irradiance will continue to

decrease, although other UV-influencing factors are likely to change at the same time.
EXECUTIVE SUMMARY

xxiv

• In polar regions, high UV irradiances lasting for a few days have been observed in association with episodes

of low total ozone. Erythemal (i.e., sunburning) irradiance averaged over several days has been increased by

~70% over sub-Antarctic regions when ozone-depleted air from the polar region passes over. In Antarctica during

October, instantaneous enhancements that can exceed a factor of three have been observed. Over northern Europe

and Alaska, the observed enhancements were smaller.

The Future of the Ozone Layer and Surface UV Radiation

The previous (2002) Assessment noted that climate change would influence the future of the ozone layer. Now we

ask: How has our understanding of this issue progressed? How has our increased understanding led to an improved estimate of the state of the future ozone layer?

Previous Assessments have relied heavily on two-dimensional models for projections of future ozone, and these

models are used again here. The 2002 Assessment noted the emerging use of three-dimensional Chemistry-Climate

Models; we have now used these models extensively in this current 2006 Assessment.

• It is unlikely that total ozone averaged over the region 60°S-60°N will decrease significantly below the low

values of the 1990s, because the abundances of ozone-depleting substances have peaked and are in decline.

Two-dimensional models and three-dimensional Chemistry-Climate Models suggest that the minimum total column

ozone values in this region have already occurred. This conclusion is consistent with the observation that over this

region, ozone has not declined further in the 2002-2005 period. 

• The decrease in ozone-depleting substances is the dominant factor in the expected return of ozone levels to

pre-1980 values. Changes in climate will influence if, when, and to what extent ozone will return to pre-1980

values in different regions.

– The current best estimate is that global (60°S-60°N) ozone will return to pre-1980 levels around the middle of

the 21st century, at or before the time when stratospheric abundances of ozone-depleting gases return to

pre-1980 levels.

– Model simulations suggest that changes in climate, specifically the cooling of the stratosphere associated

with increases in the abundance of carbon dioxide, may hasten the return of global column ozone to pre-1980

values by up to 15 years, depending on the climate-gas scenario used. However, this would not be considered

a “recovery” of stratospheric ozone from ozone-depleting gases, because perceptible depletion due to anthropogenic ozone-depleting substances will still be contributing to the ozone levels.

• The Antarctic ozone hole is expected to continue for decades. Antarctic ozone abundances are projected to

return to pre-1980 levels around 2060-2075, roughly 10-25 years later than estimated in the 2002 Assessment.

The projection of this later return is primarily due to a better representation of the time evolution of ozonedepleting gases in the polar regions. In the next two decades, the Antarctic ozone hole is not expected to improve

significantly.

• Large ozone losses will likely continue to occur in cold Arctic winters during the next 15 years. The high variability in the volume of polar stratospheric clouds in these cold winters limits our ability to predict their future

occurrences. According to Chemistry-Climate Models, Arctic ozone levels are expected, on average, to return to

pre-1980 levels before 2050.

• Climate change will also influence surface UV radiation through changes induced mainly to clouds and the

ability of the Earth’s surface to reflect light. Aerosols and air pollutants are also expected to change in the future.

These factors may result in either increases or decreases of surface UV irradiance, through absorption or scattering. As ozone depletion becomes smaller, these factors are likely to dominate future UV radiation levels.
EXECUTIVE SUMMARY

xxv

ADDITIONAL SCIENTIFIC EVIDENCE AND RELATED INFORMATION

For the purpose of this report, ozone-depleting substances are considered as being either long-lived (more than 6

months in the atmosphere) or very short-lived (less than 6 months in the atmosphere), and these are considered in the next

two sections. The atmospheric lifetime determines the likelihood of halogens emitted in the troposphere reaching the

stratosphere, and so being able to deplete stratospheric ozone. The cumulative effective abundance of halogens in the

stratosphere is quantified as equivalent effective stratospheric chlorine (EESC).1

Long-Lived Ozone-Depleting Substances

ATMOSPHERIC TRENDS

• The decline in the abundance of methyl chloroform and methyl bromide, which are shorter lived than chlorofluorocarbons (CFCs), contributed the most to the decline in effective equivalent tropospheric chlorine

levels. By 2005, the abundances of the total combined anthropogenic ozone-depleting gases in the troposphere

had decreased by 8-9% from the peak value observed in the 1992-1994 time period. The total decline was ~120

parts per trillion (ppt) between 2000 and 2004, of which ~60 ppt was due to the decline of methyl chloroform and

~45 ppt due to methyl bromide. The CFCs together accounted for less than 23 ppt of the decline. The contribution of hydrochlorofluorocarbons (HCFCs) was to add 12 ppt.

• Total tropospheric chlorine-containing chemicals (approximately 3.44 parts per billion (ppb) in 2004) continued to decrease. Recent decreases (20 ppt/year or 0.59% in 2003-2004) have been at a slightly slower rate than

decreases in earlier years (23 ppt/year or 0.64% in 1999-2000) primarily because of the reduced contribution from

methyl chloroform. The declines in total chlorine (Cl) during 2000-2004 were slightly faster than projected for

these years in the Ab scenario of the previous (2002) ozone assessment report (baseline scenario following the

1999 Beijing Amendments). 

– Chlorofluorocarbons, consisting primarily of CFC-11, -12, and -113, accounted for 2.13 ppb (~62%) of

total Cl in 2004 and accounted for a decline of 9 ppt of Cl from 2003-2004 (or nearly half of the total Cl

decline in the troposphere over this period). Atmospheric mixing ratios of CFC-12, which account for about

one-third of the current atmospheric chlorine loading, have been constant within 1% (5 ppt) since 2000, and

some in situ and Northern Hemisphere column (i.e., overhead) measurements show that peak values were

attained in 2003. CFC-11 mixing ratios are decreasing at approximately 0.8%/year (1.9 ppt/year) and CFC-113

mixing ratios are decreasing at approximately 1%/year (0.8 ppt/year), which is twice as fast as in 1999-2000.

– Hydrochlorofluorocarbons (HCFCs), which are substitutes for CFCs, continue to increase in the atmosphere. HCFCs accounted for 214 ppt, or 6%, of total tropospheric chlorine in 2004 versus 180 ppt

(5%) of total Cl in 2000. HCFC-22 is the most abundant of the HCFCs and is currently (2000-2004)

increasing at a rate of 4.9 ppt/year (3.2%/year). HCFC-141b and HCFC-142b mixing ratios increased by 1.1

ppt/year (7.6%/year) and 0.6 ppt/year (4.5%/year) over this same period, or at about half the rates found for

these two gases in 1996-2000. The rates of increase for all three of these HCFC compounds are significantly

slower than projected in the 2002 Ozone Assessment (6.6, 2.6, and 1.6 ppt/year for HCFC-22, HCFC-141b,

and HCFC-142b, respectively). 

1 Equivalent effective stratospheric chlorine (EESC), a metric noted in the previous Assessment, has found widespread use. EESC is a

gauge of the overall stratospheric burden of ozone-depleting halogen. It is derived from ground-based measurements of ozonedepleting substances, with consideration given to the number of chlorine and bromine atoms in ozone-depleting substances, the rates

at which different ozone-depleting substances release their halogen once they reach the stratosphere, and the higher per-atom efficiency for bromine relative to chlorine in destroying ozone.
EXECUTIVE SUMMARY

xxvi

– Methyl chloroform has continued to decrease and contributed 13.5 ppt (or more than half) of the overall

decline observed for total tropospheric Cl in 2003-2004. It is currently still the largest contributor to the

decline in tropospheric chlorine.

• The stratospheric chlorine burden derived from ground-based total column and space-based measurements

of inorganic chlorine is now in decline. This is consistent with the decline in tropospheric chlorine from longlived halocarbons. The burden of total stratospheric chlorine derived from satellite measurements agrees, within

±0.3 ppb (about 12%), with the amounts expected from surface data when the delay due to transport is considered.

The uncertainty in this burden is large relative to the expected chlorine contribution from shorter-lived gases.

• Total organic tropospheric bromine from halons and methyl bromide peaked in about 1998 at 16.5 to 17 ppt

and has since declined by 0.6-0.9 ppt (3-5%). This observed decrease was solely a result of declines observed for

methyl bromide. Bromine from halons continues to increase, but at slower rates in recent years (0.1 ppt Br/year in

2003-2004). 

– Atmospheric amounts of methyl bromide declined beginning in 1999, when industrial production

was reduced. By mid-2004, mixing ratios had declined 1.3 ppt (14%) from the peak of 9.2 ppt measured

before 1999. Reported production of methyl bromide for emissive uses decreased by 50% during this

same period.

– Both the recently observed decline and the 20th century increase inferred for atmospheric methyl bromide were larger than expected. Although industrial emissions of methyl bromide were thought to account

for 20% (range 10-40%) of atmospheric methyl bromide during 1992-1998 (i.e., before production was

reduced), observed concentrations are consistent with this fraction having been 30% (range 20-40%). This

suggests that fumigation-related emissions could have a stronger influence on atmospheric methyl bromide

mixing ratios than estimated in past Assessments, though uncertainties in the variability of natural emission

rates and loss, and in the magnitude of methyl bromide banked in recent years, limit our understanding of this

sensitivity.

– Mixing ratios calculated from updated emission estimates are in good agreement with the measurements for halon-1211, but they exceed all the halon-1301 measurements since 1980 by more than 10%.

Atmospheric increases in halon-1211 (0.06 ppt/year) in 2000-2004 were about half those in 1996-2000. It is

currently unclear whether halon-1301 continues to increase. 

EMISSION ESTIMATES

• Global emissions of CFC-11 (88 Gg/yr, where 1 Gg = 109 grams), CFC-12 (114 Gg/yr), and CFC-113 (6 Gg/yr)

in 2003 were approximately 25%, 25%, and 3% of their maximum values around 1986. Emissions of CFC11, CFC-12, and CFC-113 have all continued to decrease since 2000. 

• Regional emission estimates for CFCs, methyl chloroform, and carbon tetrachloride have been reported for

the first time. Differences between developed and developing regions are indicative of the differing schedules of

phase-out. However, the patchiness in the present coverage (especially in developing regions such as Southeast

Asia) and the uncertainties in the regional estimates mean that useful comparisons between summed regional emissions and global emissions derived from trends cannot currently be made. Regional emission estimates of methyl

chloroform indicate that global emissions after 2000 may have been roughly 22 Gg/year, which is not statistically

different from the estimate of 12.9 Gg/year (for 2002) obtained from industry/United Nations Environment

Programme data.

• While emissions of HCFC-22 have remained nearly constant from 2000 to 2004, emissions of HCFC-141b

and HCFC-142b decreased by approximately 15% over the same period.
EXECUTIVE SUMMARY

xxvii

• Hydrofluorocarbon-23 (HFC-23) emissions estimated from atmospheric measurements have increased from

about 6 Gg/yr in 1990 to about 13 Gg/yr in 2001 (an increase of approximately 120%). These emissions are a

byproduct of HCFC-22 production. HFC-23 mixing ratios (18 ppt in 2004) have continued to increase, at approximately 0.7 ppt/yr (4%/yr), in 2001-2004.

Halogenated Very Short-Lived Substances

• Observations show that the total inorganic bromine (Bry) in the stratosphere is more abundant than can be

accounted for by the sum of bromine delivered to this region by “long-lived” brominated source gases. Very

short-lived substances (VSLS) are defined as trace gases whose local tropospheric lifetimes are comparable to or

shorter than tropospheric transport time scales, such that their tropospheric distributions are non-uniform. Here,

VSLS are considered to have atmospheric lifetimes of less than 6 months.

• Various lines of evidence show that brominated VSLS contribute about 5 ppt (with estimates ranging from 3

to 8 ppt) to total stratospheric Bry:

– The estimate of total stratospheric Bry abundance existing in the late 1990s is about 18-25 ppt. This value is

derived from different observations of bromine monoxide (BrO). This is greater than the 16-17 ppt of bromine

delivered to the stratosphere by “long-lived” brominated source gases (i.e., the halons and methyl bromide,

CH3Br) during this period.

– Measurements of organic brominated very short-lived source gases in the tropical upper troposphere amount

to about 3.5 ppt. Compounds resulting from source gas degradation and other sources of tropospheric inorganic bromine may contribute comparable amounts. 

• Levels of total stratospheric bromine continue to show evidence for a trend that is consistent with that of tropospheric total bromine. Further studies are required to determine if the recent decline in tropospheric bromine

will be reflected in stratospheric bromine abundance.

• It is unlikely that iodine is important for stratospheric ozone loss in the present-day atmosphere. There is

little evidence for measurable iodine in the form of iodine oxide (IO) in the lower stratosphere (i.e., about 0.1 ppt).

This difference in behavior compared with bromine may be partly attributed to the short photochemical lifetime

and lower abundances of iodine source gases, and aerosol uptake of iodine in the stratosphere.

• The sum of the chlorine content from very short-lived source gases in the tropical upper troposphere is now

estimated to be about 50 ppt. While 50 ppt of inorganic chlorine (Cly) from VSLS would represent only 1-2% of

Cly from long-lived source gases (~3500 ppt), it would represent a significant contribution to background Cly due

only to naturally occurring methyl chloride (550 ppt). The value of 100 ppt for Cly from short-lived species in the

previous Assessment also included a contribution of 50 ppt from phosgene.

• The majority of known brominated and iodinated very short-lived source gases are predominantly of natural origin and are nearly unchanged in the past few decades, as indicated by firn air studies. Some have small

anthropogenic sources and have increased slightly in the Northern Hemisphere (e.g., certain brominated trihalomethanes) and a few are almost exclusively anthropogenic (notably n-propyl bromide (n-PB)). 

• Chlorinated very short-lived source gases originate largely from anthropogenic emissions, although natural

sources also contribute. There is evidence for significant recent declines in concentrations of some chlorinated very

short-lived source gases, notably chloroform, dichloromethane, and tetrachloroethene.
EXECUTIVE SUMMARY

xxviii

• The route by which VSLS are most likely to reach the stratosphere is transport out of the boundary layer by

deep convection in tropical regions (particularly if co-located with regions of high emissions over tropical

oceans) followed by quasi-horizontal transport into the lower stratosphere. The fraction of source gas and

product gas transported into the stratosphere also depends on the local chemical lifetime, where source gases

decompose, and the loss of product gases by heterogeneous processes involving aerosols and clouds. 

Global Ozone (60°S-60°N)

TOTAL COLUMN OZONE

• Global mean total column (i.e., overhead) ozone values for 2002-2005 were 3.5% below 1964-1980 average

values. The 2002-2005 values are similar to the 1998-2001 values. This behavior is evident in all available global

datasets, although differences of up to 1% between annual averages exist between some individual sets.

• There are differences between the two hemispheres in the evolution of total column ozone at midlatitudes: 

– Averaged for the period 2002-2005, total column ozone for the Northern Hemisphere (NH) and Southern

Hemisphere (SH) midlatitudes (35°-60°) are about 3% and 5.5%, respectively, below their 1964-1980 average

values and are similar to their 1998-2001 values. The NH shows a minimum around 1993 followed by an

increase. The SH shows an ongoing decrease through the late 1990s followed by the recent leveling off.

– There are seasonal differences between the NH and the SH in ozone changes over midlatitudes. Changes

since the pre-1980 period over northern midlatitudes (35°N-60°N) are larger in spring, whereas those over

southern midlatitudes (35°S-60°S) are nearly the same throughout the year.

• Total column ozone over the tropics (25°S-25°N) remains essentially unchanged. These findings are consistent

with the findings of the previous Assessments.

VERTICAL OZONE DISTRIBUTION

• Upper stratospheric ozone declined during 1979-1995, but has been relatively constant during the last decade.

Measurements from Stratospheric Aerosol and Gas Experiment (SAGE I+II) and Solar Backscatter Ultraviolet

spectrometer (SBUV(/2)) satellite instruments show significant declines through 1995 when averaged over 60°N60°S and altitudes of 35 to 50 km. The net ozone decrease was ~10-15% over midlatitudes, with smaller, but

significant changes over the tropics. Available independent Umkehr, lidar, and microwave ozone measurements confirm these findings.

• Lower stratospheric ozone declined over the period 1979-1995, but has been relatively constant with significant variability over the last decade. At midlatitudes of both the Northern and Southern Hemispheres, measurements by the SAGE I+II and SBUV(/2) satellite instruments showed declines of up to 10% by 1995 between 20 and

25 km altitude. These decreases did not continue in the last decade. 

• In the lowermost stratosphere of the Northern Hemisphere, between 12 and 15 km, a strong decrease in

ozone was observed between 1979 and 1995, followed by an overall increase from 1996 to 2004 leading to

no net long-term decrease at this level. These changes in the lowermost stratosphere have a substantial influence on the column. The Southern Hemisphere midlatitude data do not show similar increases since 1995 at

these altitudes.
EXECUTIVE SUMMARY

xxix

UNDERSTANDING PAST CHANGES IN GLOBAL OZONE

• There is good overall agreement between observed long-term changes in extrapolar ozone and model simulations that include the effects of increasing halogens. The models generally reproduce the observed ozone changes

as a function of altitude, latitude, and season, confirming our understanding that halogen changes are the main

driver of global ozone changes. This link is supported by the statistical fit of globally averaged ozone observations

with equivalent effective stratospheric chlorine (EESC; see footnote at the beginning of this section of the Executive

Summary). However, there are discrepancies between the modeled and observed changes. Notably, the models

tend to perform less well in simulating changes in the SH compared with the NH.

• Empirical and model studies have shown that changes in tropospheric and stratospheric dynamics have been

partially responsible for the observed NH midlatitude winter ozone decline from 1979 to the mid-1990s and

the ozone increase thereafter. Whether this is due to dynamical variability or results from a long-term trend in

stratospheric circulation is not yet clear. Estimates of these dynamical effects on long-term trends range from

~20% up to 50% for the winter period.

• The inclusion of additional inorganic bromine (Bry) from very short-lived substances (VSLS) in models

yields simulations with larger ozone destruction at midlatitudes and polar regions, compared with studies

including only long-lived bromine source gases. In both regions, the enhanced ozone loss occurs in the lower

stratosphere via interactions of this bromine with anthropogenic chlorine. Midlatitude ozone loss is most enhanced

during periods of high aerosol loading. Ozone loss through chemical reactions involving bromine and odd-hydrogen

(HOx) is also enhanced at midlatitudes under all conditions. The impact on long-term midlatitude ozone trends

(1980-2004), assuming constant VSLS Bry, is calculated to be small because aerosol loading was low at the start

and end of this time period.

• Several independent modeling studies confirm that dilution of ozone-depleted polar air makes a substantial

contribution to midlatitude ozone depletion, especially in the Southern Hemisphere, where polar ozone loss

is much larger. Long-term annually averaged model-based estimates indicate that about one-third of the depletion

in the Northern Hemisphere (with large interannual variation), and about one-half in the Southern Hemisphere,

results from polar loss. This is supported by the observation that polar ozone depletion in the Antarctic spring is

strongly correlated with southern summer midlatitude ozone depletion. 

• Identification of the solar cycle signal in observed ozone has been improved because of the absence of major

volcanic eruptions over the past 15 years. The deduced solar cycle variation in column ozone has a mean amplitude of 2-3% (from minimum to maximum) in low to midlatitudes from the extended data series. 

Polar Ozone

ARCTIC

• Arctic spring total ozone values over the last decade are lower than values observed in the 1980s, but somewhat larger than those in the 1990s; they are highly variable depending on dynamical conditions. For current

halogen levels, anthropogenic chemical loss and variability in ozone transport are about equally important for yearto-year Arctic ozone variability. Lower-than-average Arctic vortex temperatures result in larger halogen-driven

chemical ozone losses. 

• For the “coldest” Arctic winters, the volume of air with temperatures low enough to support polar stratospheric clouds (called “VPSC”) has increased significantly since the late 1960s. This change is much larger than

expected from the direct radiative effect of increasing greenhouse gas concentrations. The reason for the change is

not clear, and it could be due to long-term natural variability or an unknown dynamical mechanism.
EXECUTIVE SUMMARY

• Column ozone chemical loss in the 2004/2005 Arctic winter was among the largest ever diagnosed. The

2004/2005 Arctic stratosphere was exceptionally cold, particularly at altitudes below 18 km, leading to a value of

VPSC 25% larger than the previous record value. Various independent studies and methods suggest that the chemical column ozone loss in 2004/2005 was among the largest ever observed. 

• The chemical loss of column ozone for Arctic winters exhibits a near-linear relation with VPSC during each

winter. A similar relation between ozone loss and VPSC is now seen for two independent analyses of chemical

ozone loss, increasing our confidence in the robustness of this relation. Observations of the large Arctic ozone loss

in 2004/2005 are in line with the relationship established for prior winters. 

ANTARCTIC

• Large Antarctic ozone holes continue to occur. The severity of Antarctic ozone depletion has not continued to

increase since the late 1990s and, since 2000, ozone levels have been higher in some years. These recent

changes, evident to different degrees in different diagnostic measures, result from increased dynamical wave

activity and not from decreases in ozone-depleting substances.

• In September 2002, the first-ever observed major stratospheric sudden warming occurred in the Southern

Hemisphere. This early spring warming caused a drastic reduction of the ozone hole area and resulted in

a less severe ozone hole. This warming resulted from anomalously strong dynamical wave activity in the

Southern Hemisphere. The Antarctic winter of 2004 was also dynamically very active and had higher ozone

levels than in previous years.

GENERAL

• Large interannual variability in polar stratospheric temperatures complicates the interpretation of temperature trends. Previously reported estimates of temperature trends in polar regions have differed from assessment to

assessment. In retrospect, it is evident that temperature trends derived over time scales of 1-2 decades, though they

may appear statistically significant, are not robust because of large interannual and decadal variability in observed

temperatures. Therefore, changes in reported temperature trends do not necessarily indicate systematic changes in

physical or chemical processes.

• Calculated chemical loss rates of polar ozone substantially increase in models that assume: 1) more efficient

ozone destruction by reactions involving two chlorine monoxide (ClO) molecules (referred to as the ClO

dimer cycle) and 2) higher concentrations of bromine. The higher concentrations of bromine are assumed

because of new evidence that natural bromocarbons play a considerable role in the stratospheric bromine budget.

More efficient ozone destruction by the ClO dimer cycle is supported by analyses of the first stratospheric observations of the ClO dimer. These two changes to models improve the comparison between measured and modeled

Arctic ozone loss rates that was noted in the past Assessment. These changes also improve comparison between

theory and observation of the timing of ozone loss over Antarctica.

• For the first time, measurements show unambiguously that nitric acid trihydrate (NAT) polar stratospheric

cloud (PSC) particles can nucleate above the ice frost point, and there is additional evidence of their widespread occurrence. Widespread, low-number-density NAT clouds can lead to denitrification and enhanced ozone

loss. Incorporating NAT nucleation above the ice frost point into chemical transport models has improved denitrification simulations, but discrepancies in interannual variability remain, probably because the NAT nucleation

mechanisms are not fully understood.

xxx
EXECUTIVE SUMMARY

xxxi

Impact of Climate Change

• The stratospheric cooling observed during the past two decades has slowed in recent years.

– Satellite and radiosonde measurements reveal an overall decrease in the global-mean lower stratospheric

temperature of approximately 0.5 K/decade over the 1979-2004 period, with a slowdown in the temperature

decline since the late 1990s. The overall temperature decrease is punctuated by transient warmings of the

stratosphere associated with the major volcanic eruptions in 1982 and 1991. Consistent with the previous

Assessment, model calculations indicate that the observed ozone loss is the predominant cause of the

global-mean lower stratospheric cooling observed over this period. 

– The cooling of the lower stratosphere is evident at all latitudes. The annual-mean temperature in the Arctic

lower stratosphere has decreased by approximately 1 K over the 1979-2004 period. There continues to be a

large interannual variability in the winter and springtime, which makes it difficult to determine reliable trends

in the Arctic lower stratosphere. 

– A considerable interannual variability in the temperature of the Antarctic lower stratosphere has become evident in recent years; this has reduced the estimate of the decadal-scale cooling trend in that region from that

in the previous Assessment. 

– Satellite observations show large global-mean cooling trends in the upper stratosphere (1 to 2 K/decade over

1979-2004), but there is little additional decline since the middle 1990s. The conclusion of the previous

Assessment, that the observed upper stratosphere trends are due to approximately similar contributions from

decreases in ozone and increases in well-mixed greenhouse gases, is reaffirmed.

• Future increases of greenhouse gas concentrations will contribute to the average cooling in the stratosphere.

Estimates derived from climate models (AOGCMs, coupled ocean-atmosphere general circulation models) and

Chemistry-Climate Models (CCMs) with interactive ozone consistently predict continued cooling of the global

average stratosphere. The predicted cooling rate within the next two decades is dependent on the prescribed scenario and the type of model used for the Assessment. At 50 hPa (an altitude of about 20 km), an average of all

AOGCMs gives approximately 0.1 K/decade, while CCMs predict a larger cooling of about 0.25 K/decade caused

by the interactive consideration of ozone changes. All models calculate a stronger cooling at 10 hPa (about 30 km),

averaging approximately 0.5 K/decade. Simulations of polar temperatures in the future are less certain than global

mean temperatures because of greater interannual variability.

• Chemical reaction rates in the atmosphere are dependent on temperature, and thus the concentration of

ozone is sensitive to temperature changes. Decreases in upper stratospheric temperature slow the rate of photochemical ozone destruction in this region. Hence the concentration of upper stratospheric ozone increases in

response to cooling. Cooling of the polar lower stratosphere would lead to more efficient chlorine activation on

aerosols and polar stratospheric clouds, and enhanced ozone destruction. Therefore, the concentration of ozone in

the springtime polar lower stratosphere would decrease in response to cooling.

INTERACTIONS BETWEEN STRATOSPHERE AND TROPOSPHERE

• Changes to the temperature and circulation of the stratosphere affect climate and weather in the troposphere. Consistent evidence for this coupling comes from both observational analyses and models, but the mechanisms responsible are not well understood. The dominant tropospheric response comprises changes in the strength

of the midlatitude westerlies (the prevailing winds from the west). Observations and models suggest that Antarctic

ozone depletion, through its effects on the lower stratospheric vortex, has contributed to the observed strengthening

of tropospheric winds and Antarctic surface cooling during December-February.
EXECUTIVE SUMMARY

xxxii

• Updated datasets of stratospheric water vapor concentrations now show differences in long-term behavior.

Recent trend analyses, which are based on only two available multiyear datasets, cast doubt on the positive stratospheric water vapor trend that was noted in the previous Assessment. Balloonborne water vapor measurements at

Boulder, Colorado, for the period 1980-2005 show a significant increase of 5-10% per decade over altitudes of 15-

28 km. Global water vapor measurements from the Halogen Occultation Experiment (HALOE) satellite instrument

for 1991-2005 do not show a corresponding positive lower stratospheric trend. Interannual water vapor changes

derived from HALOE data exhibit quantitative agreement with temperature variations near the tropical tropopause.

In contrast, the long-term increases inferred from the Boulder data are larger than can be explained by observed

tropopause temperature changes or past increases in tropospheric methane.

The Ozone Layer in the 21st Century

• The slowing of the decline and leveling off of midlatitude upper stratospheric (35-45 km) ozone over the past

10 years has very likely been dominated by changes in equivalent effective stratospheric chlorine (EESC).

Gas-phase chemistry, modulated by changes in temperature and other gases such as methane, directly controls

ozone in this region, and observed ozone increases are similar to those modeled from EESC decreases.

• Over the past 10 years, changes in EESC have likely contributed to the slowing of midlatitude total column

ozone decline and the leveling off of column ozone. Transport also has played an important role, particularly in

the lowermost stratosphere, making attribution of ozone changes to solely EESC difficult. For northern midlatitudes, increases in ozone have been greater than expected from EESC decreases alone; whereas over southern

midlatitudes, ozone changes are more consistent with the range expected from EESC decreases.

• The decline in EESC has not caused the large annual variations observed in Arctic ozone depletion. Indeed,

the onset of ozone recovery in the Arctic has not been detected. The large interannual variations driven by

changes in meteorology are likely to preclude the detection of the first stage of recovery.

MODEL PROJECTIONS

Two-dimensional models and three-dimensional coupled Chemistry-Climate Models (CCMs), both of which incorporate many of the factors that affect ozone as well as their feedbacks, have been used in this Assessment to make projections of the evolution of ozone throughout the 21st century. The model projections guiding the conclusions below were

based on the surface time series of halocarbons from the “Ab” scenario (baseline scenario following the Beijing

Amendments) from the 2002 ozone assessment, as well as the well-mixed greenhouse gas concentrations from the Third

Assessment Report of the Intergovernmental Panel on Climate Change, Climate Change 2001: The Scientific Basis. To

focus on long-term changes, the projected ozone anomalies have been decadally smoothed. 

• The CCMs used to project future ozone abundances have been critically evaluated and more emphasis has

been given to those models that best represent the processes known to strongly affect column ozone abundances. The CCMs vary in their skill in representing different processes and characteristics of the atmosphere.

However, there is sufficient agreement between the majority of the CCMs and the observations, so that some confidence can be placed in their projections.

• Averaged over 60°S to 60°N, total column ozone is projected to increase by 1% to 2.5% between 2000 and

2020, reaching 1980 values at the same time or before EESC in that region declines to 1980 values (2040-

2050). By 2100, ozone should be up to 5% above 1980 values. Both two-dimensional models and CCMs suggest

that minimum total column ozone values have already occurred in this region. Ozone exceeds pre-1980 values at

the time EESC returns to 1980 values in nearly all models that include coupling between well-mixed greenhouse

gases and temperature (interactive two-dimensional models and the CCMs). Similarly the elevated ozone levels in

2100 are observed in all interactive two-dimensional models and in the one CCM that extended to 2100. 
EXECUTIVE SUMMARY

xxxiii

• Antarctic springtime ozone is projected to increase by 5% to 10% between 2000 and 2020, reaching 1980

values close to when Antarctic EESC decreases to 1980 values (2060-2075), and changing little thereafter.

Different diagnostic indicators of ozone depletion show different sensitivities to EESC. The most rapid change

(decrease) occurs in the ozone mass deficit and the slowest change (increase) occurs in ozone minimum values and

October ozone anomalies. Minimum ozone values remain roughly constant between 2000 and 2010 in many

models. The projected onset of decreases in the ozone mass deficit occurs between 2000 and 2005, whereas the

projected onset of increases in minimum Antarctic ozone does not occur until after 2010 in many models.

• Arctic springtime ozone is projected to increase by 0% to 10% between 2000 and 2020, reaching 1980 values

much earlier than when Arctic EESC decreases to 1980 values (2060-2070). By 2100, Arctic ozone is projected to be substantially above 1980 values. While the large interannual variability in projected Arctic ozone

obscures the date when the ozone turnaround due to decreasing EESC occurs, this is projected to occur before

2020. The increase in Arctic ozone does not follow EESC as closely as in the Antarctic, and in the majority of

CCMs, Arctic ozone exceeds 1980 values before the Antarctic. There is no indication of future large decreases in

Arctic column ozone in any of the model simulations. There is large uncertainty in predictions of Arctic ozone

because of the smaller ozone depletion and the larger interannual variability in the Arctic stratosphere.

• The projected ozone in 2100 is sensitive to future levels of the well-mixed greenhouse gases. For example,

expected future increases in nitrous oxide (N2O) will increase stratospheric nitrogen oxides (NOx), which may

exacerbate ozone depletion. However, the expected stratospheric cooling induced by increasing concentrations of

greenhouse gases, primarily carbon dioxide (CO2), is expected to slow gas-phase ozone depletion reactions and,

thereby, increase ozone. The net effect on ozone amounts will depend on future levels of the different well-mixed

greenhouse gases. The importance of this temperature feedback is demonstrated by the non-interactive twodimensional models, which predict that extrapolar column ozone will be less than or near 1980 values through the

latter half of the century.

• Human activities are expected to affect stratospheric ozone through changes in tropospheric emissions of

trace gases. Enhanced methane (CH4) emission (from wetter and warmer soils) is expected to enhance ozone production in the lower stratosphere, whereas an increase in nitrous oxide emission (from extended use of artificial

fertilizer) is expected to reduce ozone in the middle and high stratosphere. Also, changes in non-methane

hydrocarbons and NOx emissions are expected to affect the tropospheric concentrations of hydroxyl radical

(OH) and, hence, impact the lifetimes and concentrations of stratospheric trace gases such as CH4 and organic

halogen species.

• Future changes of stratospheric water vapor concentrations are uncertain. If water vapor concentrations

increase in the future, there will be both a radiative and a chemical effect. Modeling studies suggest increased

water vapor concentrations will enhance odd hydrogen (HOx) in the stratosphere and subsequently influence

ozone depletion. Increases in water vapor in the polar regions would raise the temperature threshold for the formation of polar stratospheric clouds, potentially increasing springtime ozone depletion.

UV Radiation and Its Changes

• Some unpolluted sites show that UV radiation has been decreasing since the late 1990s. However, at some

midlatitude stations in the Northern Hemisphere, surface UV irradiance continued to increase at rates of a

few percent per decade. The observed increases and their significance depend on location, wavelength range, and

the period of measurements. These increases cannot be explained solely by ozone depletion and could be attributed

to a decreasing tendency in aerosol optical extinction and air pollution since the beginning of the 1990s and partly

to decreasing cloudiness, as estimated from satellites.
EXECUTIVE SUMMARY

xxxiv

• Tropospheric aerosols are responsible for the overestimation of UV irradiance from satellite instruments

(e.g., TOMS, the Total Ozone Mapping Spectrometer) that use solar backscattered ultraviolet radiation to

derive surface UV irradiance. At clean sites the agreement with ground-based measurements is good. However,

over more polluted locations the bias can be as large as 40%, because the lowermost atmosphere containing the

absorbing aerosols is not adequately probed by the measurements. The presence of clouds and snow or ice cover

can also lead to significant biases. New algorithms have been developed to improve the parameterization of aerosol,

snow, and ice effects on satellite-derived surface UV irradiance, as well as of cloud effects using Advanced Very

High Resolution Radiometer (AVHRR) and Meteosat images, showing on average good agreement with groundbased UV observations. Although the TOMS instrument is no longer available, the continuity of satellite-derived

global UV data is maintained with the new Ozone Monitoring Instrument (OMI) onboard the Aura satellite.

• Clouds are the major factor limiting the detectability of long-term changes in UV radiation due to ozone

changes or other factors. Even if ozone trends were linear, at least 10-15 years of measurements would be needed

for detecting a trend in UV radiation. At most extrapolar locations, the effects on UV-B radiation from ozone

depletion are relatively small, and the influence of clouds is dominant.

• Air pollutants may counterbalance the UV radiation increases resulting from ozone depletion. Observations

confirm that UV-absorbing air pollutants in the lower troposphere, such as ozone, nitrogen dioxide (NO2) and sulfur

dioxide (SO2), attenuate surface UV irradiance by up to ~20%. This effect is observed at locations near the emission sources. Air pollution exerts stronger attenuation in UV compared with attenuation in total solar irradiance.

Ozone Depletion Potentials and Global Warming Potentials

• The effectiveness of bromine compared with chlorine for global ozone depletion (on a per-atom basis), typically referred to as α, has been re-evaluated upward from 45 to a value of 60. The calculated values from three

independent two-dimensional models range between 57 and 73, depending on the model used and depending on the

assumed amount of additional bromine added to the stratosphere by very short-lived substances. 

• Ozone Depletion Potentials (ODPs) have been re-evaluated, with the most significant change being a 33%

increase for bromocarbons due to the update in the estimate for the value of α. A calculation error in the previous ozone assessment leading to a 13% overestimate for the ODP of halon-1211 has been corrected.

• Direct and indirect Global Warming Potentials (GWPs) have been updated. The direct GWPs were revised

for changes in radiative efficiencies of HFC-134a, carbon tetrafluoride (CF4), HFC-23, HFC-32, HFC-227ea,

and nitrogen trifluoride (NF3), as well as changes in the lifetimes of trifluoromethylsulfurpentafluoride

(SF5CF3) and methyl chloride (CH3Cl). In addition, the direct GWPs for all compounds are affected by slight

decreases in the CO2 absolute Global Warming Potentials for various time horizons. Indirect GWPs have been

updated primarily to reflect the later return of ozone-depleting substances to 1980 levels estimated in this Assessment

compared with the earlier Assessment and to account for the increased bromine efficiency factor.

IMPLICATIONS FOR POLICY FORMULATION

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

been conveyed to decisionmakers through a series of international state-of-understanding assessment reports. The

research findings in the Scientific Assessment of Ozone Depletion: 2006 that are given above and are summarized here

provide direct current scientific input to governmental, industrial, and other policy decisions associated with protection of

the ozone layer:
EXECUTIVE SUMMARY

xxxv

• Our basic understanding that anthropogenic ozone-depleting substances have been the principal cause of the

ozone depletion over the past decades has been strengthened.

• The Montreal Protocol is working: There is clear evidence of a decrease in the atmospheric burden of ozonedepleting substances and some early signs of stratospheric ozone recovery. As noted earlier in this Executive

Summary:

– Abundances of ozone-depleting substances, taken together, are clearly decreasing in the lower atmosphere.

Nearly all the abundances of individual ozone-depleting substances are also decreasing.

– There are clear indications that the equivalent effective stratospheric chlorine (which is a standardized measure

of ozone-depleting capacity of bromine and chlorine in the stratosphere; see footnote in the previous section

of this Executive Summary) has also started to decrease.

– In the last decade, the depletion of the global ozone layer has not worsened.

– Measurements from some unpolluted locations show that UV radiation has decreased since the late 1990s, in

accordance with observed ozone increases at those sites.

• The unusually small 2002 Antarctic ozone hole was characterized by a smaller area and higher ozone levels

than observed during the past decade. This was due to an unusually strong meteorological event (a “major

stratospheric sudden warming”) and was not due to changes in ozone-depleting gases. The 2003 and 2005 ozone

holes exhibited severe depletions as observed since the early 1990s. Severe Antarctic ozone losses will very likely

continue to be observed for at least the next 10-20 years because of the expected slow decline of the long-lived

ozone-depleting gases.

• Bromine is now estimated to be approximately 60 times as effective as chlorine in global ozone depletion, on

a per-atom basis. This value is larger than the effectiveness of 45 used in the 2002 Assessment. This increase in

the effectiveness of bromine increases the Ozone Depletion Potential of bromine-containing compounds, evaluated

using the semi-empirical method that has been used in previous Assessments.

• Long-term recovery of the ozone layer from the effects of ozone-depleting substances is expected to span

much of the 21st century and is estimated to occur later than projected in the previous Assessment (2002). An

important milestone in the recovery process is the time when combined chlorine and bromine amounts (equivalent

effective stratospheric chlorine, EESC) decline to pre-1980 values.

– The date when equivalent effective stratospheric chlorine at midlatitudes returns to pre-1980 levels is now

calculated to be 2049, for the case of global compliance with the Montreal Protocol with no significant exceptions (Scenario A1). This date is about 5 years later than projected in the previous (2002) Assessment. This

projected later date primarily results from (i) an increase in CFC-11 and CFC-12 emissions due to the larger

recent estimates of amounts currently contained in equipment and products (banks) and (ii) an increase in

HCFC-22 emissions due to larger estimated future production, as reported in the 2005 Intergovernmental

Panel on Climate Change/Technology and Economic Assessment Panel (IPCC/TEAP) Special Report on

Safeguarding the Ozone Layer and the Global Climate System: Issues Related to Hydrofluorocarbons and

Perfluorocarbons.

– The return to pre-1980 conditions of equivalent effective stratospheric chlorine for the Antarctic vortex is

projected to occur around 2065, more than 15 years later than the return of midlatitude equivalent effective

stratospheric chlorine to pre-1980 levels. This projected later recovery is because, unlike in previous

Assessments, we now recognize that the age of air is greater in the Antarctic lower stratosphere, which affects

the amount of ozone-depleting gases available for ozone depletion. The equivalent effective stratospheric

chlorine evaluation for the Antarctic vortex has not been presented in previous Ozone Assessments.
EXECUTIVE SUMMARY

xxxvi

• Potential options for accelerating the recovery of the ozone layer have been evaluated. The table provided

below shows hypothetical estimates of the upper limits of improvements that could be achieved if global anthropogenic production of ozone-depleting substances were to stop after 2006, if emissions were eliminated from

existing banks at the end of 2006, or if global anthropogenic emissions of ozone-depleting substances were to stop

after 2006. Some options show greater effectiveness for accelerating recovery of the ozone layer than in previous

Assessments, for reasons noted below.

– The percentage cumulative (integrated) reductions in equivalent effective stratospheric chlorine for halons

and CFCs, integrated from 2007 until the 1980 level is reattained (shown in Column B), are larger than previously reported. This is because recent bank estimates from the 2005 Intergovernmental Panel on Climate

Change/Technology and Economic Assessment Panel (IPCC/TEAP) Special Report on Safeguarding the

Ozone Layer and the Global Climate System: Issues Related to Hydrofluorocarbons and Perfluorocarbons

are significantly larger and likely more reliable than values presented in previous Assessments for CFC-11,

CFC-12, and halon-1211.

– The percentage reduction in integrated equivalent effective stratospheric chlorine for HCFCs shown in

Column A is larger than previously reported. This is because of significantly larger estimates of future

HCFC-22 production in Article-5 countries.

– The percentage reduction in integrated equivalent effective stratospheric chlorine for methyl bromide in

Column A is larger than previously reported. This is because of the upward revision of the fraction of anthropogenic emissions relative to total methyl bromide emissions, as well as being due to the upward revision in

the ozone-depletion effectiveness of bromine atoms compared with chlorine atoms mentioned earlier.

– In addition to the percentage reductions in integrated equivalent effective stratospheric chlorine, these

scenarios can reduce the time for equivalent effective stratospheric chlorine to drop below 1980 values. A

hypothetical elimination of all emissions of ozone-depleting substances after 2006 (Column C) would advance

this time by about 15 years, from 2049 to 2034. A hypothetical elimination of all emissions from production

of ozone-depleting substances after 2006 (Column A) would advance it by about 6 years, to 2043.

Table 1. Percentage reductions in integrated equivalent effective stratospheric chlorine relative to the

baseline (A1) scenario that can be achieved in the hypothetical cases. 1

Compound or Compound Group Column A: Column B: Column C:

All Emissions All Emissions All Emissions

Eliminated from Eliminated from Eliminated after

Production after 2006 Existing Banks at 2006

End of 2006

Chlorofluorocarbons (CFCs) 0.3 11 11

Halons 0.5 14 14

Carbon tetrachloride (CCl4) 3 (a) 3

Methyl chloroform (CH3CCl3) 0.2 (a) 0.2

Hydrochlorofluorocarbons (HCFCs) 12 4 16

Methyl bromide (CH3Br) (anthropogenic) 5 (a) 5

1 Column A corresponds to a hypothetical elimination of all emissions from production after 2006. Column B corresponds to a hypothetical elimination of all emissions from banks existing at the end of 2006 (for example, capture and destruction). Column C corresponds to a hypothetical elimination of all emissions after 2006 and is approximately equal to the sum of columns A and B. 

(a) For these compounds, banks are uncertain and therefore emissions are equated to production in these calculations.
EXECUTIVE SUMMARY

xxxvii

– A methyl bromide phase-out has been in effect since 2005 in developed countries, with critical-use exemptions granted in 2005 and 2006 at levels that are 30-40% of the 2003-2004 production levels. Two additional

hypothetical cases of critical and quarantine and pre-shipment (QPS) exempted uses of methyl bromide were

considered. In the analysis of both cases, equivalent effective stratospheric chlorine is integrated above the

1980 level from 2007 until it returns to the 1980 level. The size of the critical use exemptions is similar to the

estimated use of methyl bromide for QPS use.

> If critical-use methyl bromide exemptions continue indefinitely at the 2006 level compared to a cessation of these exemptions in 2010 or 2015, midlatitude integrated equivalent effective stratospheric

chlorine would increase by 4.7% or 4.0%, respectively.

> If production of methyl bromide for QPS use were to continue at present levels and cease in 2015, midlatitude integrated equivalent effective stratospheric chlorine would decrease by 3.2% compared with

the case of continued production at present levels.

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

layer. Emissions associated with continued or expanded exemptions, QPS, process agents, and feedstocks may

also delay recovery. 

• The role of very short-lived halogenated substances in stratospheric ozone depletion is now believed to be of

greater importance than previously assessed. This suggests that significant anthropogenic production of

such substances could enhance ozone depletion. Current understanding of the Ozone Depletion Potentials of

these classes of substances is:

– The Ozone Depletion Potentials of n-propyl bromide are 0.1 for tropical emissions and 0.02-0.03 for emissions restricted to northern midlatitudes. These are unchanged from the previous Assessment.

– New analyses suggest upper-limit Ozone Depletion Potentials for CF3I of 0.018 for tropical emissions and

0.011 for midlatitude emissions. The previous Assessment had an upper limit of 0.008.

– Any chlorinated very short-lived source gas with a lifetime of ~25 days, one chlorine atom, and a similar

molecular weight to CFC-11, has an Ozone Depletion Potential of about 0.003.

• Understanding the interconnections between ozone depletion and climate change is crucial for projections of

future ozone abundances. The ozone-depleting substances and many of their substitutes are also greenhouse

gases; changes in ozone affect climate; and changes in climate affect ozone. These issues were recently the subject

of the IPCC/TEAP 2005 Special Report on Safeguarding the Ozone Layer and the Global Climate System: Issues

Related to Hydrofluorocarbons and Perfluorocarbons, and some aspects of the coupling between ozone depletion

and climate change have been considered in this 2006 Ozone Assessment. An important development has been the

emerging use of three-dimensional models that incorporate the interaction between chemistry and climate.
CHAPTER 1

Long-Lived Compounds

Contributors:

P. Bernath

D. Blake

T. Blumenstock

J. Butler

P.-F. Coheur

G. Dutton

D. Etheridge

L. Froidevaux

P. Krummel

J. Mühle

S. O’Doherty

D. Oram

E. Remsberg

R. Rhew

J.M. Russell III

C. Trudinger

D. Waugh

R. Zander

Lead Authors:

C. Clerbaux

D.M. Cunnold

Coauthors:

J. Anderson

A. Engel

P.J. Fraser

E. Mahieu

A. Manning

J. Miller

S.A. Montzka

R. Nassar

R. Prinn

S. Reimann

C.P. Rinsland

P. Simmonds

D. Verdonik

R. Weiss

D. Wuebbles

Y. Yokouchi

Final Release: February 2007

From Scientific Assessment of Ozone Depletion: 2006

CHAPTER 1

LONG-LIVED COMPOUNDS

Contents

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

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

1.2 HALOGENATED OZONE-DEPLETING GASES IN THE ATMOSPHERE . . . . . . . . . . . . . . . . . . . . . . . . . . 1.7

1.2.1 Updated Atmospheric Observations of Ozone-Depleting Gases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.7

1.2.1.1 Chlorofluorocarbons (CFCs) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.7

1.2.1.2 Hydrochlorofluorocarbons (HCFCs) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.10

1.2.1.3 Halons . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.12

1.2.1.4 Carbon Tetrachloride (CCl4) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.14

1.2.1.5 Methyl Chloroform (CH3CCl3) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.14

1.2.1.6 Methyl Chloride (CH3Cl) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.15

1.2.1.7 Methyl Bromide (CH3Br) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.18

1.2.2 Lifetimes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.22

1.2.2.1 Halocarbon Lifetime Updates and Variability . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.22

1.2.2.2 Uncertainties in Ozone-Destroying Halocarbon Lifetimes . . . . . . . . . . . . . . . . . . . . . . . . . 1.22

1.2.3 Total Atmospheric Chlorine and Bromine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.27

1.2.3.1 Total Organic Chlorine in the Troposphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.27

1.2.3.2 Total Chlorine in the Stratosphere and Hydrogen Chloride (HCl) . . . . . . . . . . . . . . . . . . . . 1.27

1.2.3.3 Organic Bromine in the Troposphere from Methyl Bromide and Halons . . . . . . . . . . . . . . 1.32

1.2.4 Effective Equivalent Chlorine and Equivalent Effective Stratospheric Chlorine . . . . . . . . . . . . . . . . 1.33

1.2.5 Fluorine in the Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.35

1.2.5.1 Total Stratospheric Fluorine and Hydrogen Fluoride (HF) at 55 km . . . . . . . . . . . . . . . . . . 1.35

1.2.5.2 Total and Stratospheric Column Inorganic Fluorine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.36

1.3 EMISSIONS ESTIMATED FROM THE HALOCARBON MEASUREMENTS . . . . . . . . . . . . . . . . . . . . . . 1.38

1.3.1 Regional Emissions and Inverse Modeling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.38

1.3.2 Global Emissions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.41

1.3.2.1 CFCs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.41

1.3.2.2 HCFCs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.44

1.3.2.3 Halons . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.44

1.3.2.4 Carbon Tetrachloride . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.45

1.3.2.5 Methyl Chloroform . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.45

1.4 OTHER TRACE GASES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.46

1.4.1 Carbon Dioxide (CO2) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.46

1.4.2 Methane (CH4), Including Isotopic Data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.47

1.4.3 Nitrous Oxide (N2O) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.47

1.4.4 Hydrofluorocarbons (HFCs) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.48

1.4.4.1 HFC-134a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.48

1.4.4.2 HFC-23 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.48

1.4.4.3 HFC-143a, HFC-125, and HFC-152a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.49

1.4.5 Perfluorocarbons, Sulfur Hexafluoride (SF6), and SF5CF3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.49

1.4.5.1 CF4 and CF3CF3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.49

1.4.5.2 CF3CF2CF3 and c-C4F8 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.50

1.4.5.3 SF6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ..1.51

1.4.5.4 SF5CF3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.51

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.52

LONG-LIVED COMPOUNDS

1.1

SCIENTIFIC SUMMARY

• The fully amended and revised Montreal Protocol is continuing to be very successful in reducing the emissions and atmospheric abundances of most ozone-destroying gases. By 2005, the total combined abundance

of anthropogenic ozone-depleting gases in the troposphere had decreased by 8-9% from the peak value

observed in the 1992-1994 time period.

Tropospheric Chlorine

• Total tropospheric chlorine-containing chemicals (approximately 3.44 parts per billion (ppb) in 2004) continued to decrease. Recent decreases (−20 parts per trillion per year (ppt/yr) or −0.59% in 2003-2004) have been

at a slightly slower rate than in earlier years (−23 ppt/yr or −0.64% in 1999-2000) primarily because of the reduced

contribution from methyl chloroform. The declines in total chlorine (Cl) during 2000-2004 were slightly faster than

projected for these years in the Ab (most likely, or baseline) scenario of the previous (2002) Assessment (WMO,

2003).

• Chlorofluorocarbons (CFCs), consisting primarily of CFC-11, -12, and -113, accounted for 2.13 ppb (~62%)

of total Cl in 2004 and accounted for a decline of 9 ppt Cl from 2003-2004 (or nearly half of the total Cl

decline in the troposphere over this period). Atmospheric mixing ratios of CFC-12, which account for about

one-third of the current atmospheric chlorine loading, have been constant within 1% (5 ppt) since 2000, and some in

situ and Northern Hemisphere column measurements show that peak values were attained in 2003. CFC-11 mixing

ratios are decreasing at approximately 0.8%/yr (1.9 ppt/yr) and CFC-113 mixing ratios are decreasing by approximately 1%/yr (0.8 ppt/yr), which is twice as fast as in 1999-2000.

• Hydrochlorofluorocarbons (HCFCs), which are substitutes for CFCs, continue to increase in the atmosphere.

HCFCs accounted for 214 ppt or 6% of total tropospheric chlorine in 2004 versus 180 ppt of Cl (5% of total

Cl) in 2000. HCFC-22 is the most abundant of the HCFCs and is currently (2000-2004) increasing at a rate of 4.9

ppt/yr (3.2%/yr). HCFC-141b and HCFC-142b mixing ratios increased by 1.1 ppt/yr (7.6%/yr) and 0.6 ppt/yr

(4.5%/yr) over this same period or at about half the rates found for these two gases in 1996-2000. The rates of

increase for all three of these HCFC compounds are significantly slower than projected in the previous Assessment

(6.6, 2.6, and 1.6 ppt/yr for HCFC-22, HCFC-141b, and HCFC-142b, respectively) (WMO, 2003). HCFC-123 and

-124 are currently measured, but contribute less than 1% to total chlorine from HCFCs.

• Methyl chloroform (CH3CCl3) has continued to decrease and contributed 13.5 ppt (or more than half) of the

overall decline observed for total tropospheric Cl in 2003-2004. It is currently still the largest contributor to

the decline in tropospheric chlorine. Globally averaged surface mixing ratios were 22.6 ppt in 2004 versus 46.4

ppt in 2000.

• An imbalance exists between the emissions inferred from the atmospheric measurements and the emissions

needed to account for the stratospheric and ocean sinks and the more recently discovered soil sink of CCl4.

CCl4 continues to decline at an approximately steady rate of 1 ppt/yr (about 1%/yr).

• The globally averaged abundance of methyl chloride of 550 ± 30 ppt is the same as that given in the previous

Assessment, and there has been no consistent trend in the atmospheric values in recent years. Mixing ratios at

midlatitudes of the two hemispheres are 5-10% lower than the high values measured in 1997-1998 (which were

probably related to large-scale effects of forest fires). Firn air measurements show an increase of approximately 50

ppt (10%) from 1940 to 1990.

Stratospheric Chlorine and Fluorine

• The stratospheric chlorine burden derived by the ground-based total column and space-based measurements

of inorganic chlorine is now in decline. This is consistent with the decline in tropospheric chlorine from longlived halocarbons (CFCs, carbon tetrachloride, methyl chloroform, HCFCs, methyl chloride, and halon1211). The burden of total stratospheric chlorine derived from satellite measurements agrees, within ±0.3 ppb

(about 12%), with the amounts expected from surface data when the delay due to transport is considered. The uncertainty in this burden is large relative to the expected chlorine contribution from shorter-lived gases (0.05 to 0.1 ppb,

see Chapter 2).
LONG-LIVED COMPOUNDS

1.2

• Column and satellite measurements of hydrogen fluoride continued to show the reduction in the upward

trend that began in the late 1990s. This reduction is consistent with the change in tropospheric fluorine burden

derived from CFCs, HCFCs, and hydrofluorcarbons (HFCs).

Total Tropospheric Bromine

• Total organic bromine from halons and methyl bromide (CH3Br) peaked in about 1998 at 16.5 to 17 ppt and

has since declined by 0.6 to 0.9 ppt (3 to 5%). This observed decrease was solely a result of declines observed for

methyl bromide. Bromine (Br) from halons continues to increase, but at slower rates in recent years (+0.1 ppt Br/yr

in 2003-2004).

• Atmospheric amounts of methyl bromide declined beginning in 1999, when industrial production was reduced.

By mid-2004, mixing ratios had declined 1.3 ppt (14%) from the peak of 9.2 ppt measured before 1999. Reported

production of methyl bromide for emissive uses decreased by 50% during this same period.

• Both the recently observed decline and the 20th-century increase inferred for atmospheric methyl bromide

were larger than expected. Although industrial emissions of methyl bromide were thought to account for 20 (10-

40)% of atmospheric methyl bromide during 1992-1998 (i.e., before production was reduced), the observations are

consistent with this fraction having been 30 (20-40)%. This suggests that fumigation-related emissions could have

a stronger influence on atmospheric methyl bromide mixing ratios than estimated in past Assessments, though

uncertainties in the variability of natural emission rates and loss, and in the magnitude of methyl bromide banked in

recent years, influence our understanding of this sensitivity.

• Mixing ratios calculated from updated emission estimates are in good agreement with the measurements for

halon-1211 but they exceed all the halon-1301 measurements since 1980 by more than 10%. Atmospheric

increases in halon-1211 (0.06 ppt/yr or 1%/yr) in 2000-2004 were about half those in 1996-2000. It is currently

unclear whether atmospheric mixing ratios of halon-1301 continue to increase.

Equivalent Effective Stratospheric Chlorine (EESC) and Effective Equivalent Chlorine (EECl)

• The decrease in EECl over the past 10 years has been 20% of what would be needed to return EECl values to

those in 1980 (i.e., before the Antarctic stratospheric ozone hole). EECl has decreased at a mean rate of 29 ppt

EECl per year from 1994 to 2004 in the lower atmosphere. This amounts to a total decline of 277 ppt or 8-9% for

EECl over this period. The reduction in stratospheric EESC is somewhat less because it takes a few years for nearsurface trends to be reflected in the stratosphere. These declines are slightly larger than projected in the previous

Assessment primarily because of the decline in methyl bromide.

• The decline in the short-lived methyl chloroform and methyl bromide contributed the most to the decline in

the effective equivalent chlorine in the atmosphere. The total decline was about 120 ppt between 2000 and 2004,

of which about 60 ppt was due to the decline of methyl chloroform and about 45 ppt due to methyl bromide. The

CFCs together accounted for less than 23 ppt of this decline. The contribution of HCFCs was 12 ppt in the other

direction, i.e., an increase.

Emission Estimates

• Global emissions of CFC-11 (88 Gg/yr), CFC-12 (114 Gg/yr), and CFC-113 (6 Gg/yr) in 2003 were approximately 25%, 25%, and 3% of their maximum values around 1986. Emissions of CFC-11, CFC-12 and CFC113 have all continued to decrease since 2000.

• Reported regional emission estimates for CFCs, methyl chloroform, and CCl4 have been summed for the first

time. Differences between developed and developing regions are indicative of the differing schedules of phase-out.

However, the patchiness in the present coverage (especially in developing regions such as Southeast Asia) and the

uncertainties in the regional estimates, mean that useful comparisons between summed regional emissions and

global emissions derived from trends cannot currently be made. Regional emission estimates of methyl chloroform

indicate that global emissions after 2000 may have been roughly 22 Gg/yr, which is not statistically different from

the estimate of 12.9 Gg/yr (for 2002) obtained from industry/United Nations Environment Programme data.
LONG-LIVED COMPOUNDS

1.3

• While emissions of HCFC-22 have remained nearly constant from 2000 to 2004, emissions of HCFC-141b

and HCFC-142b decreased by approximately 15% over the same period.

• HFC-23 emissions estimated from atmospheric measurements have increased from about 6 Gg/yr in 1990 to

about 13 Gg/yr in 2001 (an increase of approximately 120%). These emissions are a byproduct of HCFC-22

production. 

Hydrofluorocarbons (HFCs)

• The atmospheric abundances of all measured HFCs are increasing due to their rapid introduction as CFC

and HCFC replacements. HFC-134a concentrations reached 30 ppt in 2004 and are increasing at 3.9 ppt/yr

(13%/yr). Globally averaged concentrations of HFC-125 and HFC-152a were both approximately 3.1 ppt in 2004,

increasing by about 23%/yr and 17%/yr, respectively. HFC-23 mixing ratios in the Southern Hemisphere were 18

ppt in 2004, having increased at a mean rate of 0.7 ppt/yr (4%/yr) in 2001-2004.

Sulfur Hexafluoride (SF6)

• Measurements of SF6 suggest its global average concentration in 2003 was 5.2 ppt and it was growing by 0.23

ppt/yr (4%/yr). Similar percentage rates of increase have been estimated from column measurements over

Switzerland. Temporal extrapolation of the past 20 years of measurements results in significantly smaller atmospheric concentrations in the future than those produced in the most likely emission scenario from the previous

Assessment.

1.1 INTRODUCTION

Ozone depletion observed in the stratosphere over

the past three decades has resulted from the accumulation of trace gases in the atmosphere. This was due to

industrial activities and to a much-reduced extent to agricultural practices, which led to increases in the concentrations of chlorine, bromine, nitrogen, and hydrogen

radicals in the stratosphere. These radicals are produced

from long-lived source gases that, because their atmospheric lifetimes are long compared with the transport

time to reach the stratosphere, are able to be transported

intact into the stratosphere. The Montreal Protocol and

its Amendments and Adjustments have, however, led

to reductions in the tropospheric abundances of the

chlorine- and bromine-containing source gases as a result

of reductions in the production and consumption of manufactured halogenated source gases.

Direct observations of most chemical species that

cause stratospheric ozone depletion began in the late

1970s or later. The ground-based in situ and flask measurements have provided the most complete picture of the

evolution of these species because of the regularity of

the measurements. The complete 20th century and earlier pictures of the accumulation of ozone-depleting

chemicals in the atmosphere also rely on calculations

based on historic emission estimates, and analysis of air

trapped in polar firn and ice or measurements of air

archives. Ground-based remote sensing techniques, in

particular numerous Network for the Detection of

Atmospheric Composition Change (NDACC, formerly

Network for the Detection of Stratospheric Change or

NDSC) sites equipped with Fourier transform infrared

(FTIR) instruments, have also provided multidecadal

information of relevance to this chapter. Satellite observations have also provided information on the most abundant ozone-depleting substances since the mid-1980s.

Although the long-term trends are difficult to measure

with space-based techniques due to the generally short

lifetime of the space missions, they can be inferred by

combining data from multiple instruments operating in

different years. This method was applied by Harries et

al. (2001), using data from the Infrared Radiation Interferometer Spectrometer (IRIS) sensor in 1970 and the

Interferometric Monitor for Greenhouse Gases (IMG)

sensor in 1997, to identify the long-term changes in the

Earth’s outgoing longwave radiation, notably due to the

rise of the chlorofluorocarbons (CFCs) in that time

period. More recently, Rinsland et al. (2005) have compared data from the Atmospheric Trace Molecule

Spectroscopy (ATMOS) missions in 1985 and 1994 and

the Atmospheric Chemistry Experiment - Fourier Transform Spectrometer (ACE-FTS) data in 2004 to determine

lower-stratospheric trends of several target species

during the last two decades.

The atmospheric history of ozone-depleting substances in years before regular measurements began have

been examined using Antarctic firn air samples (Butler et

al., 1999; Sturges et al., 2001; Sturrock et al., 2002; Trudinger et al., 2004; Reeves et al., 2005). These analyses

have revealed that atmospheric mixing ratios of the various CFCs and halons underwent large increases beginning in the 1950s to 1970s. As a result of international

production and consumption restrictions, all of these gases

now exhibit substantially reduced rates of atmospheric

growth, and, for many of them, mixing ratio declines. The

hydrochlorofluorocarbons (HCFCs) first appeared in the

atmosphere in the 1960s to the 1980s and their mixing

ratios continued to increase through 2004. Firn air results

also indicate that nearly all of the CFCs, halons, methyl

chloroform, and carbon tetrachloride (CCl4) present in

today’s atmosphere are the result of anthropogenic production, though the existence of small natural sources or

some anthropogenic production and use in years predating

the oldest firn air samples cannot be ruled out.

This chapter provides a scientific update to various

sections of Chapter 1 of the previous World Meteorological Organization Assessment Report (WMO, 2003;

hereafter referred to as WMO 2002). It reports the measurements of long-lived compounds obtained from many

typical observation means (Table 1-1), including firn air

analysis, ground-based in situ and remote sensing stations,

and satellite observations, for ozone-depleting substances

and for other climate-related gases through 2004. The

latter gases, which include the hydrofluorocarbons, are

also discussed in Section 1.4, because they can contribute

indirectly to ozone change and ozone recovery through

their effects on the radiative forcing of the atmosphere.

Trace gases with lifetimes shorter than 0.5 years (in particular nitrogen and hydrogen radicals) are reported in

Chapter 2. The observations provide important constraints

on past emissions and thereby on future emissions.

Chapter 1 compares the current observed trends with the

expectations based on standard emission scenarios.

Further results on EECl (effective equivalent chlorine) and

EESC (equivalent effective stratospheric chlorine) calculations from assumed future emission scenarios and for

calculations of the remaining banks may be found in

Chapter 8. The reader is referred to Table 1-4 for the chemical formulae of the gases discussed.

LONG-LIVED COMPOUNDS

1.5
LONG-LIVED COMPOUNDS

Table 1-1. Measurements of long-lived compounds.*

Name Technique

Location

Firn air

Law Dome

Antarctica Firn ice

On-Site Sampling — Ground-Based

AGAGE – Advanced Global Atmospheric Gases Experiment

Cape Grim, Tasmania (41°S, 145°E) In situ GC measurements

Cape Matatula, Samoa (14°S, 171°E)

Ragged Point, Barbados (13°N, 59°W)

Mace Head, Ireland (53°N, 10°W)

Trinidad Head, United States (41°N, 124°W)

SOGE – System for Observation of Halogenated Greenhouse Gases in Europe

Jungfraujoch, Swiss Alps (46.5°N, 8.0°E) In situ GC measurements

Ny Ålesund, Norway (78°N, 11.5°E)

Monte Cimone, Italy (44° N, 10.5°E)

NOAA/ESRL – National Oceanic and Atmospheric Administration/Earth System Research Laboratory

South Pole, Antarctica 1,2,3,4 Flasks-GC, and in situ-GC

Palmer Station, Antarctica 4 Flasks-GC

Cape Grim, Tasmania 2,3,4 Flasks-GC

Cape Matatula, Samoa 1,2,3,4 Flasks-GC, and in situ-GC

Mauna Loa, United States 1,2,3,4 Flasks-GC, and in situ-GC

Cape Kumukahi, United States 2,3,4 Flasks-GC

Niwot Ridge, United States 1,2,3 Flasks-GC, and in situ-GC

Wisconsin, United States 2,4 Flasks-GC

Harvard Forest, United States 4 Flasks-GC

Barrow, United States 1,2,3,4 Flasks-GC, and in situ-GC

Alert, Canada 2,3,4 Flasks-GC

UCI – University of California at Irvine 

40-45 Pacific sites from 71°N to 47°S sampled approximately 

twice per season. www.physsci.uci.edu/~rowlandblake/ Flasks-GC

samplinglocationsmap.html

UEA – University of East Anglia

Cape Grim, Tasmania Flasks-GC, Archived air

Remote Sensing — Ground-Based

Kiruna, Sweden (67.8°N, 20.4°E) Ground-based infrared FTS

Jungfraujoch, Swiss Alps (46.5°N, 8.0°E) Total column measurements

Kitt Peak, Arizona (32.0°N, 111.6°W)

Izaña, Tenerife (28.3°N, 16.5°W)

Remote Sensing — Satellite

HALOE - Halogen Occultation Experiment

U.S. experiment (1991-2005) Infrared gas filter correlation

Latitudinal coverage from 80°S to 80°N/1 year and broadband filter

radiometry (solar occultation)

Stratospheric-mesospheric profiles

IMG - Interferometric Monitor for Greenhouse Gases

Japanese experiment (1996-1997) Spaceborne infrared FTS (nadir)

Global coverage Tropospheric columns

1.6
1.2 HALOGENATED OZONE-DEPLETING

GASES IN THE ATMOSPHERE

1.2.1 Updated Atmospheric Observations

of Ozone-Depleting Gases

In this section, measurements of halogenated gases

by the ground-based networks are summarized. As indicated in Table 1-1, these network measurements include

those by the Advanced Global Atmospheric Gases

Experiment (AGAGE), the System for Observation of

Halogenated Greenhouse Gases in Europe (SOGE,

although the measurements used here were from

Jungfraujoch and Ny Ålesund only), the National Oceanic

and Atmospheric Administration / Earth System Research

Laboratory (NOAA/ESRL, formerly NOAA/CMDL) and

the University of California at Irvine (UCI). In a number

of cases this also includes measurements of air samples

collected at Cape Grim, Australia since 1978 that have

been archived (Langenfelds et al., 1996). The section

extends the time series described in WMO 2002 (Montzka

and Fraser et al., 2003) beyond 2000 and emphasizes the

measurements from 2001 to the end of 2004. The time

series for these gases shown in the Figures and Tables

throughout this Chapter are from monthly means of data

from individual sites that have been combined to produce

monthly hemispheric and global means. Annual means of

these are also used in some Figures and Tables.

Table 1-2 shows annual means of the measured

ozone-depleting halocarbons. The precisions and accuracies of the measurements may mostly be judged by the differences between the reported values. Only for CFC-12,

HCFC-22, and halon-1301 are the differences in the trends

observed by the various groups statistically significant.

All uncertainties are one sigma unless otherwise specified. Terms used to describe measured values throughout

Chapter 1 are mixing ratios (for example parts per trillion,

ppt, pmol/mol), mole fractions, and concentrations. These

terms have been used interchangeably and, as used here,

they are all considered to be equivalent.

Methyl chloride and methyl bromide have substantial natural sources and the emissions of these gases are

also discussed in their subsections. The emissions of the

other halocarbons are discussed in Section 1.3.

1.2.1.1 CHLOROFLUOROCARBONS (CFCS)

CFC-11 (CCl3F), CFC-12 (CCl2F2), and CFC-113

(CCl2FCClF2) are routinely measured by the three

ground-based gas-sampling networks (see Table 1-1)

using gas sampling techniques. Results from these independent laboratories generally agree to within 2% (Figure

1-1). CFC-11 and CFC-12 are also measured using

ground-based infrared solar absorption spectroscopy, such

as at the Jungfraujoch station. From space, CFCs have

been measured by a variety of shuttle- or satellite-based

instruments operating in the infrared spectral region.

Most of these data concern CFC-11 and CFC-12. They

were obtained from thermal infrared nadir sounders

(Harries et al., 2001; Coheur et al., 2003; Dufour et al.,

2005), limb instruments (Bacmeister et al., 1999;

Hoffmann et al., 2005), or solar occultation sounders

(Khosrawi et al., 2004; Rinsland et al., 2005).

Recent measurements from the gas sampling networks indicate that the CFC-12 global trend is now

approximately zero, and two of the three networks plus

the remote sensing measurements at Jungfraujoch indiLONG-LIVED COMPOUNDS

1.7

Table 1-1, continued.

MIPAS - Michelson Interferometer for Passive Atmospheric Sounding

European experiment (launched in 2002) Spaceborne infrared FTS (limb)

Global coverage Strato-mesospheric profiles

ACE/FTS - Atmospheric Chemistry Experiment

Canadian experiment (launched in 2003) Spaceborne infrared FTS (solar occultation)

Latitudinal coverage from about 80°S to 80°N/1 year Upper troposphere to the lower thermosphere

MLS - Microwave Limb Sounder

U.S. experiment (launched in 2004) Microwave sounder (limb)

Latitudinal coverage from 82°N to 82°S daily Upper troposphere to the lower thermosphere

* The list provided here is restricted to the measurements reported in this chapter. GC = gas chromatographic.

1 Data from in situ instruments at these stations are used by NOAA to derive global tropospheric means for CFCs and CCl4. 2 Data from flasks collected at these stations are used by NOAA to derive global tropospheric means for CH3CCl3. 3 Data from flasks collected at these stations are used by NOAA to derive global tropospheric means for HCFCs and halons. 4 Data from flasks collected at these stations are used by NOAA to derive global tropospheric means for CH3Br.
cate that Northern Hemispheric mole fractions are now

decreasing. These observations are consistent with the

analyses of Rinsland et al. (2005) using ATMOS and ACE

satellite data. The measured annual column change above

Jungfraujoch in 2003-2004 is −0.11 × 1014 molec cm-2

(−0.16%/yr) (Zander et al., 2005) compared with 0.50 ×

1014 molec cm-2 (+0.71%/yr) in 1998 (WMO 2002) and

0.27 × 1014 molec cm-2 (+0.39%/yr) in 1999-2000. Trends

from column measurements in %/yr should be almost

directly comparable with trends for similar periods

derived from the sampling networks, because of the long

lifetimes of the species being discussed. Averaged over

Kiruna (68°N) and Izaña (28°N), the column trend from

2000 to 2005 is −0.09 ± 0.10%/yr (update from Kopp et

al., 2003, and Schneider et al., 2005). Since the CFC-12

lifetime is approximately 100 years, the measured rate of

change indicates that there are still significant emissions

of CFC-12. Differences between measured Northern

Hemisphere and Southern Hemisphere mole fractions

have been decreasing since approximately the end of the

1980s (Figure 1-1). The hemispheric difference in 2004

was still approximately 3 ppt (0.6%), indicating that the

remaining emissions are still occurring predominantly in

the Northern Hemisphere.

LONG-LIVED COMPOUNDS

1.8

 

200

300

400

500

mole fraction (ppt)

CFC-12

AGAGE

ESRL

UCI

500

520

540

560

mole fraction (ppt)

 

 

100

150

200

250

mole fraction (ppt)

CFC-11

AGAGE

ESRL

UCI

245

255

265

275

mole fraction (ppt)

1980 1985 1990 1995 2000 2005

 

20

40

60

80

mole fraction (ppt)

CFC-113

AGAGE

ESRL

UCI

75

80

85

90

mole fraction (ppt)

Year

Figure 1-1. Hemispheric monthly means of the major chlorofluorocarbons CFC-12, CFC-11, and CFC-113

(crosses for Northern Hemisphere and triangles for Southern Hemisphere). Measurements from the AGAGE

network (Prinn et al., 2000 updated), the NOAA/ESRL network (Montzka et al., 1999 updated; Thompson et al.,

2004), and UCI (D.R. Blake et al., 1996; N.J. Blake et al., 2001) are shown. To increase visibility, recent measurements are depicted on a larger scale in the inserts (the scales are on the right-hand sides of the panels).
LONG-LIVED COMPOUNDS

Table 1-2. Mole fractions and growth rates of ozone-depleting gases.

Chemical Common or Annual Mean Laboratory, Method

Formula Industrial Name Mole Fraction (ppt) Growth (2003-2004)

2000 2003 2004 (ppt yr) (% yr)

CFCs

CCl2F2 CFC-12 543.0 544.2 543.8 −0.6 −0.1 AGAGE, in situ

534.1 535.2 535.0 −0.2 0.0 NOAA*

535.6 537.9 539.7 1.8 0.3 UCI

CCl3F CFC-11 260.6 254.9 253.1 −1.8 −0.7 AGAGE, in situ

262.5 256.6 254.7 −1.9 −0.7 NOAA*

261.2 256.1 253.7 −2.4 −0.9 UCI

263.9 257.3 254.7 −2.7 −1.0 SOGE, in situ (Europe)

CCl2FCClF2 CFC-113 81.8 79.6 78.7 −0.9 −1.1 AGAGE, in situ

82.1 80.5 79.5 −0.9 −1.2 NOAA*

81.1 79.5 79.1 −0.4 −0.5 UCI

CClF2CClF2 CFC-114 17.21 17.24 17.31 0.07 0.4 AGAGE, in situ

including -114a (Cape Grim only)

CFC-114 14.90 14.72 −0.18 −1.2 UCI

CClF2CF3 CFC-115 8.10 8.26 8.34 0.08 1.0 AGAGE, in situ

8.19 8.18 8.29 0.11 1.3 SOGE, in situ (Europe)

HCFCs

CHClF2 HCFC-22 142.8 159.0 163.9 4.9 3.0 AGAGE, in situ

141.7 157.9 162.1 4.2 2.6 NOAA, flasks

161.9 169.1 7.2 4.4 UCI

CH3CCl2F HCFC-141b 12.82 16.70 17.42 0.72 4.2 AGAGE, in situ

12.66 16.58 17.18 0.60 3.6 NOAA, flasks

15.05 18.57 19.20 0.60 3.2 SOGE, in situ (Europe)

CH3CClF2 HCFC-142b 12.46 14.74 15.38 0.64 4.2 AGAGE, in situ

11.72 13.99 14.53 0.54 3.8 NOAA, flasks

CHCl2CF3 HCFC-123 0.050 0.060 0.064 0.004 6.5 AGAGE, in situ

(Cape Grim only)

CHClFCF3 HCFC-124 1.40 1.63 1.64 0.01 0.1 AGAGE, in situ

1.46 1.71 1.72 0.01 0.1 SOGE, in situ (Europe)

Halons

CBr2F2 halon-1202 0.044 0.040 0.038 −0.002 −5.1 UEA, flasks

(Cape Grim only)

CBrClF2 halon-1211 4.12 4.28 4.33 0.05 1.2 AGAGE, in situ

3.98 4.13 4.15 0.02 0.0 NOAA, flasks

4.43 4.68 4.77 0.09 1.9 UEA, flasks

(Cape Grim only)

4.32 4.52 4.62 0.08 1.8 SOGE, in situ (Europe)

CBrF3 halon-1301 2.95 3.04 3.14 0.10 3.2 AGAGE, in situ 

2.57 2.61 2.60 −0.01 0.0 NOAA, flasks

2.26 2.38 2.45 0.07 2.9 UEA, flasks

(Cape Grim only)

3.04 3.06 3.16 0.10 3.2 SOGE, in situ (Europe)

CBrF2CBrF2 halon-2402 0.425 0.427 0.427 0.0 0.0 UEA, flasks

(Cape Grim only)

0.48** -- 0.48 NOAA, flasks

1.9
The sampling networks indicate that global mole

fractions of CFC-11 decreased by approximately 2.0 ± 0.1

ppt/yr (0.8%/yr) from 2000 to 2004 (Table 1-2). This rate

of decrease is larger than that reported in WMO 2002

(Montzka and Fraser et al., 2003), 1.4 ppt/yr (0.5%/yr),

for 1996 to 2000. Zander et al. (2005) indicate that the

CFC-11 column change and the corresponding annual

trend have remained constant since the late 1990s over

the Jungfraujoch, amounting to −1.85 × 1013 molec cm-2

(−0.59%/yr in 2004). A slightly smaller downtrend is to

be expected in the CFC-11 column than in the surface

measurements, by approximately 0.1%/yr, because the

proportion of CFC-11 that is in the stratosphere is predicted to have been continuing to increase slowly. The

decline in the tropospheric CFC-11 mole fractions from

2000 to 2004 is approximately 20% less than that projected

in the Ab (baseline, most likely) scenario of WMO 2002

(Table 1-14, Chapter 1, WMO (2003)). The measured

hemispheric difference for CFC-11 has been decreasing

(Figure 1-1), but it was still approximately 2 ppt (0.8%) in

2004. This and the 45-year lifetime of CFC-11 indicate

that significant Northern Hemispheric emissions are still

occurring.

CFC-113 has decreased by approximately 1% (0.8

ppt) from 2003 to 2004. This is larger than the rate of

decrease of 0.5%/yr reported for 1999-2000 in WMO 2002

(Montzka and Fraser et al., 2003). Since the lifetime is

estimated to be 85 years, the recent decrease indicates that

global emissions of CFC-113 are now small. Consistent

with this, the hemispheric gradient is less than 1 ppt

(Figure 1-1). A global decrease of CFC-113 of 1%/yr provides an upper limit to its atmospheric lifetime of 100

years.

CFC-114 (CClF2CClF2) and CFC-115 (CClF2CF3)

have been measured in Cape Grim (41°S) air samples since

1978. As noted by Sturrock et al. (2001), and by Mangani

et al. (2000) for Antarctica, the mixing ratio of CFC-114

was not changing in the atmosphere in mid-2000. Recent

measurements indicate that there have only been small but

insignificant changes of CFC-114/-114a (Table 1-2). In

contrast, CFC-115, which is used with HCFC-22 (CHClF2)

in refrigeration blends, has continued to increase at about

1%/yr (0.08 ppt/yr), which is only half of the growth rate

reported in WMO 2002 (Montzka and Fraser et al., 2003).

1.2.1.2 HYDROCHLOROFLUOROCARBONS (HCFCS)

Measurements of HCFCs from the gas sampling

networks have been updated through 2004 (O’Doherty et

al., 2004; Krummel et al., 2004). Furthermore, recent

remote measurements of HCFC-22 (Rinsland et al., 2005)

and HCFC-142b (Dufour et al., 2005), which show genLONG-LIVED COMPOUNDS

1.10

Table 1-2, continued.

Chemical Common or Annual Mean Laboratory, Method

Formula Industrial Name Mole Fraction (ppt) Growth (2003-2004)

2000 2003 2004 (ppt yr) (% yr)

Chlorocarbons

CH3Cl Methyl chloride 532.1 523.4 527.8 4.4 0.8 AGAGE, in situ

(Ireland/Cape Grim average)

529.5 523.0 526.0 3.0 0.6 SOGE, in situ (Europe)

CCl4 Carbon 96.4 93.5 92.6 −0.8 −1.0 AGAGE, in situ

tetrachloride 100.2 97.2 96.1 −1.1 −1.1 NOAA*

99.2 96.3 95.1 −1.2 −1.2 UCI

CH3CCl3 Methyl 45.6 26.5 22.0 −4.5 −18.6 AGAGE, in situ

chloroform 45.8 26.5 22.0 −4.5 −18.7 NOAA, flasks

47.7 28.3 23.9 −4.4 −16.8 UCI

44.3 26.7 22.2 −4.5 −18.5 SOGE, in situ (Europe)

Bromocarbons

CH3Br Methyl 10.3 8.8 9.1 0.3 3.3 AGAGE, in situ

bromide (Ireland only)

8.7 8.1 7.9 −0.2 −2.4 NOAA, flasks

All values are estimates for the global troposphere unless specified otherwise.

* These NOAA values consist of a combination of flask and in situ measurements.

**Earlier halon-2402 data from NOAA are for 1996, not 2000.
eral consistency with the AGAGE and the NOAA/ESRL

values, have recently become available from the ACE

satellite mission.

Global mixing ratios of the three most abundant

HCFCs have all increased due to sustained emissions

(resulting from their being substitutes for CFCs and other

ozone-depleting substances (ODSs) (Figure 1-2). The

annual global mole fraction of HCFC-22 (CHClF2) was

approximately 165 ppt in 2004, with an averaged annual

growth rate of about 4.9 ± 0.5 ppt/yr (3.2 ± 0.3%/yr) (2000-

2004, with more weight being given to the more precise

trend estimates from AGAGE and NOAA/ESRL).

Combined vertical column abundances of HCFC-22

above Kiruna (68°N) and Izaña (28°N) yield a trend of

3.5 ± 0.3%/yr (1.1 × 1014 molec cm-2 yr-1) for 2000-2004

(updated from Kopp et al., 2003; Schneider et al., 2005).

LONG-LIVED COMPOUNDS

1.11

1980 1990 2000 2010

YEAR

0

10

20

30

40

MOLE FRACTION (ppt)

HCFC-141b

1980 1990 2000 2010

YEAR

0

10

20

30

MOLE FRACTION (ppt)

HCFC-142b

1980 1990 2000 2010

YEAR

0

50

100

150

200

MOLE FRACTION (ppt)

HCFC-22

Figure 1-2. Annual global mean mole fractions for HCFC-141b, HCFC-142b, and

HCFC-22 derived from the AGAGE (green)

and NOAA/ESRL (red) networks of groundbased sites (O’Doherty et al., 2004; Montzka

et al., 1999). The early measurements are

based on analysis of the air archived in

Australia (for HCFC-22: from Miller et al.,

1998, green line before 1995; for HCFC-141b

and HCFC-142b: from Oram et al., 1995, blue

lines before 1995); global means have been

calculated from these values using the 12-box

model. Also shown are tropospheric annual

values at the beginning of the years 1990 to

2005 from scenario Ab given in Table 1-14 of

WMO 2002 (black). Atmospheric lifetimes

used in the calculations for the three species

were 9.3, 17.9, and 12.0 years, respectively.
Column abundances above the Jungfraujoch station (last

reported for 1986 to 2001 in Figure 1-2 of WMO 2002) have

been measured since 1986. The column abundances found

at the extremes of the Jungfraujoch observational time base

are 0.90 × 1015 molec cm-2 in January 1986 and 2.42 × 1015

molec cm-2 in December 2004. This corresponds to an

increase of a factor of nearly 2.7 over this 19-year period

(Zander et al., 2005). This factor is only a few percent larger

than (i.e., not significantly different from) that determined

from the Australian archived air (Langenfelds et al., 1996),

AGAGE, and NOAA measurements.

There has been a substantial slowdown in the rates

of accumulation of HCFC-141b (CH3CCl2F) and HCFC142b (CH3CClF2), particularly in the last two years. The

global mean mixing ratio of HCFC-142b has increased to

about 15 ppt in mid-2004, compared with about 12 ppt in

mid-2000, with an annual average growth rate of 0.6 ± 0.1

ppt/yr (4%/yr, 2003-2004). This is approximately half of

the previously reported growth rate of 1.1 ppt/yr averaged

from 1999-2002, based on data from both AGAGE and

NOAA/ESRL measurements (O’Doherty et al., 2004).

The growth rate of HCFC-141b has also slowed from an

annual average rate of about 1.6 ppt/yr in 1999-2000 to

0.6 ± 0.1 ppt/yr (4%/yr) in 2003-2004, resulting in a global

mean surface mole fraction of 17.3 ± 0.2 ppt in mid-2004.

Northern Hemisphere mean mole fractions were approximately 2.5 ppt and 1.4 ppt higher than in the Southern

Hemisphere, for HCFC-141b and 142b respectively, in

mid-2004.

Figure 1-2 shows that the slowdown in the growth

rates of all three gases in the atmosphere was not anticipated in the projections made in the Ab scenario in WMO

2002. The observed growth rates for 2000-2004 were 4.9,

1.1, and 0.6 ppt/yr for HCFC-22, HCFC-141b, and HCFC142b, respectively, versus projected growth rates of 6.6,

2.6, and 1.6 ppt/yr.

Figure 1-3 shows monthly mean baseline mixing

ratios (and occasional pollution events) for HCFC-123

(CHCl2CF3) measured at Cape Grim from 1998 through

2004 (Krummel et al., 2004). Its mixing ratio in 2004 is

0.064 ppt. A strong annual cycle is evident because of the

annual cycle in hydroxyl and the short atmospheric lifetime

(1.3 years) of HCFC-123. Overall the atmospheric mixing

ratios have been increasing over the seven-year period at

approximately 6 ± 1%/yr. These are the only reported measurements of HCFC-123 since a value of 0.03 ppt for 1996

by Oram (1999). Despite its small mixing ratio, there is

considerable interest in this gas because as a refrigerant it is

operated at sub-ambient temperatures and leakage is therefore exceptionally small. Moreover what does get into the

atmosphere is rapidly destroyed. Equally important it has a

very low Ozone Depletion Potential (ODP).

Measurements of HCFC-124 (CHClFCF3), which

were first reported in WMO 2002 (Montzka and Fraser et

al., 2003), have been updated from AGAGE measurements

at the two AGAGE stations at Mace Head, Ireland, and Cape

Grim, Tasmania (Prinn et al., 2000). Trends and annual

global means are reported in Table 1-2. The increase of

HCFC-124 has been slowing rapidly since 1998 (see also

WMO 2002). No new measurements of HCFC-21 (CHCl2F)

have been reported since WMO 2002. The atmospheric histories of several HCFCs have been reconstructed from

analyses of air trapped in firn, and show that concentrations

at the start of the 20th century were less than 2% of current

mixing ratios (Sturrock et al., 2002).

1.2.1.3 HALONS

From the introduction of halons in the early 1960s,

halon use grew steadily worldwide until the Montreal

Protocol required an end to their consumption in developed countries by the end of 2003; however, production in

developed countries was allowed for essential uses and to

supply the needs of developing countries. Global production of halon-1211 (CBrClF2) and halon-1301 (CBrF3)

peaked in 1988 at 43 Gg/yr and 13 Gg/yr, respectively

LONG-LIVED COMPOUNDS

1.12

1998 1999 2000 2001 2002 2003 2004 2005

Year

0.4

0.8

1.2

.6

0.02

0.04

0.06

0.08

HCFC-123 at Cape Grim, Tasmania

CHCl2CF

mixing ratio (ppt) mixing ratio (ppt) Figure 1-3. Monthly mean baseline mixing ratios

(ppt) of HCFC-123 measured at Cape Grim, Tasmania (upper panel) (Krummel et al., 2004). The

lower panel shows the continuous series of the measurements with the pollution events included.
(UNEP, 2003). In developing countries, halon-1211 production began in the 1980s and showed a growth curve

similar to that in developed countries until Multilateral

Fund projects began to reverse that trend in the 1990s

(UNEP, 2003). A third halon, halon-2402 (CBrF2CBrF2),

was used predominantly in the former Soviet Union. No

information on the production of halon-2402 before 1986

has been found. Fraser et al. (1999) developed emission

projections for halon-2402 based on atmospheric measurements. They reported that the emissions grew steadily

in the 1970s and 1980s, peaking in the 1988-1991 time

frame at 1.7 Gg/yr. They found these results to be qualitatively consistent with the peak production of 28,000 ODP

tonnes reported by the Russian Federation under Article

VII of the Montreal Protocol (or assuming all production

was halon-2402 and an ODP of 6, a peak production of

approximately 4.650 Gg/yr).

Figure 1-4 shows the comparisons of the global

mean mole fractions for halon-1211 and halon-1301 from

the twelve times per day measurements by AGAGE (green

lines) and the multiple samples per month analyzed by

NOAA/ESRL (red lines). Also shown are the values

derived from measurements of the air that has been

archived several times per year in Australia and analyzed

at the University of East Anglia (UEA) (blue lines; Oram

et al., 1995; Fraser et al., 1999).

Observations from NOAA/ESRL, AGAGE, and

UEA since 2000 suggest that the rate of increase of halon1211 has continued to slow down, although halon-1211

mixing ratios continue to increase in the global atmosphere. The ground-based sampling network measurements

of halon-1211 show that the global tropospheric mixing

ratio of halon-1211 in 2004 was approximately 4.4 ppt,

with a calibration uncertainty of approximately 0.3 ppt

(Table 1-2). The mean rate of increase from 2000 to 2004

was 0.06 ± 0.01 ppt/yr (1.4%/yr). This is significantly

lower than the rate of increase in 1999-2000, which was

0.12 ppt/yr (WMO 2002, Montzka and Fraser et al., 2003).

For halon-1301, the AGAGE, UEA, and NOAA/-

ESRL measurements show a large disparity in mixing

ratio and rate of change during 2000-2004. The NOAA

data show a negligible change (0.01 ± 0.01 ppt/yr),

whereas AGAGE and UEA show an average rate of

increase of 0.05 ± 0.01 ppt/yr (1.8%/yr). This discrepancy

was not present in the pre-2000 measurements, perhaps

because AGAGE measurements only began in 1998. Thus

it is currently unclear whether there has been a significant

slowdown in the growth rate compared with that reported

in WMO 2002 of 0.06 ppt/yr for 1999-2000 (Montzka and

Fraser et al., 2003). Because of this, the observations do

not provide as strong a constraint on the emissions as they

do for the CFCs, HCFCs, and halon-1211.

The measurements of the Australian air archive by

UEA indicate that halon-2402 has remained approximately

constant over the 2000-2004 period, and a fourth halon,

1202 (CBr2F2), which is not controlled under the Montreal

Protocol, has declined by approximately 14%. The projections for this period given in WMO 2002 were that

halon-2402 would have declined slowly and halon-1202

would have declined by about 40%. Because the atmospheric lifetime of halon-1202 is approximately three years

(Fraser et al., 1999), some residual production of halon1202 is indicated, perhaps from over-bromination during

the small remaining production of halon-1211.

All four halons have been measured in firn air from

Dome C (75°S, Antarctica), Devon Island (75°N, Canada),

and Greenland (75°N) (Reeves et al., 2005). At the base

of the firn, the halon concentrations are below detection

(<0.001 ppt), confirming that these species are entirely

anthropogenic. The global halon mole fractions derived

from these firn air samples are consistent with those

LONG-LIVED COMPOUNDS

1980 1990 2000 2010

YEAR

0

2

4

6

MOLE FRACTION (ppt)

HALON-1211

1980 1990 2000 2010

YEAR

0

2

3

4

MOLE FRACTION (ppt)

HALON-1301

Figure 1-4. Annual global mean mole fractions for

halon-1211 and halon-1301 derived from measurements at multiple sites in the AGAGE (green lines)

and NOAA/ESRL (red diamonds, an update to Butler

et al., 1998; red lines, from Montzka et al., 1999) networks. Also shown are annual global means calculated from measurements of archived air sampled at

Cape Grim, Australia (blue lines, updated from Fraser

et al., 1999) using the 12-box model. These measurements are compared against tropospheric annual

means for the beginning of the years 1990 to 2005

given in Table 1-14 (scenario Ab) of WMO 2002 (black

lines). The greenish gold lines for halon-1211 and

halon-1301 have been calculated from the UNEP

Halons Technical Options Committee emission scenarios (UNEP, 2003) (modified for halon-1211, see

text) using a 12-box model.

1.13
obtained from an analysis of the Cape Grim air archive

(Fraser et al., 1999; Reeves et al., 2005).

By 2000, total tropospheric bromine from halons

was about 8 ppt, having approximately doubled over the

previous 10 years (Figure 1-5). The measurements show

clear signs that the atmospheric accumulation of total

halon bromine is slowing. However the Ab scenario in

WMO 2002 (Montzka and Fraser et al., 2003) projected

that it should stop growing by approximately 2006,

because of reductions in halon-1211 starting in 2003 (the

black line in Figure 1-4). The WMO 2002 Ab scenario

provides a reasonable fit to the halon-1211 observations

(Figure 1-4) but there is little evidence that halon-1211

stopped increasing in the 2000-2004 period. Therefore it

seems likely that the total halon bromine maximum will

be delayed by a year or two.

1.2.1.4 CARBON TETRACHLORIDE (CCl4)

The major use for CCl4 during the 1980s was

believed to be as a feedstock for the production of the

CFCs (Simmonds et al., 1988). Therefore as production

of the CFCs was substantially reduced during the 1990s

because of the Montreal Protocol, atmospheric mixing

ratios of CCl4 decreased (Figure 1-6). Global surface

mixing ratios of CCl4 are shown to have reached a maximum of about 106 ± 1 ppt in late 1990 and they have

declined subsequently at a steady rate of 1.0 ± 0.1 ppt/yr

(about 1%/yr) as a result of reduced emissions (Figure 1-

6). Calibration differences between the different reporting

groups since 1995 are 3-4%. The annual global mean in

2004 was 92-96 ppt (group average is 94.6 ppt) depending

on which group made the measurements. For this gas in

particular, a comparison of the measurements at the

common measurement site in Samoa confirms that the

NOAA/AGAGE difference is due to calibration and not

due to the effects of differing site locations used to produce global means. Significant emissions of CCl4 are still

occurring and these sustain the interhemispheric difference of 1.4 ± 0.3 ppt, which has been relatively constant

since 1995.

In WMO 2002 (Montzka and Fraser et al., 2003),

the estimated atmospheric lifetime of carbon tetrachloride

was reduced to 26 years because of ocean losses (YvonLewis and Butler, 2002). More recently, soil losses measured in North America have led to an estimated lifetime

contribution from soil losses globally of 90 years, with an

uncertainty range of 50 to 418 years (Happell and Roche,

2003; see also Borch et al., 2003). Combining this lifetime contribution with those due to ocean losses (94 years,

with a range of 82 to 191 years) and an atmospheric lifetime from the 1998 WMO Assessment (35 years, with a

range of 21 to 43 years; WMO, 1999) led Happell and

Roche (2003) to infer an overall lifetime for carbon tetrachloride of 20 years. However, based on the discussion in

the emissions section (Section 1.3.2.4), it was decided not

to change the recommended overall lifetime from the

WMO 2002 value of 26 years at this time.

1.2.1.5 METHYL CHLOROFORM (CH3CCl3)

Because of Montreal Protocol restrictions, there has

been a steady decline from an all-time high in methyl chloroform concentrations of more than 130 ppt in 1992

(Figure 1-6). The decline was due to a rapid decline in the

emissions during the 1990s (McCulloch and Midgley,

2001) that resulted from the combination of an efficient

replacement of methyl chloroform as an industrial solvent

and its relatively short atmospheric lifetime of 5 years.

The global mean surface mixing ratio of methyl chloroform has steadily decreased to 22.6 ± 1.0 ppt averaged

over 2004 from 46.4 ppt in 2000 (averages of the AGAGE,

NOAA, and UCI values given in Table 1-2). The decline

since 1998 (Montzka et al., 2000) has occurred at an

approximately exponential rate consistent with relatively

small emissions. Furthermore, elevated concentrations

have practically ceased to exist during pollution events at

background sites such as Mace Head (Ireland), although

LONG-LIVED COMPOUNDS

1.14

1960 1970 1980 1990 2000 2010 2020

Year

0

2

4

6

8

10

Greenland

Antarctica

WMO2002

Bromine (ppt) Figure 1-5. Tropospheric bromine from four halons

(1211, 1301, 2402, 1202) at high northern (red) and

high southern (purple) latitudes, based on modelderived emissions that best-fit the firn air data from

Greenland and Antarctica and the air archive data

from Cape Grim, Tasmania (Reeves et al., 2005).

Those estimates are compared with model estimates

(green) for the Southern Hemisphere (Fraser et al.,

1999), using the Ab halon emission scenario prepared

for WMO 2002 (Montzka and Fraser et al., 2003).
other European remote sites have reported distinct methyl

chloroform excursions (Gros et al., 2003).

Background concentrations of CH3CCl3 from both

hemispheres have been converging steadily from an interhemispheric gradient of more than 10% to an essentially

constant annual mean difference of 2.1 ± 0.4% (corresponding to an uncertainty of ±0.1 ppt) since 1999. In fact,

during the warm season, Northern Hemispheric background concentrations can even be lower than in the

Southern Hemisphere at the same time, due to the different

hemispheric hydroxyl radical (OH) abundances. Calibration scales between the networks of NOAA/ESRL and

AGAGE used to differ by several percent (WMO 2002,

Montzka and Fraser et al., 2003), but since the newest scale

updates by AGAGE (SIO-98) and NOAA (Hall et al.,

2002) came into force, an excellent agreement exists. At

stations with concurrent measurements from both networks, only small mean offsets of 0.4 ± 1.8% (1 sigma)

are observed (AGAGE-NOAA) between 2000 and 2005.

1.2.1.6 METHYL CHLORIDE (CH3Cl)

Methyl chloride (CH3Cl) is the most abundant

chlorine-containing compound in the atmosphere, with

globally averaged concentrations of about 550 ppt. Before

WMO 2002, there had been an imbalanced budget and

large missing source, but in WMO 2002 tropical and subtropical plants were proposed as potential sources to fill

the gap. Recently, abiotic release of CH3Cl by senescent

or dead leaves has been proposed as a new significant

source in the tropics and subtropics. The short- and longterm trends of atmospheric CH3Cl, and source/sink

estimates for biomass burning, oceans, rice paddies, OH

reaction, and loss to polar oceans have been updated.

LONG-LIVED COMPOUNDS

 

80

85

90

95

100

105

110

mole fraction (ppt)

CCl

AGAGE

ESRL

UCI

1980 1985 1990 1995 2000 2005

Year

 

0

40

80

120

160

mole fraction (ppt)

CH CCl

AGAGE

ESRL

UCI

10

20

30

40

50

mole fraction (ppt)

Figure 1-6. Hemispheric monthly mean of CH3CCl3 and CCl4 (crosses for Northern Hemisphere and triangles

for Southern Hemisphere) from the AGAGE stations (Prinn et al., 2000, 2005), the NOAA/ESRL network

(Thompson et al., 2004; Montzka et al., 1999 updated), and UCI (D.R. Blake et al., 1996; N.J. Blake et al.,

2001). Recent measurements for methyl chloroform are depicted in magnified form in the insert, using the

scale on the right-hand side of the panel.

1.15
Atmospheric Distribution and Trends

Observations from NOAA/ESRL (since 1995) and

AGAGE (since 1998) are consistent with the global

average concentration (~550 ± 30 ppt) and the latitudinal

variation (higher concentrations at the lower latitudes)

reported in WMO 2002 (Simmonds et al., 2004; Yoshida

et al., 2004). These observations also suggest interannual

variability for CH3Cl, with significant enhanced concentrations during 1998 and a subsequent steady decrease until

2001, a trend that may be explained by large-scale forest

fires in Indonesia and Canada during 1997-1998

(Simmonds et al., 2004; see also Table 1-2).

Aydin et al. (2004) and Trudinger et al. (2004) independently produced atmospheric histories of CH3Cl for

the 20th century based on Antarctic firn air measurements.

These two studies both revealed an increase of about 10%

in the 50 years prior to 1990 for atmospheric CH3Cl mixing

ratios over Antarctica, with a rather steady level after 1980

(~480-530 ppt). These tendencies are similar to the finding

of Butler et al. (1999) in WMO 2002. The firn reconstruction of Kaspers et al. (2004), however, showed higher concentrations and a decreasing trend of 1.2 ppt/yr between

1975 and 2001; the difference from the results of the above

studies remains unexplained. The ice core measurements

during the past 300 years reported by Aydin et al. (2004)

suggested a cyclic natural variability on the order of 110

years, which is in phase with the 20th-century rise inferred

from firn air measurements.

Sources

Table 1-3 lists the revised estimates of source and

sink strengths for atmospheric methyl chloride. Tropical

forests remain a major source in the global atmospheric

budget. High concentrations of atmospheric CH3Cl have

been observed in the tropical rainforests in Southeast Asia

(up to 1500 ppt; Yokouchi et al., 2000) and above the

canopy of tropical rainforests in South America (up to 809

ppt; Moore et al., 2005), supporting the significance of

tropical forests as a potentially large CH3Cl source.

However, this source’s strength and geographical distribution are still highly uncertain, mostly due to the large diversity of tropical vegetation.

Hamilton et al. (2003) reported abiotic release of

CH3Cl by senescent or dead plant material, both foliar and

woody, and that these emissions rose dramatically when

temperature increased. They estimated the global annual

CH3Cl emission from the weathering of leaf litter, between

30°N and 30°S, to be 30-2500 Gg/yr, possibly exceeding

the source from live tropical vegetation.

For biomass burning, Andreae and Merlet (2001)

provided a revised inventory (650 ± 325 Gg/yr) based

on evaluation and integration of the available biomassburning emission data. The value is closer to the lower

limit of the estimate range (655-1125 Gg/yr) reported in

WMO 2002, suggesting that biomass burning emissions

may have been overestimated in WMO 2002. Threedimensional (3-D) model studies (Lee-Taylor et al.,

2001; Yoshida et al., 2004) have also shown that a

smaller burning source is more consistent with observations.

The oceanic source estimate, which once had been

understood as the major CH3Cl source, has been revised

downward again. On the basis of measurements of the

solubility of CH3Cl in seawater, Moore (2000) reported a

global annual flux from warm waters ranging from 465 to

495 Gg/yr. Recalculation of the oceanic CH3Cl flux with

the empirical relationship between saturation and SST by

Yoshida et al. (2004) gave a net flux of 350 Gg/yr, which

requires an oceanic emission of 380-500 Gg/yr to compensate for the oceanic uptake of 30-150 Gg/yr (discussed

below). These values are at the lower end of the WMO

2002 estimate range (325-1300 Gg/yr).

Large CH3Cl emission from some salt marshes has

been attributed to the plants growing in these regions

(Rhew et al., 2002; but see also Cox et al., 2004), and the

emissions are probably influenced by the high concentrations of available chloride ions (Cl–

). In addition to whiterot fungal sources, ectomycorrhizal fungi recently have

been found to emit CH3Cl (0.003-65 μg g dry-1d-1; Redeker

et al., 2004), but this source’s global emission strength has

not been estimated. Moore et al. (2005) suggested the possibility that some of the emission from woodrot fungi

might compete with microbial uptake in the soil (see

below). Thus, uncertainties for the fungal emissions of

CH3Cl have been added. Finally, the emissions from rice

paddies have been revised downward from 5.0 to 2.4-4.9

Gg/yr using a new emission algorithm (Lee-Taylor and

Redeker, 2005).

Sinks

The largest sink of CH3Cl from the atmosphere is

the reaction with the OH radical. The rate constant of the

CH3Cl + OH reaction was re-evaluated to be 2.4 × 10-12

e1250/T or 3.6 × 10-14 cm3 molec-1 s

-1 at 298 K (Sander et al.,

2003). The annual loss of CH3Cl due to reaction with OH,

calculated with detailed OH distributions, was 3800-4100

Gg/yr (Lee-Taylor et al., 2001; Yoshida et al., 2004), which

exceeds the best estimate from WMO 2002 of 3180 Gg/yr.

Soils are another important sink for CH3Cl (Khalil

and Rasmussen, 2000), and several bacterial strains capable

of growth on CH3Cl have been isolated recently from a

variety of soils (McDonald et al., 2002; Miller et al., 2004).

In addition to the ubiquitous bacteria capable of growing

LONG-LIVED COMPOUNDS

1.16
aerobically on CH3Cl, the potential importance of anoxic

biodegradation as a CH3Cl sink was shown by Freedman et

al. (2004), but its global strength has not been quantified.

The oceanic uptake to cold waters (i.e., polar

oceans) was re-examined using the measured solubility of

CH3Cl (Moore, 2000), and the estimate has been revised

upward from 37-113 Gg/yr in WMO 2002 to 93-145 Gg/yr.

Studies by Tokarczyk et al. (2003a, b) using a stableisotope incubation technique indicated that microbial

activity plays an important role in CH3Cl degradation in

coastal seawater (Bedford Basin, Nova Scotia) as well as

in the Southern Ocean. Thus, CH3Cl-degrading bacteria

are found in both terrestrial and marine environments. A

common set of genes and proteins is found in certain terrestrial and marine methyl halide-oxidizing bacteria

(Schafer et al., 2005). McDonald et al. (2002) suggested

that these bacteria may play an important role in mitigating

ozone depletion resulting from CH3Cl.

There is no update for the other sinks, such as the

reaction with Cl– and loss to the stratosphere.

Atmospheric Budget

With high CH3Cl emission from plants and dead

leaves in the tropics and subtropics, global sinks of 4300-

6700 Gg/yr could be balanced by known sources (Table 1-

3), although uncertainties in emissions are so large that

the possibility of missing sources or sinks cannot be denied.

LONG-LIVED COMPOUNDS

1.17

Table 1-3. Estimated source and sink strengths for atmospheric methyl chloride from observations

and three-dimensional model studies.

Source or Sink Type Source or Sink 3-D Model Studies

Estimatea Lee-Taylor et al. (2001)b Yoshida et al. (2004)c

(Gg/yr)

Sources

Tropical and subtropical plants 820-8200d

Tropical senescent or dead leaves 30-2500e

Biomass burning 325-1125d,f 733 611

Oceans 380-500g,h 477i 508

Salt marshes 65-440d 170 170

Fungi 43-470d 128 -

Wetlands 48j 48 48

Rice paddies 2.4-4.9k - -

Fossil fuel burning 5-205d

Waste incineration 15-75d 162 162

Industrial processes 10c

Subtotal (1743-13,578) (4098) (4399)

Sinks

OH– reaction 3800-4100h,k 3994

Loss to stratosphere 100-300d -

Cl reaction 180-550d - -

Soil 100-1600d 256 256

Loss to cold waters (polar oceans) 93-145g - 149

Subtotal (4273-6695) (4106) (4399)

a Best estimates that appeared in the 2002 WMO Assessment table (WMO, 2003) are not given here, due to the difficulty of reasonable selection. b Mean value of their best estimates is given. c Model mean is given. d WMO, 2003. e Hamilton et al., 2003. f Andreae and Merlet, 2001. g Moore, 2000. h Yoshida et al., 2004. i Net ocean flux is given.

j Varner et al., 1999; no range was given. k Lee-Taylor and Redeker, 2005.

3850

2380 2900
By adding tropical terrestrial biogenic sources of 2400-

2900 Gg/yr, model simulations of atmospheric CH3Cl

showed general agreement with the observations (LeeTaylor et al., 2001; Yoshida et al., 2004).

An isotopic study (Gola et al., 2005) shows that the

atmospheric isotopic ratio of CH3Cl is not well explained

by known sources and sinks unless the flux from senescent and dead leaves is quite large (1800-2500 Gg/yr). The

values have large uncertainties, however, due to high variability of the stable carbon isotope ratios of biomass

burning emission and tropical plants, as well as limited

information on the isotope ratio of the oceanic source. In

addition, no measurements of CH3Cl emission from senescent and dead leaves have been made in the field.

Our understanding of the budget of CH3Cl has been

revised based upon the considerations discussed above,

but it is still difficult to assess how atmospheric CH3Cl

levels might be affected by changes in global climate and

land-use patterns, mostly due to large uncertainties in the

strengths and geographical distribution of tropical terrestrial sources.

1.2.1.7 METHYL BROMIDE (CH3Br)

Methyl bromide is the most abundant brominated

gas regulated by the Montreal Protocol. Additional observational and modeling studies of methyl bromide have

been published since 2002 Assessment (WMO, 2003).

Recent atmospheric measurements have become available, as have additional results from the analysis of firn

air. Methyl bromide also was measured in ice core air bubbles for the first time. New techniques have become available for the analysis and quantification of sources and

sinks, and progress has been made in refining estimates of

methyl bromide fluxes from the natural environment.

Recent Atmospheric Trends

Since the previous Assessment (WMO, 2003), three

different groups have published multiyear observational

records for CH3Br at different surface sites (Yokouchi et

al., 2002; Montzka et al., 2003; Simmonds et al., 2004).

The updated studies indicate a global mean mixing ratio

for CH3Br of 9.2 ppt in the three years before 1999, which

was the first year industrial production declined as a result

of the Montreal Protocol. The observations also show that

the tropospheric mixing ratio of CH3Br started declining

in 1999, with the largest decline observed in the Northern

Hemisphere (Figure 1-7). By mid-2004, the global tropospheric mixing ratio of CH3Br had declined by approximately 1.3 ppt (14%) from the amount observed in the

years before anthropogenic production began decreasing

(9.2 ppt in 1996-1997). Reported industrial production of

methyl bromide for emissive uses decreased by 50%

during this same period (UNEP, 2002a, updated). Methyl

bromide is unique among ODS regulated by the Montreal

Protocol because it is emitted in substantial quantities from

natural processes. As a result, attributing the observed

declines to reduced industrial production is less straightforward than for other ODS. The coincidence of the

observed decline with reduced anthropogenic production

(as a result of Montreal Protocol), and the observation that

the largest declines were observed in the NH argue that

the atmospheric decline is primarily the result of reduced

anthropogenic emissions. The decline observed through

2004 was somewhat larger than projected by scenario Ab

(see black curve in Figure 1-4) from WMO 2002 (Montzka

and Fraser et al., 2003).

The enhanced decline observed for CH3Br from the

mid-1990s to mid-2004 could imply that its atmospheric

abundance is more sensitive to changes in anthropogenic

production than suggested in the best-estimate budget in

WMO 2002 (Montzka and Fraser et al., 2003), or that nonanthropogenic sources or sinks of CH3Br changed systematically and simultaneously. For example, declines in

CH3Br mixing ratios from 1998 to 2001 may have been

accelerated as the atmosphere recovered from high biomass burning levels in 1998 (Simmonds et al., 2004),

although the influence of burning on CH3Br trends over

longer periods (1996-2004) is not well defined.

Interlaboratory differences for the absolute calibration of methyl bromide in the atmosphere are smaller than

they were in WMO 2002. The range of results from the

three laboratories reporting recent atmospheric data is ±2%.

Preindustrial Mixing Ratios

Methyl bromide was measured in air trapped in ice

cores for the first time (Saltzman et al., 2004). The results

from Siple Dome (81°S, 149°W) suggest that ambient air

mixing ratios of CH3Br over Antarctica during 1700-1900

were 5-6 ppt, or substantially lower than the 8 ppt observed

during the 1990s (Figure 1-8). Independent measurements

of CH3Br in firn air at Law Dome suggest mixing ratios

over Antarctica of 5.5 ppt in the 1930s (Trudinger et al.,

2004). These firn air results are consistent with those

obtained previously from numerous sites across Antarctica: South Pole, Siple Dome, Dome Concordia, and

Dronning Maud Land (Butler et al., 1999; Sturges et al.,

2001). Despite the variations in snow temperature,

humidity, marine influence, and snow accumulation rates

at these different Antarctic sites, similar atmospheric hisLONG-LIVED COMPOUNDS

1.18
tories have been derived for CH3Br from firn and ice-core

results from samples at all these locations. This consistency argues against processes that significantly influence

CH3Br mixing ratios during extended periods in contact

with Antarctic snow and ice. In the Northern Hemisphere,

however, this is not true; firn air results from different sites

in Greenland show unusual variations that can be

explained only by invoking in situ production of CH3Br

(Butler et al., 1999; Sturges et al., 2001).

Although consistent histories have been inferred

from firn air samples collected at South Pole and Law

Dome, the nature of the firn at Law Dome allows the derivation of a history with finer time resolution. For example,

the results from Law Dome imply that the 20th-century

mixing ratio increases for CH3Br in the high-latitude

Southern Hemisphere occurred mostly from 1950 to 1980

(Trudinger et al., 2004), not gradually over the entire century (Butler et al., 1999).

LONG-LIVED COMPOUNDS

1.19

5.0

6.0

7.0

8.0

9.0

10.0

11.0

12.0

1994.0 1996.0 1998.0 2000.0 2002.0 2004.0 2006.0

Year

Atmospheric methyl bromide (ppt)

ALT (82N, 1)

ALT (82N, 2)

SUM (73N, 2)

BRW (71N, 2)

MHD (53N, 2)

MHD (53N, 3)

LEF (46N, 2)

HFM (43N, 2)

THD (41N, 2)

NWR (40N, 2)

SGB (35N, 1)

MLO (19N, 2)

KUM (19N, 2)

SMO (14S, 2)

CGO (40S, 2)

CGO (40S, 3)

P SA (65S, 2)

SPO (90S, 2)

Ab, WMO2002

Global, Surf. (2)

Figure 1-7. Measured mixing ratios of methyl bromide at different surface sites in the Northern and Southern

Hemispheres. Most results are displayed as 12-month running means to highlight the observed long-term

mixing ratio changes (larger seasonal variations that are observed at some sites are minimized here as a

result). Results from the National Institute for Environmental Studies (NIES) are displayed as annual means

connected by lines. Sampling latitudes and the data source are indicated in parentheses next to the site

acronym: from NIES (1; Yokouchi et al., 2002), NOAA/ESRL (2; Montzka et al., 2003), and AGAGE (3; Simmonds

et al., 2004). Sites are listed in the legend with the highest NH latitudes first: ALT (Alert, 82.5°N); SUM (Summit,

72.6°N), BRW (Barrow, 71.3°N), MHD (Mace Head, 53.3°N), LEF (Wisconsin, 45.9°N), HFM (Harvard Forest,

42.5°N), THD (Trinidad Head, 41.0°N), NWR (Niwot Ridge, 40.0°N), SGB (Sagami Bay, 35.0°N), MLO (Mauna

Loa, 19.5°N), KUM (Kumukahi, 19.5°N), SMO (American Samoa, 14.2°S), CGO (Cape Grim, 40.4°S), PSA,

(Palmer, 64.9°S), and SPO (South Pole, 90.0°S). A global mean derived from most of the NOAA/ESRL (2) data

is shown as the thick white line; the thin white line is the global surface mixing ratio of CH3Br from scenario Ab

of WMO 2002 (Montzka and Fraser et al., 2003).
Updated Information on Nonindustrial Sources and 

Sinks

Work has continued since WMO 2002 on identifying and quantifying the contribution of non-industrial

fluxes of methyl bromide to and from the atmosphere.

While progress has been made in improving our understanding of these fluxes, best estimates of their magnitudes

remain essentially unchanged from Table 1-9 in the previous Assessment (Montzka and Fraser et al., 2003).

Regarding the oceans, previously it was learned that

aqueous concentrations and saturations of methyl bromide

in the temperate open ocean depend upon sea surface temperature (Groszko and Moore, 1998; King et al., 2000).

Additional studies have attempted to refine this relationship (King et al., 2002), though it is clear that in some

waters, additional unidentified factors influence methyl

bromide saturation levels (Sturrock et al., 2003; Tokarczyk

and Moore, 2006). These studies have confirmed that the

ocean on average is a net sink of atmospheric methyl bromide. The role of biology in driving seasonal changes in

flux through both production and degradation of methyl

bromide has been further confirmed and quantified in temperate (Yvon-Lewis et al., 2002), polar (Tokarczyk et al.,

2003b; Yvon-Lewis et al., 2004), and coastal (Sturrock et

al., 2003) waters. The findings support the conclusions of

WMO 2002 regarding the lifetime of methyl bromide with

respect to oceanic losses (1.9 (1.1-3.9) years).

The global contribution of plant emissions to the

atmospheric burden of methyl bromide remains poorly

quantified, given the plethora of species not yet studied.

Improvements in quantifying this source may be possible

with the identification of a gene that encodes an enzyme

capable of producing methyl halides from aqueous chloride, bromide, and iodide ions (Rhew et al., 2003a). This

finding may allow a genetic basis for understanding and

predicting methyl halide production by plants. Other

studies relating to plants have led to a slightly reduced

estimate of emissions from rice paddies (Lee-Taylor and

Redeker, 2005).

Uptake by soils has been studied further since the

previous Ozone Assessment (WMO, 2003). A laboratory

study of methyl bromide uptake by boreal soils found

uptake rates similar to those measured in the past (Rhew

et al., 2003b). This work was the first to allow the quantification of both uptake and emission fluxes through use

of isotopically labeled methyl halides. Field measurements continue to show that terrestrial ecosystems can act

both as a source and sink of CH3Br. While net emission

has been observed in coastal grasslands and wetlands in

Tasmania, Australia, net uptake was observed at similar

sites with exposed soil or leaf litter (Cox et al., 2004). A

study of CH3Br losses to a temperate forest soil over the

1999 growing season revealed a loss that was apparently

an order of magnitude less than previous uptake estimates

for this soil type (Varner et al., 2003). Also, a multiyear

study in freshwater peatlands showed a close balance

between emissions and uptake (White et al., 2005), in contrast with earlier studies that suggested large net sources

from these ecosystems. Discrepancies between these

recent studies and prior estimates may be caused by competing production and consumption mechanisms at individual field sites. Emissions from fungi, for example, may

affect the net CH3Br flux measured at different sites

LONG-LIVED COMPOUNDS

1.20

1700 1750 1800 1850 1900 1950 2000

Year

0

2

4

6

8

0

 ice, Siple Dome

 firn, South Pole

 firn, Law Dome

 flask, South Pole

WMO2002

CH3Br (ppt)

Figure 1-8. The history of methyl bromide abundance in the southern polar

atmosphere, as measured in air samples from Antarctic ice and firn and

from flasks collected at the South Pole

(orange symbols, Butler et al. (1999)

profiles dated with the Trudinger et al.

(2004) firn diffusion model; blue symbols, Trudinger et al., 2004; green symbols, Saltzman et al., 2004; purple

symbols, Montzka et al., 2003; http://-

www.cmdl.noaa.gov/noah/) compared

with calculated Southern Hemisphere

CH3Br levels based on the Ab emission

scenario in WMO 2002 (black line)

(Montzka and Fraser et al., 2003).
(Varner et al., 2003; Redeker et al., 2004). Such discrepancies underscore the potential for isotopic studies to provide improved estimates of gross CH3Br fluxes.

Ambient air measurements in a coastal region confirm previous findings from smaller-scale experiments that

methyl bromide can be emitted from coastal areas

(Sturrock et al., 2003; Cox et al., 2005). Enhanced mixing

ratios were also observed at multiple sites on the west coast

of California, though these measurements may have been

influenced by agricultural activities in the region (Low et

al., 2003). Unusual enhancements of CH3Br have been

observed during spring in the Arctic, however, that do not

appear to be explained by surface-based emissions or other

known sources (Wingenter et al., 2003). The authors proposed a photochemical mechanism by which CH3Br may

be produced in the background atmosphere in substantial

amounts, though analogous reactions have been shown to

be quite slow and unlikely.

Estimates of global carbon monoxide (CO) emissions from fires and their interannual variability have been

derived based upon satellite data (Duncan et al., 2003).

These estimates are similar in magnitude to those

described in Andreae and Merlet (2001), who applied

burning-specific emission factors to all major pyrogenic

processes (grassland and forest fires, the burning of biofuels and agricultural residues, and the making and burning

of charcoal) to derive a global emission for methyl bromide of 29 Gg CH3Br/yr. Though this newer source estimate is slightly larger than the 20 (10-40) Gg CH3Br/yr

quoted in the previous Assessment for biomass burning

(Montzka and Fraser et al., 2003), it falls well within the

uncertainties of the older estimate.

On the Imbalance of CH3Br Sources and Sinks and 

the Contribution of Industrially Derived CH3Br

Despite the recent work on identifying and quantifying sources and sinks, best estimates of the magnitudes

of sources cannot sustain measured mixing ratios during

the past 25 years given the best estimates of loss (WMO,

2003). The firn air and ice bubble air data imply that this

discrepancy existed before CH3Br was produced industrially (Reeves, 2003; Saltzman et al., 2004). Though observational evidence for a significant CH3Br source in the

tropics is currently lacking, modeling studies have suggested that such a source could diminish the discrepancy

in a way that is reasonably consistent with observations

(Lee-Taylor et al., 1998; Warwick et al., 2006). Others

(Reeves, 2003) have noted that the discrepancy could also

be reduced if the CH3Br lifetime were longer than 0.7 year

(0.5-0.9 year) (WMO, 1999; WMO, 2003).

The firn air and ice bubble air results also suggest

an atmospheric increase in the Southern Hemisphere

during the 20th century that is larger than expected from

increases in emissions of anthropogenically produced

CH3Br alone (Reeves, 2003; Saltzman et al., 2004). The

increase can be explained if emissions from anthropogenic

activities accounted for a larger fraction of the total flux

than presumed in the “best-estimate” budget (1992 fumigation emissions of 41 Gg/yr in 1992 relative to a total

flux of 204 Gg/yr; Montzka and Fraser et al., 2003), or if

other emissions had also increased over this time, such as

biomass burning. The large uncertainties in CH3Br fluxes

from biomass burning over time prevent deriving tight

constraints on the CH3Br fluxes from industrially derived

production from the 20th century firn air and ice bubble

results (Saltzman et al., 2004).

Previous Assessments have suggested that anthropogenic CH3Br emissions in 1992 of 41 Gg/yr derived

from industrial production accounted for 20 (10-40)% of

the total annual flux at that time (this range encompasses

a total flux estimated by total sources or sinks; WMO,

2003). Alone, the 20th-century increase in atmospheric

CH3Br inferred from firn air and ice bubble air, or the large

decline observed in ambient air since 1999, do not tightly

constrain this fraction additionally. Both the recent observed decline and the 20th-century increase inferred for

atmospheric methyl bromide were larger than expected,

however, given our understanding of the CH3Br budget

and declines in reported production (UNEP, 2002a;

Yokouchi et al., 2002; Montzka et al., 2003; Reeves, 2003;

Saltzman et al., 2004; Trudinger et al., 2004). Because

CH3Br emitted from fumigation changed more than any

other known flux during both periods, the results suggest

a contribution from fumigation-related emissions in the

middle to upper half of the range quoted in the past, or 30

(20-40)%. This suggests that fumigation-related emissions could have a stronger influence on atmospheric

methyl bromide mixing ratios than estimated in past

Assessments, though uncertainties in the variability of natural emission rates and loss, and in the magnitude of

methyl bromide banked in recent years, influence our

understanding of this sensitivity.

The higher contribution of anthropogenic CH3Br to

the total flux may be explained in part by a mean release

fraction of CH3Br after application to soils for agriculture

fumigation purposes that is greater than 50%. In past

WMO Ozone Assessments, a mean release fraction of 50%

had been suggested for this application; the Methyl

Bromide Technical Options Committee (MBTOC), however, suggests a mean release fraction of 70% (40-87%;

UNEP, 2002a). Such a larger fraction would imply emissions from all types of fumigation of 51 Gg instead of 41

Gg in 1992; this additional emission would partially reduce

the imbalance between known sources and sinks. A relaLONG-LIVED COMPOUNDS

1.21
tively larger contribution of fumigation-related emissions

could also be explained if the total atmospheric lifetime of

CH3Br were longer than 0.7 year (Yokouchi et al., 2002;

Reeves, 2003).

1.2.2 Lifetimes

1.2.2.1 HALOCARBON LIFETIME UPDATES AND

VARIABILITY

After being emitted, most environmentally important trace gases, including methane and other organic compounds, carbon monoxide, nitrogen oxides, and sulfur

gases, are removed from the atmosphere by oxidation. The

strengths of the tropospheric sinks are essentially determined by the concentrations of oxidants, in particular the

highly reactive hydroxyl radical (OH) that dominates the

removal of many gases, including methane (CH4), other

hydrocarbons, CO, and natural and anthropogenic hydrohalocarbons. For that reason, global OH levels are often

taken as a proxy for the oxidizing efficiency of the atmosphere. Through the OH feedback, even changes in the tropospheric emissions of gases that are too short lived to

reach the stratosphere can affect the stratosphere when

they influence tropospheric OH. The most important such

gases are CO and nitrogen oxides. Lifetimes of many

ozone-depleting substances are also determined by photolysis rates in the stratosphere. Hence, lifetimes of these

gases can depend on the intensity of the Brewer-Dobson

circulation as well. Changes in the oxidizing power of the

atmosphere could have large impacts on air pollution,

aerosol formation, greenhouse radiative forcing, and stratospheric ozone depletion. The quantification of possible

recent changes in tropospheric OH continues to be an

active and controversial area of research. The evidence

for such changes is based primarily on inverse models of

CH3CCl3 emissions and tropospheric observations that

have deduced OH variations on the order of ±10% in recent

decades but no systematic long-term trends (Bousquet et

al., 2005; Prinn et al., 2005). In the future, the Intergovernmental Panel on Climate Change/Technology and Economic Assessment Panel (IPCC/TEAP Chapter 2; Velders

and Madronich et al., 2005) concludes that tropospheric

OH in the 21st century may change by −18% to +5%,

depending on the choices of emission scenarios.

The atmospheric lifetimes for the vast majority of

the gases listed in Table 1-4 are unchanged from WMO

2002. The list of gases and lifetimes in this Table is not

limited only to gases with atmospheric lifetimes longer

than 6 months, because it was decided to put all the lifetimes needed for Chapters 1 and 2 of this Assessment in

a single table. This section describes the basis for changes

that were made either due to new laboratory experiments

or due to including additional gases. The best estimate

lifetime for methyl chloride (CH3Cl) was reduced due to

the discovery of additional sinks. Analysis by Tokarczyk

et al. (2003a, b) of the partial atmospheric lifetime with

respect to ocean removal for CH3Cl (4.1 years) reduces

the total lifetime to 1.0 years.

The atmospheric lifetime for methyl chloroform has

been left at 5.0 years, the same as in WMO 2002 (Montzka

and Fraser et al., 2003), because the suggestion by Prinn

et al. (2005) for a lifetime of 4.9 years and a tropospheric

lifetime due to reaction with OH of 6.0 years (6.1 years

used in this Assessment) was not deemed to represent a

significant change.

Although the atmospheric lifetime of the hydrofluoroether CHF2OCF3 (HFE-125) is unchanged at 136 years,

the much longer atmospheric lifetime (1800 years) implied

by the theoretical analysis of Wu et al. (2004) suggests

that further study is needed for this compound.

The atmospheric lifetime of trifluoromethylsulfurpentafluoride (SF5CF3) was modified to a range of 650-

950 years to better account for the findings in Takahashi

et al. (2002). It did not seem warranted to use a single lifetime of 800 years. However, it should be noted that laboratory data by Chim et al. (2003) and Limão-Vieira et al.

(2004) both imply an atmospheric lifetime closer to 1000

years for SF5CF3.

A number of additional gases have been included in

Table 1-4 based on new laboratory data for their reaction rates

with OH. These include: CFH2CH2OH, CHF2CF2OCHF2,

CF3CHFCF2OCH2CF2CF3, CF3OC(O)H, C2F5OC(O)H,

CF3CHFCF2CH2OH, C2F5C(O)CF(CF3)2, and nC3F7OC(O)H. Some other compounds were not included because only room temperature reaction data were

available.

Also included are several compounds that were otherwise included in the Global Warming Potential (GWP)

analysis in WMO 2002: c-C3F6 (τ = 1000 years), NF3 (740

years), (CF3)2CHOCHF2 (3.1 years), and (CF2)4CH(OH)

(0.85 year).

1.2.2.2 UNCERTAINTIES IN OZONE-DESTROYING

HALOCARBON LIFETIMES

In order to estimate the remaining banks (and their

uncertainties) of CFCs and other ozone destroying halocarbons, it is useful to have the best possible estimates of

the uncertainties in their lifetimes. For this purpose, the

model-based stratospheric lifetime estimates from the

1998 Assessment (Table 1-4 in WMO, 1999) are combined

with values derived from stratospheric measurements

LONG-LIVED COMPOUNDS

1.22
LONG-LIVED COMPOUNDS

1.23

Table 1-4. Trace gas lifetimes for selected halocarbons and perfluorinated gases. For completeness,

estimates for local lifetimes for some very short-lived (lifetime (τ) < 0.5 year) species are included (in italics).

As discussed in Chapter 2, the atmospheric lifetimes for these species (defined as the ratio of burden to

emission) (Prather and Ehhalt et al., 2001; Montzka et al., 2003) depend on the location and time of emission. Thus, these local lifetimes should not be used in semi-empirical ODP, GWP, or EESC calculations for

these gases.

Industrial Designation or Chemical Formula Lifetime Notes

Common Name (years)

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 1.0 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, 19

Bromodifluoromethane CHBrF2 5.8 4

Bromochloromethane CH2BrCl 0.41 8, 15, 16

Bromodichloromethane CHBrCl2 0.19 7, 15, 16

Dibromochloromethane CHBr2Cl 0.21 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, 20

Diiodomethane CH2I2 Minutes 7, 15, 16

Chloroiodomethane CH2ClI Hours 7, 15, 16

Trifluoroiodomethane CF3I 0.01 2, 15

Halogen-substituted ethenes

Trichloroethene C2HCl3 0.01 15

Perchloroethene C2Cl4 0.27 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
LONG-LIVED COMPOUNDS

1.24

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 0.08 15

1,2 Dichloroethane CH2ClCH2Cl 0.19 10, 15

Methyl chloroform CH3CCl3 5.0 9

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

Iodoethane C2H5I 0.01 7, 15, 16, 20

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 CH3CF2CCl2F 26.4 4

HCFC-225ca CHCl2CF2CF3 1.9 11

HCFC-225cb CHClFCF2CClF2 5.8 11

n-Propyl bromide (n-PB) CH3CH2CH2Br 0.03 15, 16

n-Propyl iodide CH3CH2CH2I 0.001 7, 15, 16, 20

Isopropyl iodide CH3CHICH3 0.003 7, 15, 20

Halogen-substituted higher alkanes

HFC-365mfc CH3CF2CH2CF3 8.6 1, 14

HFC-356mcf CH2FCH2CF2CF3 1.2 1, 14

HFC-356mff CF3CH2CH2CF3 8.1 1, 14

HFC-338pcc CHF2CF2CF2CHF2 12.3 1, 14

Table 1-4, continued.

Industrial Designation or Chemical Formula Lifetime Notes

Common Name (years)
LONG-LIVED COMPOUNDS

1.25

FC-C318 (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

Perfluorodecalin C10F18 2000 29

HFC-55-10mcff CF3CF2CH2CH2CF2CF3 7.7 2

FC-51-14 (Perfluorohexane) C6F14 3200 2

Perfluorocyclopropane c-C3F6 >1000 2

Fluorinated alcohols and ketones

CF3CH2OH 0.41 15

CF3CF2CH2OH 0.39 15, 23

(CF3)2CHOH 2.0 4

CF3(CF2)3CH2OH 0.45 23

CFH2CH2OH 0.04 27

CF3CHFCF2CH2OH 0.34 22

C2F5C(O)CF(CF3)2 0.04 21

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

HCFE-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 2.0 2

HFE-365mcf3 CH3OCH2CF2CF3 0.11 2, 15

Table 1-4, continued.

Industrial Designation or Chemical Formula Lifetime Notes

Common Name (years)
LONG-LIVED COMPOUNDS

1.26

HFE-374pc2 a CH3CH2OCF2CHF2 5.0 2

CHF2CF2OCHF2 26.5 24

CF3CHFCF2OCH2CF2CF3 9.1 24

(CF3)2CHOCHF2 3.1 2

CF3CH(OCF3)CHF2 3.1 2

(CF3)2CFOCH3 3.4 2

CF3OC(O)H 3.6 25

C2F5OC(O)H 3.6 26

n-C3F7OC(O)H 2.6 26

HFE-7100 (HFE-44-9) CH3OC4F9 5.0 2

HFE-7200 (HFE-56-9) C2H5OC4F9 0.77 2

HFE-43-10pccc124 b CHF2OCF2OCF2CF2OCHF2 6.3 2

HFE-236ca12 (HG-10) CHF2OCF2OCHF2 12.1 2

HFE-338pcc13 (HG-01) CHF2OCF2CF2OCHF2 6.2 2

Other fluorinated species

Trifluoromethylsulfurpentafluoride SF5CF3 650-950 18

Sulfur hexafluoride SF6 3200 2

Nitrogen trifluoride NF3 740 2

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. (2006); and stratospheric loss lifetimes inferred from IPCC (2001).

2. Prather and Ehhalt et al. (2001) and Ramaswamy et al. (2001).

3. See Section 1.2.1.6 for further discussion related to methyl halide global lifetimes. Includes estimate of ocean sink from Tokarczyk et al. (2003a, b).

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. Based on WMO 2002, including oceanic loss term with 94-year partial lifetime observed in saturation data and ascribed to an unidentified process

(Yvon-Lewis and Butler, 2002) but also accounting for soils sink with 90-year partial lifetime (Happell and Roche, 2003). However, see Section

1.3.2.4 for additional considerations based on atmospheric measurements.

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 Qiu 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 from Chapter 2 (this Assessment). 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 (τ < 0.5 year) 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), Chim et al. (2003), and Limão-Vieira et al. (2004). Range in lifetime depends on whether solar proton events reduce the lifetime.

19. Bayes et al. (2003).

Table 1-4, continued.

Industrial Designation or Chemical Formula Lifetime Notes

Common Name (years)
by Volk et al. (1997, Table 5). This provides lifetime

uncertainties that are almost entirely independent of the

estimates that can be produced by using the atmospheric

measurements from the ground-based networks combined

with emissions estimated from industrial production figures. The uncertainties are obtained by taking the root

mean square sum of the variances from the model-based

estimates and the variances given by Volk et al. The results

are lifetime uncertainty ranges of 35-57 years for CFC-11,

79-113 years for CFC-12, and 66-114 years for CFC-113.

However, it was indicated in Section 1.2.1.1 that CFC-113

is decreasing in the atmosphere at a rate of 1%/yr, which

implies a maximum lifetime of approximately 100 years.

A range of 66-100 years is then derived for CFC-113. The

stratospheric lifetime uncertainty range for CCl4 calculated in the same way is 28-46 years. It may be noted that

the recommended lifetimes in the previous section do not

lie exactly in the middle of these uncertainty ranges. That

is because, except where new measurements had been

made, there was no justification for changing the values

used in WMO 2002.

For the major HCFCs, the lifetime uncertainty

ranges are obtained by combining the uncertainty in the

hydroxyl value of 14% given by Prinn et al. (2001) with

the uncertainties in the reaction rates between the HCFCs

and hydroxyl from Sander et al. (2006). This gives lifetime uncertainties of ±17% for HCFC-22, ±20% for

HCFC-141b, and ±24% for HCFC-142b.

1.2.3 Total Atmospheric Chlorine and

Bromine

1.2.3.1 TOTAL ORGANIC CHLORINE IN THE

TROPOSPHERE

Total organic chlorine (CCly) from long-lived

source gases continued to slowly decrease in the troposphere during 2001-2004 (Montzka et al., 2003;

O’Doherty et al., 2004) (Figure 1-9 and Table 1-5). In

mid-2004 tropospheric CCly was approximately 3.44 parts

per billion (ppb), or about 0.25 ppb Cl (nearly 7%), below

the peak observed in 1992-1994. The rate of decline in

CCly from 2003 to 2004 (−20 ppt Cl/yr or −0.59%/yr) was

slightly less than in previous years (−23 ppt Cl/yr or

−0.64%/yr), owing to the reduced influence of methyl

chloroform. In 2004, mixing ratio changes for methyl

chloroform still contributed substantially to the decrease

in CCly (−13.5 ppt Cl/yr), though this contribution was

nearly matched by the decline observed in the aggregate

of CFCs plus CCl4 (−12.7 ppt Cl/yr). In 2004, the CFCs

accounted for 62% of CCly in long-lived gases in the lower

atmosphere; this fraction has increased slightly over the

past decade as Cl from shorter-lived gases has diminished

more rapidly than Cl from CFCs.

Increases in CCly from HCFCs slowed during 2000-

2004. CCly from the three major HCFCs increased at a

rate of nearly 6 ppt Cl/yr during 2004, or almost half the

rate of annual increase measured in 1996 (Table 1-5).

HCFCs accounted for 214 ppt of CCly in mid-2004, or

about 6%.

1.2.3.2 TOTAL CHLORINE IN THE STRATOSPHERE AND

HYDROGEN CHLORIDE (HCl)

Stratospheric chlorine time series are different from

tropospheric time series in many respects. Stratospheric

chlorine levels lag tropospheric levels, due to the time it

takes for the trends to propagate. During the propagation

of tropospheric trends into the stratosphere, mixing leads

to a distribution of transit times in any air parcel in the

stratosphere, which is described by the concept of “age of

the air.” This mixing will inevitably make long-term

trends in stratospheric time series less sharp, if they are

driven by tropospheric trends (see, e.g., Waugh et al., 2001,

and Engel et al., 2002). The peak in the tropospheric chlorine is thus expected to occur later in the stratosphere and

be broader and have somewhat lower mixing ratios in its

LONG-LIVED COMPOUNDS

1.27

20. Cotter et al. (2003).

21. Taniguchi et al. (2003).

22. Chen et al. (2003).

23. Hurley et al. (2004) estimate the lifetimes of F(CF2)nCH2OH, for n = 1-4, to be 0.45 year.

24. Chen et al. (2005).

25. Chen et al. (2004a).

26. Chen et al. (2004b).

27. Rajakumar et al. (2005).

28. Li et al. (2006).

29. Shine et al. (2005).

a Referred to as HFE-374pcf2 in past Assessments. b Referred to as H-Galden 1040x in past Assessments.

Table 1-4, continued.
LONG-LIVED COMPOUNDS

1.28

Table 1-5. Contributions of halocarbons to total chlorine (Cl) in the troposphere. An average of AGAGE

and NOAA/ESRL global means was used to derive these mid-year mixing ratios. Rates of change were calculated as the difference between the indicated year and the previous year. Uncertainties in rates of change (1

standard deviation) are estimated from the difference in rates derived from observations by the different global

networks, NOAA/ESRL and AGAGE. The uncertainty in total Cl is approximately 25 ppt Cl (1 standard deviation) throughout the period 1996-2004 (based on the sum of differences in global tropospheric means derived

for individual gases between the AGAGE and NOAA/ESRL networks).

Contribution to Rate of Change in Total Cl 

Total Cl (ppt Cl) Total Cl (%) (ppt Cl/yr)

1996 2000 2004 1996 2000 2004 1996 2000 2004

Aggregated CFCs 2160 2156 2126 59% 61% 62% 5.8 (3.4) −1.4 (2.0) −8.9 (0.4)

CH3Cl a 550 550 550 15% 16% 16% 0.0 b 0.0 b 0.0 b

CCl4 407 393 378 11% 11% 11% −4.2 (0.6) −3.6 (0.3) −3.8 (0.7)

Aggregated HCFCs 141 180 214 4% 5% 6% 10.3 (0.5) 9.2 (0.8) 5.9 (1.0)

CH3CCl3 274 137 66 8% 4% 2% −39.7 (0.2) −26.8 (1.0) −13.5 (0.4)

CH2Cl2, CHCl3, etc. c 100 100 100 3% 3% 3% 0.0 b 0.0 b 0.0 b

Halon-1211 3.5 4.0 4.2 0% 0% 0% 0.2 0.1 0.0

Total Cl 3636 3521 3437 −28 (3.5) −23 (2.4) −20 (1.4)

−0.76% −0.64% −0.59%

a CH3Cl assumed to be 550, consistent with the previous Assessment (Montzka and Fraser et al., 2003). b Growth rate assumed to be zero here, though some very short-lived gases have been declining over the past decade (see Chapter 2). c CH2Cl2, CHCl3, etc., also includes the very short-lived gases C2Cl4, C2HCl3, and CH2ClCH2Cl, and an additional 50 ppt to account for Cl from COCl2

(see Chapter 2 for additional discussion of the contribution of short-lived gases to total Cl).

2200

2400

2600

2800

3000

3200

3400

3600

3800

1975 1980 1985 1990 1995 2000 2005

Year

Mixing ratio (ppt)

-2.0

-1.0

0.0

1.0

2.0

3.0

4.0

5.0

6.0

Rate of change (%)

Scenario Ab, WMO '02

+ 100 ppt Cl 

AGAGE

NOAA

AGAGE, rate

NOAA, rate

Figure 1-9. The tropospheric abundance of organic chlorine (CCly) from

the NOAA/ESRL and AGAGE global

measurement networks (updates of

Montzka et al., 1999, and O’Doherty

et al., 2004) and its rate of change over

time. Quantities are based upon independently measured mixing ratios of

CFC-11, CFC-12, CFC-113, HCFC22, HCFC-141b, HCFC-142b, methyl

chloroform, and carbon tetrachloride.

Results for CFC-114 and CFC-115

from Prinn et al. (2000) are used in

both compilations. An additional constant 550 ppt was added for CH3Cl

and 100 ppt was added for short-lived

gases such as CH2Cl2, CHCl3, C2Cl4,

and COCl2. Rates of change are

determined from 12-month differences and are plotted with respect to

the right-hand axis. Observations are

compared with the projections from

WMO 2002 (green line).
maximum (about 150 ppt lower; see Engel et al., 2002).

Figure 1-10 shows an update of stratospheric chlorine

trends for midlatitudes (44°N) calculated by Engel et al.

(2002), based on tropospheric observations by NOAA/-

ESRL up to the year 2005, mean profiles of mean age, and

a parameterization of the width of the age spectrum. Based

on these calculations, the maximum chlorine levels were

expected in the lower stratosphere around 1995 and were

expected to have reached altitudes above 30 km by about

2000. It is expected that total chlorine is now decreasing

at all altitudes.

Total chlorine in the stratosphere could either be in

the form of organic chlorine or inorganic chlorine. Figure

1-11 is constructed based on the profiles of organic and

inorganic species derived from observations and modeling

results (see figure caption for details). Nassar et al.

(2006a) derived the total chlorine budget for 30°N-60°N,

based mainly on observations by the ACE-FTS, yielding

3.65 ppb chlorine (with an estimated accuracy of ±0.13

ppb) in 2004. This value is slightly larger than the value

of 3.54 ppb (with a precision of ±0.10 ppb) derived by

Zander et al. (1996) for 1994 from ATMOS experiment

observations, when stratospheric chlorine levels were still

increasing.

At high altitudes (above 50 km) the total chlorine

budget is dominated by hydrogen chloride (HCl), which

at this altitude may serve as a good proxy for total inorganic chlorine (Cly) and even total chlorine. The longest

continuous time series of HCl from satellite observations

is that of the Halogen Occultation Experiment (HALOE)

(Russell et al., 1993; Anderson et al., 2000). For about 55

km altitude, Figure 1-12 compares the HCl measurements

with model values of HCl that are derived from a total

chlorine time series that has been convoluted with spectra

of various widths, each having a mean age of 5.5 years.

The total chlorine values are based on the NOAA/ESRL

and AGAGE measurements including a 100 ppt contribution from very short-lived (VSL) source gases (full lines),

whereas the dashed curve (from scenario Ab in WMO

2002) excludes a contribution from VSL compounds.

These Cly time series have been reduced by a factor of

0.96, which represents model estimates of the HCl/Cly

ratio at 55 km, in order to arrive at the model curves for

HCl shown in Figure 1-12.

LONG-LIVED COMPOUNDS

1.29

Total chlorine evolution in the midlatitude stratosphere

Figure 1-10. Total chlorine calculated based on the vertical distribution of mean age (derived from observations of the age tracers sulfur hexafluoride (SF6) and carbon dioxide (CO2)) and assumptions on the width of

the age spectrum. Tropospheric data by NOAA/ESRL are used as the basis of the calculation (updated from

Engel et al., 2002). No chlorine contribution from very short-lived (VSL) species is included in this calculation.
The linear trend in HALOE values from January

1999 to November 2005 is −26 ± 12 ppt/yr (one sigma),

which is larger than, but not significantly different from,

the trend in the total chlorine based on the organic record.

The time of the peak in HCl is somewhat ill defined in the

HALOE data. Rinsland et al. (2005) showed that the HCl

change observed between ATMOS observations in 1985

and 1994 and ACE observations in 2004 are broadly consistent with the expectations due to the changes in total

stratospheric chlorine (Figure 1-12). Overall, the time

series of the lower mesospheric HCl (Cly) are qualitatively

consistent with expectations from ground-based measurements after considering lag times (~5-6 years) (Waugh et

al., 2001; Waugh and Hall, 2002) associated with transport of tropospheric air to the lower mesosphere, although

the fine-scale structure observed between 1997 and 2002

in the HALOE data remains unexplained.

There are considerable differences in the amount of

upper stratospheric HCl measured by the four instruments.

The systematic difference in HCl observations between

MLS and HALOE is about 0.2-0.4 ppb (about 7-15%,

HALOE lower), while ACE and MLS show better agreement (Froidevaux et al., 2006). Measurements of HCl by

MLS and ACE both show values between 3.4 and 3.6 ppb

in 2006 (Figure 1-12), which would imply a peak value

for total chlorine of ~3.5 to 3.8 ppb, larger than can be supplied by the 12 long-lived source gases. The HALOE

measurements of HCl are consistent with supply of chlorine from only long-lived source gases. The differences

between the instruments are not currently understood, but

it should be noted that the different measurements do agree

within the stated accuracies, indicated by the error bars

(±2 sigma) in Figure 1-12. Due to these differences in

absolute values of these observations, it is at present not

possible to state whether upper stratospheric inorganic

chlorine is consistent with supply from only the 12 longlived source gases, or it is large enough to allow for a chlorine contribution from short-lived gases of 0.05 to 0.1 ppb.

This issue is discussed further in Section 2.5.2 of Chapter

2, which focuses on possible contributions from the very

short-lived substances.

As is evident from Figure 1-11, column observations of HCl and chlorine nitrate (ClONO2) can also be

used to gain insight regarding the total inorganic chlorine

in the stratosphere. Recent high spectral resolution

infrared ground-based observations have also provided

evidence that the total loading of stratospheric chlorine

has reached its maximum, with tentative evidence, at the

1-sigma level, of the onset of a long-term decline. For

example, observations from the Network for the Detection

of Atmospheric Composition Change (NDACC) from the

3 stations with measurements spanning up to 24 years have

LONG-LIVED COMPOUNDS

Figure 1-11. The northern midlatitude (30°N-60°N)

stratospheric chlorine budget for 2004 expressed in

volume mixing ratio (VMR) (Nassar et al., 2006a). It

includes ACE-FTS measurements of HCl, chlorine

nitrate (ClONO2), chlorofluorocarbonyl (COClF),

CH3Cl, CCl4, CFC-11, CFC-12, CFC-113, HCFC-22,

and HCFC-142b, which are complemented by chlorine monoxide (ClO) observations from the SubMillimeter Receiver (SMR) on the Odin satellite

(Berthet et al., 2005), hypochlorous acid (HOCl)

measurements by the Michelson Interferometer for

Passive Atmospheric Sounding (MIPAS) (von

Clarmann et al., 2006), and phosgene (carbonyl

chloride, COCl2) observations from the balloonborne MkIV (Toon et al., 2001). ACE-FTS measurements of HCFC-22 were extended above 25 km

based on values from the Stratospheric Chemistry

Transport Model-1 (Rummukainen, 1996; Coheur et

al., 2003). CH3CCl3 and the minor contributors CFC114, CFC-114a, CFC-115, HCFC-141b, and halon1211 were based on SAGE III Ozone Loss and

Validation Experiment (SOLVE) measured profiles

(Schauffler et al., 2003) scaled to 2004. The sum of

the chlorine contributions from these species gives

the total chlorine profile, with error bars indicating

the 1-sigma precision.

1.30
provided evidence for stabilization in total chlorine

loading, with a broad peak starting at the end of the 1990s,

following a rapid buildup during the 1980s (Rinsland et

al., 2003). Figure 1-13 provides an update to the NDACC

time series from the Jungfraujoch station measurements

extending through 2004 from monthly mean time series.

It indicates that the total stratospheric inorganic chlorine

loading (Cly) has decreased slowly (−0.7 ± 0.3%/yr, 1σ)

since it peaked in late 1996, at the limit of being statistically significant at the 2σ level (an update of Mahieu et

al., 2004). Multiplying this percentage by a stratospheric

Cly burden of 3.6 ppb suggests a trend in Cly of −25 ± 11

ppt/yr. The column measurements are in good agreement

with model calculations by a two-dimensional (2-D) model

run at the University of Leeds (orange trace in Figure 1-

13, described in Rinsland et al., 2003).

Solomon et al. (2006) have reported a steeper downtrend for stratospheric chlorine monoxide (ClO) from 1996

to 2005 of −1.5 ± 0.2%/yr, based on remotely sensed millimeter wave measurements. They calculate however

that this downtrend contains an enhancement of about

−0.5%/yr, relative to Cly trends, because of increases in

upper stratospheric methane that have been measured by

HALOE. Their observations, which extend from 1984 to

present, are broadly consistent with the rise in total chlorine due to human activities and recent declines due to the

Montreal Protocol.

LONG-LIVED COMPOUNDS

1.31

1985 1990 1995 2000 2005

2.5

3.0

3.5

4.0

HCl / ppbv

Spectrum width / yr

0

2

4

MLS

HALOE

ATMOS

ACE

Mean Age = 5.5 yr

Figure 1-12. HCl monthly averages for 60°S-60°N at 0.46 hPa from MLS (red dots; Froidevaux et al., 2006

updated) and HALOE (purple circles; Anderson et al., 2000 updated), along with quarterly averages from ACE

(in green; Nassar et al., 2006a updated) and ATMOS version 3 data averages (cyan squares; Irion et al., 2002;

Rinsland et al., 2005) for 50-60 km; error bars are 2-sigma estimates of accuracy. Solid curves use total chlorine time series from tropospheric data (including 100 ppt from very short-lived (VSL) chlorinated substances)

and a mean age assumption of 5.5 years, with various age spectrum widths, as indicated; a factor of 0.96 is

applied to get HCl. The dashed black curve shows expected HCl based on the WMO 2002 scenario Ab (no

contributions from VSL compounds) and a mean age of 5.5 years, with a spectrum width of 2 years.
1.2.3.3 ORGANIC BROMINE IN THE TROPOSPHERE

FROM METHYL BROMIDE AND HALONS

Because bromine is an efficient catalyst for depleting stratospheric ozone, the Montreal Protocol set limits

on anthropogenic production and consumption of the most

abundant and long-lived brominated gases in the atmosphere: methyl bromide and the halons. Increased anthropogenic production and use of these gases led to increasing

amounts of bromine (Br) in the atmosphere during the

second half of the 20th century. Enhanced human-related

emissions of methyl bromide and halons likely dominated

changes in atmospheric bromine concentrations over the

past 50 years. By the late 1990s, about half of the Br in

the stratosphere was derived from anthropogenic activities (Montzka and Fraser et al., 2003).

Natural processes deliver bromine to the stratosphere

in the form of methyl bromide and short-lived bromocarbons or oxidation products of these gases. Long-term

changes in atmospheric Br arising from natural processes

likely were smaller than the changes in anthropogenicallyderived Br during the 20th century. The background level

of stratospheric Br accounted for by short-lived bromocarbons is an area of active research (Chapter 2).

LONG-LIVED COMPOUNDS

1.32

Figure 1-13. Top: Time series of monthly mean total column HCl (red circles) and ClONO2 (green triangles),

as derived from the Jungfraujoch (46.5°N) observational database; Cly (blue triangles) results from the summation of the corresponding HCl and ClONO2 data points. To avoid the influence of the significant variability

affecting measurements in winter- and springtime, these time series have been limited to the June to November

months. Nonparametric least-square fits to these datasets are reproduced as thick black curves; they help to

appraise the two-decade mean evolution of HCl, ClONO2, and Cly. Also shown is the sum of the HCl and

ClONO2 columns calculated by the University of Leeds 2-D model (top orange curve). Bottom: Evolution of

total surface organic chlorine derived from measurements performed by the AGAGE network. For better comparison with the Jungfraujoch data, the tropospheric chlorine trend is also shown with a 3.5-year time shift

(magenta line). Updated from Mahieu et al. (2004).
New measurements of methyl bromide have been

reported (see Section 1.2.1.7) since the previous Ozone

Assessment (WMO, 2003) and they allow trends in total

bromine from regulated gases to be assessed. The results

(Figure 1-14) indicate that the global tropospheric burden

of Br from the sum of methyl bromide and halons peaked

in 1998 and has declined since (Montzka et al., 2003).

By mid-2004, tropospheric Br was 0.6 to 0.9 ppt below

the peak; the mean decline from 1998 to 2004 was

between −0.1 and −0.15 ppt Br/yr (a range is given here

to reflect differences in recent trends for halon data from

different networks and the possibility that mean tropospheric CH3Br mixing ratios trends are 7% less than

observed at Earth’s surface). The decline is solely the

result of a decrease in global mixing ratios of methyl bromide, which declined at a mean rate of −0.24 ppt/yr from

1999 to 2004, or somewhat faster than expected. Mixing

ratios of halon-1301 and halon-1211 still increased slowly

from 1999 to 2004; the mean increase in Br from these

gases over this period was between 0.08 and 0.12 ppt/yr.

Continued increases in the halons are consistent with the

existence of substantial banks of halons that are released

at a slow rate to the atmosphere.

The decline observed for bromine from regulated gases

is significant relative to changes in chlorine, considering that

the efficiency for Br to deplete stratospheric ozone is 60

times greater than Cl on a per-atom basis (see Chapter 2 and

Chapter 8 for an updated discussion of this efficiency factor,

alpha). A decline of −0.1 to −0.15 ppt Br/yr, multiplied by

60, corresponds to a rate of −6 to −9 ppt Cl equivalents/yr, or

more than one-third of the rate of decline in atmospheric chlorine over this same period (−20 ppt Cl/yr; Table 1-5).

As was the case for Cl, initial declines in atmospheric Br have been achieved through reduced production

and emission of a short-lived gas even while the long-lived

gases continued to slowly increase.

1.2.4 Effective Equivalent Chlorine and

Equivalent Effective Stratospheric

Chlorine

Effective equivalent chlorine (EECl) and equivalent

effective stratospheric chlorine (EESC) provide a rough

means of gauging the influence of measured changes in

both chlorinated and brominated halocarbons on the ozonedepleting halogen abundance in the future stratosphere (see

LONG-LIVED COMPOUNDS

1.33

5.0

7.0

9.0

11.0

13.0

15.0

17.0

19.0

1960 1970 1980 1990 2000

Year

Bromine at Earth s surface (ppt)

Ab, WMO 2002

UEA

NOAA/ESRL

AGAGE

CH3Br only

Figure 1-14. Trends in the abundance of Br from brominated

gases regulated by the Montreal

Protocol on Substances that

Deplete the Ozone Layer (halon1211, halon-1301, halon-2402

and methyl bromide) as projected

in the WMO 2002 Ab scenario

(black line) and as measured by

different stations (red, green, and

blue lines). Although independent global surface means

are derived for halon-1211 and

halon-1301 by different laboratories, the same is not true for

methyl bromide and halon-2402.

For these gases, the source of

data included in the compilation

of total Br for all labs was solely

NOAA/ESRL (methyl bromide)

and UEA (halon-2402). The

global surface estimates are

made with samples collected

from single sites (AGAGE prior to

the mid-1990s and UEA) or multiple sites (AGAGE after the mid1990s and NOAA/ESRL).
descriptive box on EECl and EESC in Chapter 8). EECl

and EESC are derived from ground-based observations of

trace gas mixing ratios. As currently formulated, EECl and

EESC do not include contributions from very short-lived

halogen source gases. Weighting factors are applied to

measured mixing ratios to account for the variable rates at

which these gases undergo degradation in the stratosphere

and liberate halogen atoms for interaction with ozone.

Mixing ratios of bromine are multiplied by a factor alpha

(α), which reflects the enhanced per-atom efficiency for

bromine to catalyze the destruction of stratospheric ozone,

relative to chlorine (60 is used here). Finally, a lag time is

applied in estimating EESC to account for the time it takes

air to be transported from the surface to regions in the stratosphere. A lag of 3 years is used to approximate transport

times and halogen burdens in the midlatitude lower stratosphere (Daniel et al., 1995); longer lag times are more

appropriate for the polar stratosphere and the upper stratosphere (Andrews et al., 2001; Newman et al., 2006). No

time lag is applied when deriving EECl, so that dates quoted

for EECl correspond to sampling dates. Further details on

uncertainties and limitations of these quantities are discussed in Chapter 8 of this Assessment, as well as in

Chapter 1 of the previous Assessment (WMO, 2003).

Decreases in EECl and EESC continued during

2000-2004, owing to continued declines in tropospheric

mixing ratios of most regulated ozone-depleting source

gases (Figure 1-15). By 2005, EECl was 277 ppt, or 8-

9% below the peak observed in the 1992-1994 time

period. This decline represents about 20% of the decline

LONG-LIVED COMPOUNDS

1.34

2800

2900

3000

3100

3200

1990 1995 2000 2005 2010

Year

EECl (ppt)

Ab Scenario (WMO, 2003)

AGAGE + NOAA Halons + CH Br from Ab (WMO, 2003)

NOAA + CH Br from Ab (WMO, 2003)

NOAA + Observed CH Br

Figure 1-15. EECl calculated from observations of the most abundant long-lived chlorinated gases (CFC-11,

-12, -113, CH3CCl3, CCl4, HCFC-22, -141b, and -142b) and the longest-lived brominated gases (halon-1211

and halon-1301). Results for CH3Br are also included either as projected in the WMO 2002 Ab scenario (WMO,

2003) (in data represented by blue, red, and green symbols and lines) or as observed in the NOAA/ESRL network along with observational data for all other gases (black points and line). Results from global measurements programs AGAGE (blue diamonds and line; O’Doherty et al., 2004) and NOAA/ESRL (red diamonds

and line; Montzka et al., 2003) are compared with scenario Ab from the 2002 WMO Assessment report (Montzka

and Fraser et al., 2003) (green line). In the calculations presented here, all bromine atoms are multiplied by a

factor of 60, owing to evidence suggesting that Br atoms are 60 times more effective at destroying stratospheric ozone (see discussion in Chapter 8). Fractional release factors used here are taken from Table 8-1,

and all release factors have been normalized to 0.8 for CFC-11, despite reports that absolute release factors

may be substantially smaller for the middle stratosphere (Schauffler et al., 2003). Using a smaller absolute

release factor would affect all absolute mixing ratios and growth rates reported here, but would not affect relative quantities or integrated EESC calculated for future scenarios in Chapter 8.
needed for atmospheric EESC to return to 1980 levels

(Figure 1-16). The mean rate of decline during 2000-

2004 was similar to that observed during 1996-2000 but

was faster than projected in scenario Ab of WMO 2002

(Montzka and Fraser et al., 2003). This discrepancy

arises primarily from a more rapid decline in CH3Br,

though slower-than-projected increases in measured

mixing ratios of HCFCs and halon-1301, and a more

rapid decline than projected for CH3CCl3, also contributed (Figure 1-15). Note that discussions related to

trends in observationally derived EECl and EESC did

not include CH3Br in the recent WMO 1998 and 2002

reports, because of a lack of CH3Br measurements for

recent years that have since become available. Declines

in shorter-lived gases (CH3Br and CH3CCl3) continue to

be the largest contributors to declining EECl. The early

removal of shorter-lived gases means that later decreases

in EECl and EESC will likely be dominated by the

atmospheric decay of the longer-lived gases. Despite

diminished mixing ratios of CH3CCl3, this gas remained

the single largest contributor to decreasing EECl during

2000-2004 (−61 ppt EECl out of a decline from all gases

of −120 ppt EECl over this period). Declines in CH3Br

accounted for about one-third of the overall decline

(−44 ppt EECl); the CFCs together accounted for less of

the overall decline (−23 ppt EECl) as did CCl4 (−13 ppt

EECl). These decreases were offset somewhat by continued increases from HCFCs (+12 ppt EECl) and halons

(+11 ppt EECl).

Calculations of EECl discussed here include the

assumption that mixing ratios measured at the Earth’s surface provide an accurate estimate of the amount of Cl and

Br from halocarbons entering the stratosphere. For

source gases whose tropospheric loss rates are similar to

rates of vertical mixing within the troposphere, surfacebased observations may overestimate the amount and

trend of Cl and Br reaching the stratosphere. Such influences may affect CH3Br more than other gases considered in the calculation of EECl here, because it has the

shortest lifetime with respect to atmospheric photochemical loss in the troposphere (~1.7 years). Vertical gradients on the order of 0-15% have been observed for CH3Br

between the lower and upper troposphere (Schauffler et

al., 1999).

1.2.5 Fluorine in the Stratosphere

1.2.5.1 TOTAL STRATOSPHERIC FLUORINE AND

HYDROGEN FLUORIDE (HF) AT 55 KM

The majority of the fluorine in today’s atmosphere

is derived from anthropogenic activities, as there are no

known significant natural fluorine sources (Zander et al.,

1994). Stratospheric fluorine is unimportant for stratospheric ozone depletion in comparison to chlorine and

bromine, but, because it is mainly due to CFCs, HCFCs,

and HFCs, it provides an important consistency check on

the anthropogenic chlorine input. An inventory of total

fluorine (FTOT) in the stratosphere, including both organic

and inorganic species, from 82°N to 82°S over the period

of February 2004-January 2005 inclusive has been

reported by Nassar et al. (2006b), based primarily on ACEFTS measurements. Stratospheric FTOT values were 2.50-

2.59 ppb at all latitudes with a precision of ±0.04-0.07 ppb.

The highest value of mean FTOT occurred in the tropics,

which is qualitatively consistent with increasing stratospheric fluorine and the mean stratospheric circulation pattern. Figure 1-17 shows the stratospheric FTOT budget for

30°N-60°N in 2004.

The main source of stratospheric HF is photodissociation of chlorofluorocarbons (CFCs), and it is therefore

an effective tracer of fluorine input to the stratosphere. As

altitude increases, the organic fluorine compounds decomLONG-LIVED COMPOUNDS

1.35

1000

1500

2000

2500

3000

3500

1975 2000 2025 2050

EESC (ppt)

Observations

WMO 2002, Ab scenario

Past Trend

1980 level

Year in stratosphere (sampling date +3yr)

Figure 1-16. Observed and projected changes in

EESC relative to the 1980 level. The 1980 level is

chosen to represent the atmospheric EESC burden

before the ozone hole was observed. Here, EESC is

derived from observations of all gases included in the

WMO 2002 Ab scenario (Montzka and Fraser et al.,

2003). The observed decline in EESC of 277 ppt

EESC from its peak in the 1990s is 20% of the difference between the peak and 1980 EESC levels.

Dates are meant to be relevant for the midlatitude

stratosphere and are derived by adding 3 years to

the date at which samples were collected at the surface (to account for the time it takes for air at Earth’s

surface to reach the midlatitude stratosphere).
pose due to photolysis and chemistry, forming inorganic

compounds. At 55 km altitude, HF comprises ~93% of

the total inorganic fluorine (Fy) at high latitudes and ~86%

near the equator, based on National Center for Atmospheric Research (NCAR) two-dimensional model results

(Brasseur et al., 1990). The HALOE instrument has provided temporal and spatial information on the near-global

distribution of HF since October 1991. The accuracy for

a single HF profile at 55 km is ~14% and the precision is

better than 0.05 ppb (Russell et al., 1996). Figure 1-18

shows the time series of HALOE-observed near-global

HF, HALOE-derived Fy using latitude-dependent HF/Fy

ratios between 86 and 93%, (updated from Anderson et

al., 2000), and the WMO 2002 scenario Ab convolved with

an age spectrum with a mean age of six years and a width

of three years (Waugh et al., 2001). The HALOE-derived

Fy at 55 km was increasing at an approximately constant

rate until about 1996, and at a slower rate thereafter.

Overall, these changes in the Fy record are strongly consistent with expectations from ground-based measurements after considering lag times (~6 years) associated

with transport of tropospheric air to the lower mesosphere

(Waugh et al., 2001; Waugh and Hall, 2002). This lag time

is also consistent with the lag time derived for Cly in

Section 1.2.4.

1.2.5.2 TOTAL AND STRATOSPHERIC COLUMN

INORGANIC FLUORINE

Of the three species relevant for column inorganic

chlorine measurements (HF, COF2, and COClF), COClF

cannot be monitored by FTIR measurements from the

ground but Fy

* = HF + 2COF2 is an excellent surrogate of

the total Fy burden in the stratosphere (Zander et al., 1994).

Monthly mean total column time series of HF, COF2, and

LONG-LIVED COMPOUNDS

1.36

Figure 1-17. The northern midlatitude (30°N-60°N)

stratospheric fluorine budget for 2004 expressed in

volume mixing ratio (VMR) (Nassar et al., 2006b),

which includes hydrogen fluoride (HF), carbonyl fluoride (COF2), chlorofluorocarbonyl (COClF), carbon

tetrafluoride (CF4), CFC-11, CFC-12, CFC-113,

HCFC-22, and HCFC-142b measured by the ACEFTS, along with estimates of 15 minor organic fluorine species (including CFCs, HCFCs, hydrofluorocarbons (HFCs), fluorocarbons (FCs), and halons)

based on predicted values in the previous WMO

Assessment (WMO, 2003). ACE-FTS measurements of COF2 were extended above 40 km based

on the model of Kaye et al. (1991). SF6 is not shown,

but its contribution has been accounted for. The sum

of the fluorine contributions from these species gives

the total fluorine profile, with error bars indicating the

1-sigma precision of 0.05 ppb. Figure 1-18. Time series of HALOE-observed nearglobal HF (black dots) and HALOE-derived total inorganic fluorine (Fy) at 55 km (red dots). The total Fy

was derived from HALOE by applying 86-93% HF/Fy

ratios to the observed HF. The estimated precision

(±1 sigma) of the near-global HF and Fy averages is

on the order of the size of the data points. The black

solid line through the Fy data points is the WMO 2002

baseline scenario (Ab) convolved with a spectrum

with mean of 6 years and width 3 years. Updated

from Anderson et al. (2000).
Fy

* derived from the Jungfraujoch observational database

are shown in Figure 1-19. The Fy

* curve indicates a temporal slowing in 1999, mainly resulting from the COF2

behavior. Overall, there has been a slowing down of the

Fy

* accumulation, with a mean rate of increase of (1.3 ±

1.2)%/yr (1 sigma) determined over the 2000-2004 time

period. This has essentially resulted from the phasing out

of important fluorine-containing source gases (CCl2F2

(CFC-12), CCl3F (CFC-11), and CCl2FCClF2 (CFC-114)),

which at present has only been partially compensated for

by increases in substitute products such as CHClF2

(HCFC-22) and CH2FCF3 (HFC-134a) (Anderson et al.,

2000; Anderson and Russell, 2004). The sum of the HF

and 2 times COF2 contributions as deduced from 2-D

model runs for 47°N (Chipperfield et al., 1997) are also

reproduced on Figure 1-19 (upper trace). These calculations account for the evolution of the relevant source gas

concentrations characterized by past measurements and

by the Ab baseline scenario (WMO, 2003). However, it is

important to mention that, so far, the Chipperfield et al.

2-D model runs have not properly reproduced the partitioning among the two main fluorine reservoirs (with an

overestimation of the COF2 contribution to Fy). A possible cause for the latter discrepancy would be a new

pathway for destroying COF2.

A 27-year time series record of daily mean stratospheric HF columns has also been derived by Rinsland et

al. (2002) from analysis of infrared solar spectra recorded

from the U.S. National Solar Observatory on Kitt Peak

(31.9°N, 111.6°W, 2.09 km altitude). The long-term HF

evolution is modeled by assuming a seasonal cycle combined with a polynomial expression that solves for coefficients to determine linear and quadratic coefficients for

the 14-50 km column as a function of time. The measurements show a continued buildup of the stratospheric HF

column, with a spring maximum and an autumn minimum

superimposed on significant short-term variability. The

latter variations are presumably due to latitudinal transLONG-LIVED COMPOUNDS

1.37

Figure 1-19. Time series of monthly mean total column HF (red circles) and COF2 (green triangles), as derived

from the Jungfraujoch (46.5°N) observational database; Fy* (blue triangles) results from the summation of the

corresponding HF and 2 times COF2 data points. To avoid the influence of the significant variability affecting

measurements in winter- and springtime, these time series have been limited to the June to November months.

Nonparametric least-square fits to these datasets are reproduced as thick black solid and dashed curves; they

help to appraise the two-decade mean evolution of HF, COF2, and Fy*. Also shown is column Fy* above the

Jungfraujoch station calculated with a two-dimensional model (top black curve) (Chipperfield et al., 1997).
port of lower stratospheric air above the Kitt Peak station.

The updated results also indicate a rapid buildup in fluorine loading in the middle atmosphere, and that the rate of

increase has slowed during recent years.

1.3 EMISSIONS ESTIMATED FROM THE HALOCARBON MEASUREMENTS

1.3.1 Regional Emissions and Inverse

Modeling

In the last few years, inverse modeling of groundbased measurements has been increasingly used to

assess emissions of trace gases on regional to continental

scales. This approach provides a tool for the validation

of emissions and a real-world check for compliance with

international agreements. For the Montreal Protocol, in

particular, these regional studies have a potential to be

used as a tool to test the consistency of different phase-out

schemes within the developed and the developing regions

of the world. In this section, an assessment of regional

emissions, which are estimated by inverse modeling, will

be given, together with an attempt to account for uncertainties due to systematic and random errors. A global

integration is then attempted to check for compliance with

worldwide emissions derived from trends, although the

significance of this analysis is limited by large uncertainties in the summed regional emissions.

For the estimation of regional emissions by inverse

modeling of ground-based measurements, the same basic

principles are generally followed. Observations at one or

more locations within the region of interest are separated

by statistical or meteorological methods into periods that

are influenced by emissions and those that are representative of background conditions. Afterwards the regional

emissions are estimated by two principal approaches. In

the first method, emissions are calculated by comparing

concentrations of the gas with unknown sources with concurrent data on a gas for which a regional emission inventory is available (tracer-ratio method). In the second

method, the recent travel histories of the air masses arriving

at the observation site are combined with the measured concentrations. The recent history of the air can be deduced

using a range of methods including single or multiple back

trajectories, or Lagrangian particle dispersion models

(LPDM). Results of such inverse modeling-derived emission estimates for ozone-depleting trace gases for different

regions of the globe are summarized in Table 1-6.

For Europe, emissions of CFCs have been estimated

for the period of 1995-2000 by a combination of measurements from Mace Head (Ireland) and calculations using a

sophisticated LPDM model (Numerical Atmospheric

Dispersion Modeling Environment, NAME) (Manning et

al., 2003). Resulting European emissions in this period

were 8.5-9.6 Gg/yr and 13.5-15.2 Gg/yr for CFC-11 and

CFC-12, respectively. Interestingly, consumption-based

emissions (UNEP, 2002b) show considerably higher

values for CFC-11 (18.5-23.0 Gg/yr), whereas values for

CFC-12 (9.6-25.0 Gg/yr) were in the same range. Recent

European HCFC emissions were estimated to be highest

for HCFC-22 (38-41 Gg/yr) and approximately 10 Gg/yr

or less for HCFCs 141b and 142b (O’Doherty et al., 2004).

For CH3CCl3, a combined analysis of 4 years of data

from Jungfraujoch (Switzerland) and Mace Head (Ireland)

provided estimates of European emissions in the years

2000-2003 between 0.3 and 3.4 Gg/yr (Reimann et al.,

2005), which contrasts with estimated European emissions

of over 20 Gg/yr in 2000 from a short-term aircraft campaign over Germany (Krol et al., 2003). On the basis of

the much larger time period covered by the study of

Reimann et al. (2005), it is assumed that average emissions in Europe after 2000 were in the range of a few Gg/yr

and that the higher results of more than 20 Gg/yr were

probably due to short-term emissions not being representative for longer time periods.

In contrast to these observations, consumptionbased emission estimates for CH3CCl3 in Europe were

declared to be smaller than 0.1 Gg/yr after 1997. The

observed ongoing emissions could have been caused, in

part, by consumers banking CH3CCl3 in anticipation of

the restrictions of the Montreal Protocol. This hypothesis

is supported by the drastic decline in consumption-based

estimates between 1995 and 1996 (McCulloch and

Midgley, 2001), which is not represented by observationbased estimates. Additional sources of CH3CCl3 not covered by industry or Montreal Protocol consumption data

could include hazardous waste sites, illegal imports, and

small releases during HCFC production. Although these

ongoing emissions are not significant for stratospheric

ozone depletion, they point to the possibility of noncompliance with the Montreal Protocol (see Section 1.3.2.5

for further discussion). For Eastern and South Eastern

Europe (Ukraine, Bulgaria, Romania, Turkey, and Greece),

indications exist that emissions of CH3CCl3 and CFC-113

were still occurring at a non-negligible level in August

2001 (Gros et al., 2003).

In the United States, several studies haven been performed in the last decade. Between 1996 and 2003,

ongoing emissions have been detected for CFC-11 and

CFC-12. Whereas sources of CFC-11 have not changed

much over the period, estimated emissions of CFC-12 have

declined unambiguously (Barnes et al., 2003; Hurst et al.,

2006; Li et al., 2005; Millet and Goldstein, 2004). This

LONG-LIVED COMPOUNDS

1.38
behavior can be rationalized by different emission functions for these substances in the U.S. Whereas CFC-11

emissions are primarily originating from in-use foams with

small but continuing emissions from a large reservoir,

CFC-12 emissions from its use in refrigeration and air conditioning occur over a shorter time span and CFC-12 is

being replaced by non-ozone depleting substances. United

States emissions of CFC-113 between 1996 and 2002 were

small and those of CCl4 have essentially been no longer

detectable in the regional studies. On the other hand,

CH3CCl3 emissions were estimated to be between 2.2-4.4

Gg/yr, which is comparable to values in Europe on an

emissions per capita basis.

In Australia and Japan, emissions of the most

important ozone-depleting substances were estimated to

be of the order of a few Gg/yr in the last few years (Dunse

et al., 2005; Yokouchi et al., 2005). Noteworthy is that

CH3CCl3 emissions were no longer detectable in Australia

after 1998, but continuing emissions of 1.7 Gg/yr were

estimated in Japan in 2000-2003. For the developing countries in Asia available emission inventories are of limited

quality. This is unfortunate because production and usage

was still allowed within the Montreal Protocol in many of

these countries (e.g., China and India). An aircraft campaign studying the outflow of the Asian subcontinent

(Transport and Chemical Evolution over the Pacific,

TRACE-P) resulted in estimates of the emissions of several countries within the region in 2001 (Blake et al., 2003;

Palmer et al., 2003). For East Asia (China, Japan, and

Korea), emissions of 21.5 Gg/yr for CCl4, 14.2 Gg/yr for

CH3CCl3 (with a considerable share from sources around

Shanghai), 30.1 Gg/yr for CFC-11, and 39.4 Gg/yr for

CFC-12 were found. In addition, halon-1211 emissions

were 0.2 Gg/yr from Southeast Asia and 2.3 Gg/yr from

South China, primarily originating from the Shanghai

region.

Emission estimates for Russia are also based on

insufficient coverage. This is an important issue because

of large uncertainties in the production and banks of ozonedepleting substances for this country. A campaign along

LONG-LIVED COMPOUNDS

1.39

Table 1-6. Annual emissions derived from regional models and from networks (Gg/yr).

Europe United Australia Japan Asia Russia Global Global

(EU-15) States (including Total Total

Japan) from Emissions

Regional (Section

Models 1.3.2)

CFC-11 8.5-9.61 6.67 0.28 1.49 30.110 1.211 49 90

7.34

CFC-12 13.5-15.21 13.47 0.58 2.29 39.410 6411 135 122

16.44

CFC-113 NA 0.64 ND NA NA 0.811 (1.8) 10

1.46

CCl4 NA <0.64 ND 1.29 21.510 0.611 (22) 70

CH3CCl3 0.3-3.42 2.25 ND 2.09 14.210 0.411 22-46 See

243 3.67 Section

4.46 1.3.2.5

NA: not available; ND: not detected; values in parenthesis: data coverage is too weak to derive a global figure. Global emission estimates are an average from 2001-2003.

1 1995-2000 (Manning et al., 2003). 2 2000-2003 (Reimann et al., 2005). 3 2000 (Krol et al., 2003). 4 1999-2002 (Li et al., 2005); for the total emissions from the U.S., combined emissions from (Li et al., 2005) and (Barnes et al., 2003) were used,

except for CCl4, where emissions from California alone were extrapolated to the U.S. using a population figure of 273 million. 5 2001-2002 (Li et al., 2005). 6 2001-2002 (Millet and Goldstein, 2004). 7 2003 (Hurst et al., 2006) (only estimating U.S. emissions, related to 273 million inhabitants). 8 1998-2000 (Dunse et al., 2005). 9 2002 (Yokouchi et al., 2005). 10 2001 (Palmer et al., 2003). 11 2001 (Hurst et al., 2004).
the Trans-Siberian railway (Trans-Siberian Observations

Into the Chemistry of the Atmosphere, TROICA-7) indicated only very small emissions of 1.2 Gg/yr for CFC-11,

which contrasted with high estimates of 64 Gg/yr for CFC12 (Hurst et al., 2004). The latter possibly originated from

releases due to leakage from excessively large banks or

from unaccounted production. On the other hand, emissions of CCl4 and CH3CCl3 were small, in accordance with

expectations.

Summation of the regional emission estimates for

CFCs, CH3CCl3, and CCl4 (Table 1-6) provides a consistency check on the global emissions that have been calculated from the combination of the background trends in

their concentrations and their atmospheric lifetimes

(Section 1.3.2). This comparison is however hampered

by several likely sources of random and systematic errors.

The largest source of potential error is the result of the substantial gaps in the spatial and temporal coverage in the

existing regional studies. Whereas in Europe, the U.S.,

and Australia continuous measurements are being routinely exploited for deriving regional emissions, capabilities for doing this in regions with a substantial potential

for large uncounted emissions (e.g., Southeast Asia) are

still to be developed. Furthermore, the emission estimates

described contain uncertainties resulting from the meteorological information or from the emission inventories that

are often used as a priori information. In addition, errors

can arise from inadequate calibration or sampling (e.g.,

contamination) as well as insufficient measurement precision. The quantification of these errors is far from straightforward.

A comparison between several of the reported

regional emission estimates provides a possible range of

uncertainty for an individual estimate. For methyl chloroform three independent studies were performed in the

United States within the period of 2001-2003 and ranging

from the East Coast to the West Coast (Hurst et al., 2006;

Li et al., 2005; Millet and Goldstein, 2004). The three

emission estimates for the U.S. ranged from 2.2-4.4

Gg/yr, giving a standard deviation around the mean of

33%. Larger uncertainties were found in a study on

European methyl chloroform emissions (Reimann et al.,

2005), which was based on two measurement sites and

three techniques of analysis. The spread of the emission

estimates for Western Europe ranged from 0.33 to 3 times

the geometric annual means. The larger uncertainties for

Europe than for the U.S. possibly originate from traditional differences in industrial activities between the

European countries.

The substantial range of uncertainty in the regional

emissions indicated by the European estimates for CH3CCl3

may have been significantly affected by the generally small

remaining emissions detected in these recent observations.

In fact a substantially smaller spread in the estimates from

the Mace Head observations for 1995-2000 is reported in

Reimann et al. (2005). Moreover Simmonds et al. (2006),

using continuous observations at Mace Head, report a

spread in the annual estimates of Western European emissions of dichloromethane, trichloroethene, and tetrachloroethene, for all of which the current emissions are

more than ten times those of CH3CCl3, of a factor of three

in the worst case. Therefore when pollution events may be

easily separated from baseline observations, a factor of

three in the inferred regional emission uncertainty is probably to be expected. This factor is likely to increase significantly when the pollution events are small. In addition,

for some large areas of the world (e.g., Asia), where estimates are dependent on only a few short-term campaigns,

the uncertainties are also likely to be considerably larger.

A first indicative comparison between global and

the summed regional emission estimates is attempted

(Table 1-6). For CFC-11, only about 54% of the estimated

global emissions are found in the summed regional emissions, with the partitioning between CFC-11 and CFC-12

from Russia especially appearing to be suspect. For CFC12 a good consistency is obtained, with the principal contributions coming from Russia and China. For CFC-113

regional emissions are consistently low, with most studies

not being able to identify detectable emissions. For CCl4

the only detectable current emissions were from China,

but the amount was only a relatively small fraction (31%)

of the inferred global emissions for a 26-year lifetime.

Unfortunately only limited data are available from regions

such as Korea and India, which might be responsible for

significant emissions. There are indications that India in

particular is still a source of CCl4 emitted from solvent use

and from process agent applications. For example, air

from India observed at a background site in the Indian

Ocean has been found to contain elevated concentrations

of CCl4 (Lobert and Harris, 2002). Furthermore, a study

in Mumbai (India) found concentrations of CCl4 that were

noticeably above the background (Srivastava et al., 2006).

For CH3CCl3, indications of Indian emissions have also

been found during campaigns, although their extent could

not be quantified (Lobert and Harris, 2002; Srivastava et

al., 2006).

In summary, however, the present spatial and temporal coverage, especially in developing regions (e.g.,

Southeast Asia), is still too inhomogeneous to provide

quantifiably useful comparisons of the summed regional

emissions with the global emissions derived from trends

in observed concentrations.

LONG-LIVED COMPOUNDS

1.40
1.3.2 Global Emissions

This section describes time series of global emissions inferred from the measurements by the ground-based

sampling networks (see Section 1.2.1). For the AGAGE

measurements, the estimates are based on an inverse procedure using the AGAGE two-dimensional model (12

boxes: 4 in the lower troposphere, 4 in the upper troposphere, and 4 in the stratosphere), with transport parameters tuned to simulate the observed latitudinal gradients

for the CFCs (Cunnold et al., 1994; 1997). There are five

AGAGE sites with long time series of measurements (e.g.,

for the CFCs) using electron capture detectors, but there

are only two AGAGE sites (Mace Head, Ireland, and Cape

Grim, Tasmania) employing mass spectrometer detectors

that have provided shorter time series of measurements of

gases such as the HCFCs and the HFCs.

For the derivation of global emission rates from the

NOAA/ESRL measurements, the estimation procedure

starts from their reported tropospheric mean mole fractions. Those tropospheric means are derived from

weighting (usually by latitude) high-frequency results at 5

sites, or low-frequency flask measurements at 8-10 remote

sites (Montzka et al., 1999). The stratospheric contents

are derived from the tropospheric means by ratioing the

values to the AGAGE tropospheric means and multiplying

by the stratospheric contents calculated in the 12-box

model that was tuned to the AGAGE measurements.

Measurements of the halocarbons in the Australian

archived air samples have been used to extend the time

series back to 1978 for a number of the gases (e.g.,

Langenfelds et al., 1996; Fraser et al., 1996). In this case,

the 12-box model, and other two-dimensional models

(Oram et al., 1995; Fraser et al., 1999), have been used to

infer global means and global emission rates from measurements at this one location.

For most of these long-lived gases, the largest

sources of uncertainty in the emission estimates are the

species lifetimes. Smaller uncertainties arise from

absolute calibration uncertainties (except for halon-1301,

for which the calibration uncertainties are quite large) and

from inadequacies in simulating atmospheric transport.

The effects of measurement errors are substantially

reduced by reporting emissions averaged over three-year

periods. The effects of these errors may be seen in Figures

1-20 and 1-21 in differences, if any, between the emission

estimates derived from the AGAGE and the NOAA/ESRL

measurements. The upper and lower envelopes of the error

bars in these two figures for the AGAGE-based estimates,

corresponding to the shortest and the longest lifetimes

respectively, were calculated using the 12-box model using

those lifetimes. Note that the trends in the emission estimates in %/yr are not very sensitive to the lifetime uncertainties. Errors in the representativeness of the measurement sites for deriving global means also adds uncertainty

to derived emissions, but during periods when emissions

and atmospheric gradients are small (CFCs and CH3CCl3

currently), these uncertainties are dramatically reduced.

1.3.2.1 CFCS

Figure 1-20 illustrates the global emission estimates

derived from the combination of AGAGE and NOAA/-

ESRL measurements for three different years for the CFCs.

It shows that, as suggested by the latitudinal gradient information in Section 1.2.1.1, emissions of CFC-11 and CFC12 are continuing but at lower levels than in 2000 (Table

1-7). Moreover, the emissions are about 25% of those (in

1986) before the Montreal Protocol was enacted. On the

basis of the regional emission estimates summarized in

Section 1.3.1, it would appear that most of the remaining

emissions are coming from Asia and Russia. It should be

noted that the estimates for CFC-11 emissions in 2003 are

very sensitive to its lifetime. Emissions of CFC-113 have

now almost ceased (now 3% of 1986 values) and the

regional emission estimates, at the few sites where measurements have been made, support the current, low global

emissions. Except for CFC-11, the decays of the CFC

emissions from 2000 to 2003 derived from global observations are not significantly different from those predicted

by the Ab (most likely) scenario from WMO 2002. For

CFC-11, a significantly faster decline in the emissions was

projected in that scenario relative to those given in Figure

1-20. The emissions from the Ab scenario are not shown

in Figure 1-20 because those emissions were calculated

from the Ab scenario mixing ratios using a one-box model,

whereas the emissions in Figure 1-20 were calculated

using a 12-box model. The differences between emissions

calculated by the 12-box model and the 1-box model are

illustrated in Figure 8-1 of Chapter 8. It can therefore be

misleading to compare emissions calculated by the two

different models. The faster decline in the CFC-11 emissions in 2000-2004 than projected in WMO 2002 is therefore inferred from the more rapid decline in the observed

mixing ratios (see Section 1.2.1.1) versus those projected

in the Ab scenario.

An important objective of this report is to estimate

the remaining banks, which will be responsible for the

emissions of ozone-destroying halocarbons over the next

few years. Total emissions up to the present time can be

estimated from the ground-based measurements. These

are shown in Table 1-7. An assumption about how these

emissions were distributed in time prior to the start of

observations in 1978 is needed. However, because the

LONG-LIVED COMPOUNDS

1.41
LONG-LIVED COMPOUNDS

1.42

1975 1985 1995 2005

YEAR

0

200

400

600

3 YEAR SMOOTHED EMISSIONS (Gg/yr)

CFC1975 1985 1995 2005

YEAR

0

200

400

600

3 YEAR SMOOTHED EMISSIONS (Gg/yr)

CFC-12

1975 1985 1995 2005

YEAR

0

100

200

300

3 YEAR SMOOTHED EMISSIONS (Gg/yr)

CFC- 13

1975 1985 1995 2005

YEAR

0

100

200

300

3 YEAR SMOOTHED EMISSIONS (Gg/yr)

Carbon Tetrachloride

Figure 1-20. Annual global emissions for CFC-11, CFC-12, CFC-113, and carbon tetrachloride, smoothed

over three-year intervals, calculated by inverting measurements, using the 12-box model, from the AGAGE

(green lines) and NOAA/ESRL (red diamonds) ground-based networks (as in Cunnold et al., 1997). The calculations used equilibrium lifetimes of 45, 100, 85, and 26 years for CFC-11, CFC-12, CFC-113, and carbon tetrachloride, respectively. The green error bars indicate the uncertainties in the estimates due to the uncertainties

in the species lifetimes (see Table 1-7). The TEAP (UNEP, 1998) estimate of the CCl4 emissions in 1996 is

indicated in black.
LONG-LIVED COMPOUNDS

1.43

1975 1985 1995 2005

YEAR

0

20

40

60

80

100

3 YEAR SMOOTHED EMISSIONS (Gg/yr)

HCFC-141b

1975 1985 1995 2005

YEAR

0

10

20

30

40

50

3 YEAR SMOOTHED EMISSIONS (Gg/yr)

HCFC-142b

1975 1985 1995 2005

YEAR

0

100

200

300

400

3 YEAR SMOOTHED EMISSIONS (Gg/yr)

HCFC-22

Figure 1-21. Annual global emissions for HCFC-141b, HCFC-142b, and HCFC-22, smoothed over threeyear intervals, calculated by inverting the measurements (see Figure 1-2) from the AGAGE (green lines

starting from 1995) and NOAA/ESRL (red diamonds) ground-based networks and measurements of the air

archived in Australia (green lines before 1995) with a 12-box model (as in, and updated from, O’Doherty et

al., 2004). The calculations were based on equilibrium lifetimes of 9.3, 17.9, and 12.0 years for HCFC-141b,

HCFC-142b, and HCFC-22, respectively. The green error bars indicate the effects of uncertainties in the lifetimes listed in Table 1-7.
largest emissions occurred after 1979, and they have been

in the atmosphere for a shorter time, this introduces uncertainties that are small compared with the effects of atmospheric lifetime uncertainties that are shown in Table 1-7.

The remaining banks may be estimated by subtracting the

accumulated production from the numbers in Table 1-7.

Chapter 8 provides both those numbers and estimates of

the banks. The effects of using the 12-box model used

here to estimate the emissions that have already occurred

may be compared against the one-box model results given

in Chapter 8.

Table 1-7 also provides estimates of the total emissions integrated from the beginning of production to 2004.

It may be noted that these total emission estimates and

their one sigma uncertainty limits, when combined with

total production numbers that are known fairly accurately

(e.g., from UNEP 2005), can be used to constrain the

expectations of future emissions (e.g., the current global

banks of unreleased gases).

1.3.2.2 HCFCS

Figure 1-21 shows HCFC emission estimates

derived from applying the 12-box model to the AGAGE

and NOAA/ESRL ground-based network measurements.

The two sets of estimates differ because of absolute calibration differences (e.g., 4-5% for HCFC-142b, less for

HCFC-141b and HCFC-22) and some other minor differences in the most recent measurements. The emissions of

HCFC-141b (54.4 Gg/yr in 2003) and HCFC-142b (25.1

Gg/yr in 2003) are calculated to have decreased from 2000

to 2004 by approximately 15% (10 and 4 Gg/yr for HCFC141b and HCFC-142b, respectively) and this is seen by

both networks. These emission estimates correspond to a

recent decrease in the rates of growth in the atmospheric

mole fractions compared with the growth rates reported in

WMO 2002 (Figure 1-2, Montzka and Fraser et al., 2003).

HCFC-22 emissions have been increasing fairly slowly or

not at all from 2000 to 2004 (e.g., at less than 2%/yr for

AGAGE and less than 1%/yr for NOAA/ESRL). This is

down from increases of approximately 3%/yr from 1994

to 2000. Note that the ranges of uncertainty, in percent,

on the 2003 emissions of the HCFCs and HFC-134a (Table

1-7) are less than for those other gases in Table 1-7 that

have much longer emission histories.

The recent decreases in the emissions of HCFC141b and HCFC-142b, and the fact that the measured mole

fractions in 2000-2004 are less than projected (see Section

1.2.1.2) in the best-guess scenario in WMO 2002, may

mean that emissions are occurring more slowly than

expected and the future emissions from the banks will be

larger and occur over a longer time than originally anticipated.

1.3.2.3 HALONS

Difficulties in reconciling the halon measurements

with emission estimates derived from production and the

LONG-LIVED COMPOUNDS

1.44

Table 1-7. Global emission estimates. Emissions are in Gg/yr, smoothed over 3-year intervals and given to

3 significant figures, and are derived using a 12-box model from a combination of AGAGE and NOAA/ESRL

measurements. The total emissions are based solely on AGAGE measurements and they include emission

estimates for 2004. The lifetimes and their uncertainties are from Sections 1.2.2.1 and 1.2.2.2. Uncertainty

ranges, instead of standard deviations, are given because the central values typically do not lie in the centers

of the uncertainty ranges, primarily because the estimated emissions are typically proportional to the inverses

of the lifetimes.

Species Release Estimates Rate of Change Total Lifetime Lifetime 2003 Emission Total Emissions 

(Gg/yr) (2000-2003) Emissions Used Uncertainty Estimate Uncertainty Range

1986 2000 2003 (%/yr) (1920-2004) (yr) Range (yr) Range Gg %

(Gg)

CFC-11 360 100 88 −4.3 9,450 45 57-35 60-126 8,700-10,460 18

CFC-12 472 150 114 −9.1 14,300 100 113-79 101-144 14,000-15,000 7.5

CFC-113 186 18 6 −33.3 3,170 85 100-66 1-15 3,080-3,340 8.4

CCl4 130 73 71 −0.9 6,620 26 30-16 58-131 6,060-8,590 34

HCFC-22 154 258 270 +1.5 6,040 12 14.0-10.0 242-310 5,530-6,760 20

HCFC-141b 0 58.8 54.4 −2.6 589 9.3 11.2-7.4 48.4-63.5 548-651 17

HCFC-142b 1.5 29.1 25.1 −4.9 407 17.9 22.2-13.6 22.3-29.6 380-450 17

HFC-134a 0 76 104 +10.4 721 14 16.4-11.6 99-111 696-757 8.4
remaining banks of halon-containing fire extinguishing

equipment (e.g., the estimates shown on page 75 of

IPCC/TEAP 2005) were previously pointed out by Fraser

et al. (1999). Currently the remaining very small worldwide production of these halons is insignificant in relation

to the size of the global halon bank. For halon-1301, the

estimates are from the United Nations Environment

Programme/Halons Technical Options Committee (UNEP,

2003) emission scenario. Those emissions have been run

through the 12-box model to yield the gold curve in the

halon-1301 panel of Figure 1-4 (which is similar to that in

Figure 1-17 of WMO 2002). A smaller emission rate for

halon-1301 has recently been suggested by the work of

Verdonik and Robin (2004).

For halon-1211, there is less spread in the measurements and the mole fractions estimated from the

emission banks do not differ as much as they do for

halon-1301. The emission scenario used is that of TEAP

(an update to the UNEP (1999) model, see WMO 2002,

Section 1.6.1.5), which is similar to the UNEP (1999)

model. It should be noted that this scenario contains

lower emissions than those shown in IPCC/TEAP (2005,

page 75), which were from UNEP (2003). An error in

that model was recently discovered in accounting for the

production by Article 5(1) countries in 1987-1993. The

WMO 2002 Ab scenario also provides a reasonable fit to

the observations except that, although the observations

indicate a slowing of the increases, there is little evidence that halon-1211 stopped increasing in the 2000-

2004 period as the Ab model predicted.

1.3.2.4 CARBON TETRACHLORIDE

Table 1-7 and Figure 1-20 show the estimated emissions of carbon tetrachloride, for a lifetime of 26 years,

determined from the measurements by the ground-based

sampling networks combined with AGAGE 12-box model

calculations. The emission uncertainty error bars are based

on the combined uncertainties in the contributions of stratospheric, oceanic, and soil losses. A lifetime of 20 years

increases the calculated emissions by approximately 26

Gg/yr, from 71 Gg/yr in 2003 to 97 Gg/yr. Simmonds et

al. (1988 and 1998) argued that most of the CCl4 emissions in the past were fugitive emissions from the carbon

tetrachloride used in the manufacture of CFCs. For a lifetime of 20 years, not only would the CCl4 emissions have

to have been equivalent to approximately 16% of CFC

production prior to 1994, versus approximately 9% based

on International Trade Commission figures on CCl4 production in the U.S. in 1979-1985 (Simmonds et al., 1988),

but the proportion would have had to have increased to

almost 60% after 1995 as a result of the substantial reduction in CFC manufacture (cf. WMO 2002, Montzka and

Fraser et al., 2003, Figure 1-15). TEAP (UNEP, 1998)

indicated that approximately 23 Gg/yr of carbon tetrachloride emissions could arise from uses other than in CFC

manufacture; that report also estimated total emissions

from industry data of 31 to 62 Gg in 1996, with a most

likely value of 41 Gg (see Figure 1-20). Therefore it is

difficult to reconcile the atmospheric observations with an

atmospheric lifetime for CCl4 of 20 years.

In addition, recent high-frequency atmospheric

measurements (see Section 1.3.1) at a number of global

locations have only detected approximately 22 Gg/yr in

2001, which came almost entirely from Eastern Asia.

Therefore, even though emissions from India are

unknown, it is difficult to account for the approximately

97 Gg/yr in 2003 (99 Gg/yr in 2001) required for consistency with a 20-year lifetime (or even the 71 Gg/yr in 2003

required for a lifetime of 26 years). Simmonds et al. (1998)

showed that the observed latitudinal gradient from the

AGAGE observations was reasonably well simulated by a

lifetime of 42 years. Updated calculations suggest a gradient-based lifetime of approximately 35 years, unless a

substantial fraction of the emissions are coming from the

Southern Hemisphere.

Overall the budget of carbon tetrachloride remains

poorly understood. Because the currently observed

regional emissions, the TEAP emission estimate, and the

measured interhemispheric gradient seem inconsistent

with an overall lifetime of 20 years and because the uncertainty ranges for both the ocean and soil sinks allow for

the possibility of only very small sinks, no change from

the lifetime of 26 years given in WMO 2002 is recommended at this time.

1.3.2.5 METHYL CHLOROFORM

Global production and sales of methyl chloroform

from 1970 to 1995 were tabulated in a database from annually provided figures from the producing companies.

According to McCulloch and Midgley (2001), the uncertainties were 2.1% on production and 4.2% on emissions.

These uncertainties are smaller than for most other ozonedestroying gases because of the smaller number of companies that produced methyl chloroform and because of the

generally short time (less than a year) between production

and emission into the atmosphere. After 1995, emissions

were derived from figures on production and consumption

reported to the United Nations Environment Programme

(UNEP, 2002b). Global emissions have been calculated to

be 19.7 ± 2.3 Gg in 2000 and 12.9 Gg in 2002. This may be

compared against maximum emissions of 718 Gg, which

occurred in 1990 (McCulloch and Midgley, 2001).

LONG-LIVED COMPOUNDS

1.45
McCulloch and Midgley also pointed out that there

are discrepancies in the databases for 1994-1996. In particular, there is a production versus consumption discrepancy in the UNEP figures in 1995-1996. It was noted in

Section 1.3.1 that regional emission estimates suggested

an anomalous large disagreement for methyl chloroform

emissions versus global emissions in 1996. Furthermore

Prinn et al. (2005) suggested that, based on summing the

recent regional emission estimates, the global emissions

in the past ten years might be somewhat low. They pointed

out that the regional emissions could be approximately

accommodated (within the uncertainties of the measurements) by placing 5% of the global emissions in a slower

release category, with emissions being divided equally

over a 10-year interval. This would not have affected the

production numbers. The slower releases would have

increased the emissions in 2002 by 3.6 Gg to 16.5 Gg and

from 93.5 Gg to 112.3 Gg in 1996. The revised global

totals would not, however, account for the European emissions estimate of over 20 Gg in 2000 by Krol et al. (2003),

which Prinn et al. (2005) did not include because of the

much lower estimates from the European ground stations

(see also Section 1.3.1). This discrepancy, and hence the

global emissions since 2000, are not understood. These

ongoing emissions are not, however, significant for stratospheric ozone depletion.

Historically, accurate methyl chloroform emissions

data have been useful in the estimation of global OH abundance. Prinn et al. (2005) used the emission estimates just

described to calculate the annual atmospheric loss rates

for methyl chloroform that produce the best fit to the

AGAGE observations. This resulted in a time series of

annual globally averaged tropospheric OH concentrations

from 1979 to 2003 and particularly low OH values (by

10%) in 1997/1998. Monte Carlo estimates of the uncertainties in the recent emissions of methyl chloroform produce uncertainties of only a few percent in these OH estimates. Wennberg et al. (2004) pointed out that the methyl

chloroform that had entered the ocean could have been

trapped in the polar region and, because of the very steep

decline in the atmospheric concentrations in recent years,

much of it could now be returning to the atmosphere. Prinn

et al. (2005) estimated that these ocean fluxes could make

annual contributions to the methyl chloroform budget that

varied between 4 and 7 Gg/yr between 1995 and 2003.

This increased the OH estimates by approximately 5%

after 2000.

In order to derive methyl chloroform emissions

using the ground-based networks’ measurements, independent estimates of OH are required. Calculations of OH

in chemical models yield a methyl chloroform lifetime

approximately 10% less (Spivakovsky et al., 2000) than

the Prinn et al. value of 4.9 years, and other models produce larger differences (WMO 2002, Montzka and Fraser

et al., 2003). Therefore because of the relatively small

uncertainties in the emissions given by McCulloch and

Midgley (2001), it is not generally very productive to estimate emissions from the measurements.

1.4 OTHER TRACE GASES

1.4.1 Carbon Dioxide (CO2)

Carbon dioxide is the primary agent of anthropogenic climate change, accounting for 60% of the

increase in direct radiative forcing since preindustrial

times (IPCC, 2001). Climate change can affect the ozone

budget, for example due to temperature changes in the

polar vortices, and due to changes in atmospheric transport. Such changes may delay ozone recovery. In 2004,

global annual average (surface) CO2 was 377 ± 0.2 parts

per million by volume (ppmv), as determined through the

NOAA/ESRL global network.

The atmospheric growth rate of CO2 has averaged

about 0.5% per year over the nearly 50 years of in situ

measurements (Keeling and Whorf, 2005). The observed

atmospheric growth rate is explained by about one-half of

CO2 emissions from fossil fuel combustion and land use

change being absorbed at the surface of the Earth. CO2

can be taken up by either the ocean or terrestrial biosphere;

there is considerable year-to-year variability in the uptake

that, at the global scale, is closely related to the impact of

El Niño on both ocean and land. Two tracers complementary to CO2 are molecular oxygen (O2) and 13CO2. Both

tracers allow for partitioning of total uptake into terrestrial and oceanic components, and both indicate that the

majority of interannual variability is driven by terrestrial

processes (Battle et al., 2000). Over the long term (> 500

years), however, it is likely that the oceans, as the largest

carbon reservoir, will exert the strongest control on atmospheric CO2.

Future levels of atmospheric CO2 will be strongly

determined by future fossil fuel emissions. Additionally,

there are no guarantees that the current uptake rate of 50%

will be maintained, because both the terrestrial and oceanic

carbon cycles may change as a result of changing climate.

Furthermore, the mechanisms by which carbon is absorbed

at the surface, especially by the terrestrial biosphere, are

not fully understood. Simulations in which simple models

of the carbon cycle are coupled to climate models have

shown positive feedbacks between temperature and CO2

emissions in the 21st century, which can reduce or even

eliminate net CO2 uptake of anthropogenic emissions. In

LONG-LIVED COMPOUNDS

1.46
one study (Cox et al., 2000), coupling resulted in an additional 1.5 K global temperature increase by 2100, compared with the uncoupled case. Although the magnitude

of the feedbacks is uncertain, the direction is likely to be

positive, and it is clear that the carbon cycle is a first-order

uncertainty in prediction of future climate.

1.4.2 Methane (CH4), Including Isotopic

Data

Methane is important for its impact on climate both

through its radiative properties and chemically as a major

source of water vapor to the stratosphere and as a sink for

Cl. In the troposphere, its oxidation products are controlling the chemistry. The global average CH4 mixing ratio

for 2004 at the surface was 1777 ± 0.6 ppb, determined

from marine boundary layer sites in the NOAA/ESRL and

other surface monitoring networks (GLOBALVIEW-CH4,

2005). This compares with glacial values of about 400

ppb (Spahni et al., 2005) and 700 ppb in 1800 AD

(MacFarling-Meure, 2004). The precise reasons for interglacial/glacial CH4 variability as well as variability during

the Holocene (Flückiger et al., 2002; Ruddiman, 2003) are

not well known but may relate to large-scale moisture and

temperature changes affecting wetland emissions. The

rise to present-day values has resulted from increases in

emissions related to fossil fuel extraction, biomass combustion, and agriculture, and current atmospheric mixing

ratios are unprecedented in at least the last 650,000 years

(Spahni et al., 2005). The global growth rate of methane

has varied from about 14 ppb/yr in 1985 to near zero in

2000 but with significant interannual variability. Since

2000, the growth rate has averaged to 0 ± 4 ppb/yr. The

decline in growth rate has been explained as an approach

to steady state, with a constant total of global sources and

constant sink strength (Dlugokencky et al.,1998). However, despite evidence for only small variations in the

global total, individual sources like biomass burning (van

der Werf et al., 2004) and emissions from the former Soviet

Union (Dlugokencky et al., 2003) have seen substantial

changes during this period (Simpson et al., 2002).

Changes in the dominant CH4 sink, the hydroxyl radical

(OH), also affect atmospheric interannual variability and

trends (Bousquet et al., 2005; Krol et al., 1998; Prinn et

al., 2001).

The source budget of atmospheric methane is very

diverse, with no dominant source (IPCC, 2001), making

determination of the causes of atmospheric changes difficult. Major sources include boreal and tropical wetlands,

leakage during coal and natural gas extraction, enteric fermentation, fires, and rice agriculture. Recent laboratory

measurements (Keppler et al., 2004) have also shown

living plants to be a source, but their global impact is currently unknown. “Bottom-up” inventories rely on spatial

and temporal extrapolation of small-scale field measurements, leading to significant uncertainties in the total. On

the other hand, “top-down” approaches inverting atmospheric data can accurately determine continental-scale

totals, but cannot independently identify the different

sources in a given region. Prediction of future atmospheric

levels of CH4 would require, at a minimum, a much better

understanding of how different sources and sink variations

have changed the atmospheric burden.

Measurements of the carbon (13C) (Lowe et al.,

1994; Quay et al., 1999; Miller et al., 2002) and hydrogen

(

2

H or “D”) (Quay et al., 1999; Rice et al., 2001) isotopic

ratio of atmospheric methane (δ13C or δD) can be used to

help quantify the contribution of different sources to the

atmospheric growth rate. For δ13C, sources can be divided

into microbial, biomass burning, and fossil fuel sources,

which have distinct isotopic signatures. Isotopic and

mixing ratio mass balances can be used to constrain the

source distribution of these broad categories. One consistent result of carbon isotopic data is to infer a significantly

stronger global biomass burning source than bottom-up

estimates both in the modern (Quay et al., 1999; Miller et

al., 2002) and paleo atmospheres, where recent results

have shown much larger than expected combustion

sources between 0 and 1200 AD (Ferretti et al., 2005), possibly due to human activity. Though there are currently

only a few δD measurements, they, along with 14CH4 measurements (which help constrain the fossil-fuel related

source), will likely allow for better determination of

atmospheric growth rate variations. New measurements

of CH4 from space (Frankenberg et al., 2005), with

unprecedented spatial coverage, also afford an opportunity to improve our understanding of the CH4 source distribution.

The broad range of CH4 sources (roughly half

anthropogenic, half non-anthropogenic) and the fact the

source mixture has likely changed over the last century,

imply great difficulty in prediction of future mixing ratios.

Predictions range from 1575 to 3730 ppb, depending on

emission scenario (IPCC, 2001). However, these predictions do not include the possibility of large source changes

like destabilization of vast quantities of CH4 hydrates in

Arctic permafrost, which adds further uncertainty to predicted methane abundances at the end of the 21st century.

1.4.3 Nitrous Oxide (N2O)

Nitrous oxide is the major source of nitrogen

oxides (NOx) in the stratosphere, with the latter contributing significantly to ozone depletion. In addition,

LONG-LIVED COMPOUNDS

1.47
N2O, which has a lifetime of about 120 years, is an important greenhouse gas targeted by the Kyoto Protocol. Its

actual global concentration is close to 320 ppb. This constituent is emitted at the Earth surface by a large number

of sources; major natural contributions are from the

oceans and tropical forests, while most important anthropogenic emissions are from cultivated soils, cattle, and

biomass burning (IPCC, 2001).

Several studies have dealt with the historical evolution of the N2O atmospheric concentration. Nitrous oxide

levels started to decrease in the early Holocene and reached

minimum levels (less than 260 ppb) around 8,000 B.C.,

before exhibiting a slow increase to the preindustrial value

of about 265 ppb. MacFarling-Meure (2004) has produced

a 2,000-year record of atmospheric N2O variations from

measurements on air in ice from Law Dome, Antarctica.

Notable features in the N2O record are a preindustrial level

of 260-275 ppb, a slowing of growth in the 1940s and

1950s, a 5 ppb lowering during the Little Ice Age (LIA:

1300-1850 A.D.), and a 10 ppb increase between 670 and

810 A.D. The origin of this latter increase is unknown and

was not reflected in corresponding CH4 or CO2 changes.

N2O isotopic composition studies of air from

Greenland ice and Antarctic (South Pole) firn (1785-1990)

show a decline in δ15N (1.9 per mil) since preindustrial

times, consistent with a 30% increase in total N2O emissions that are primarily related to agricultural activities

(Sowers et al., 2002). A similar decrease in δ15N (1.8 per

mil) since preindustrial times has been reported from

analyses of Antarctic (Maud Dronning Land and Dome C)

firn air (Röckmann et al., 2003). The corresponding latter

data have been interpreted as due to an increasing soil

source of N2O with a time-dependent isotopic signature.

A larger trend in δ15N (2.8 per mil) of N2O since preindustrial times has been reported in firn and ice air from

Greenland and Antarctica (Berkner Island). The data are

consistent with an increase in the use of fertilizers in agriculture (Bernard et al., 2006).

Since the 2002 WMO Assessment, measurements

performed by the AGAGE and NOAA/ESRL networks

have confirmed the steady increase of the N2O atmospheric

concentration observed over the last two decades: globally averaged surface mixing ratio amounted to 318-319

ppb in January 2005 (updated from Prinn et al., 2000;

Conway et al., 2004). On the basis of flask samples collected regularly at 11 worldwide sites from 1981 to 1996

(Khalil et al., 2002), concentrations are higher on average

by 0.7 ± 0.04 ppb in the Northern Hemisphere, and differences of concentrations over land and oceans indicate

large land-based emissions. Annual rates of change of the

N2O concentrations have been monotonic but quite variable, with a mean annual increase of 0.69 ± 0.02 ppb/yr

(0.22%/yr) for 1985-1996 (Khalil et al., 2002) and 0.74 ±

0.02 ppb/yr (0.24%/yr) for 1986-2005 (updated from Prinn

et al., 2000).

Remote Fourier transform infrared (FTIR) measurements performed at the NDACC site of Jungfraujoch

have also shown a regular buildup of the N2O atmospheric

burden, with a linear fit to the monthly mean total column

dataset indicating an annual increase of 1.06 × 1016

molec/cm2, which corresponds to 0.26%/yr in 2003-2004

(Zander et al., 2005).

1.4.4 Hydrofluorocarbons (HFCs)

HFCs have rapidly replaced CFCs in refrigeration,

foam blowing, and aerosol applications. For example,

mobile air conditioning accounted for 60% of CFC-12 in

1990 and by 1994 had been replaced by HFC-134a in all

but a few countries. HFC-134a (CH2FCF3) and HFC-23

(CHF3) are the two most abundant hydrofluorocarbons in

the atmosphere, and they are continuing to grow rapidly

(Prinn et al., 2000; Sturrock et al., 2001; O’Doherty et al.,

2004). Recently Culbertson et al. (2004) have also

reported measurements of these two principal HFCs from

the analyses of flask samples collected at Cape Meares,

Oregon; Point Barrow, Alaska; and Palmer Station, Antarctica. Updated measurements of HFC-125 (CHF2CF3)

and HFC-152a (CH3CHF2), which were first reported in

the previous Assessment (WMO, 2003), also exist.

1.4.4.1 HFC-134a

HFC-134a has rapidly mostly replaced CFC-12 in

several refrigeration and air-conditioning applications.

As a consequence of increased emissions, the annual

global mean concentration in mid-2004 was about 30 ppt,

increasing at an annual average rate of approximately 3.9

ppt/yr (2000-2004) (O’Doherty et al., 2004; updates to

Montzka et al., 1996) (Figure 1-22 (a)). Annual emissions have been increasing at approximately 10 Gg/yr

from 2000 to 2004 (Figure 1-22 (b)). Because so little of

the HFC-134a has yet been destroyed in the atmosphere,

the error bars on the annual emission estimates in 2003

are just ±6 Gg/yr.

1.4.4.2 HFC-23

HFC-23 is primarily emitted as a byproduct of

HCFC-22 production, with smaller amounts released from

semiconductor etching and refrigeration and fire extinguisher losses (Oram et al., 1998). HFC-23 has been measured since 1978 in flask samples from Cape Grim (Oram

LONG-LIVED COMPOUNDS

1.48
et al., 1998) and, until 1997, approximately 10% larger

values from Oregon (Culbertson et al., 2004). The mixing

ratio at Cape Grim in 2004 was approximately 18 ppt and

the average growth rate in 2001-2004 was approximately

0.7 ppt/yr (Figure 1-23). This rate of growth is significantly less than that (approximately 0.9 ppt/yr) in 1994-

2001. The IPCC/TEAP report (Velders and Madronich et

al., 2005, Figure 2.4) shows estimated emissions, on the

basis of the Cape Grim measurements, of 13 Gg/yr in 2001

versus 6.4 Gg/yr in 1990.

1.4.4.3 HFC-143a, HFC-125, AND HFC-152a

Measurements of HFC-143a (CH3CF3) have been

reported by Culbertson et al. (2004) for flask samples collected at Cape Meares, Oregon; Point Barrow, Alaska; and

Palmer Station, Antarctica. Tropospheric mixing ratios in

the Oregon samples increased from about 0.2 ppt in 1978

to about 2 ppt in 1997. Average annual concentrations

during 1997 at Alaska and Antarctica were 1.7 and 1.3 ppt,

respectively.

HFC-125 is mainly used in refrigeration, while the

primary uses of HFC-152a are in foam-blowing and

aerosol propellants. Measurements from Cape Grim

(updated from Sturrock et al., 2001) and at Mace Head are

shown in Figure 1-23. In 2000-2004, mixing ratios have

been increasing approximately exponentially at 23%/yr

(HFC-125) and 17%/yr (HFC-152a). In mid-2004, globally averaged values derived from these measurements

were 3.1 ppt for both gases.

1.4.5 Perfluorocarbons, Sulfur

Hexafluoride (SF6), and SF5CF3

Perfluorocarbons (PFCs) and SF6 are very powerful

and long-lived greenhouse gases whose emissions contribute to national inventories of greenhouse gas emissions

that are regulated under the Kyoto Protocol. Trifluoromethylsulfurpentafloride (SF5CF3) is a similar chemical

species whose source is very uncertain and that is not

included in the Kyoto Protocol.

1.4.5.1 CF4 AND CF3CF3

Aoki and Makide (2004) have measured CF4 (FC14) levels of 80-81 ppt in Tokyo during August 2003.

Long-term trend data (1978-1997) have been reported for

CF4 and CF3CF3 (FC-116) in the Cape Meares, Oregon,

air archive, and in flask data from Point Barrow, Alaska,

and Palmer Station, Antarctica (Khalil et al., 2003). The

observed mixing ratios of CF4 and CF3CF3 in 1997 at Cape

Meares were 74 ppt and 2.9 ppt, respectively. These

mixing ratios are consistent with the values given in WMO

2002 (Montzka and Fraser et al., 2003). CF4 increases

have been determined from infrared high spectral resolution solar occultation Fourier transform spectrometer

measurements between 3 and 50 hPa (~20 to 40 km altitude) and latitudes from 50°N to 50°S during 1985, 1992,

1993, 1994, and 2004. These space-based measurements

show that the slowdown in the rate of CF4 accumulation

previously reported from surface measurements through

LONG-LIVED COMPOUNDS

1.49

1975 1985 1995 2005

YEAR

0

50

100

150

3 YEAR SMOOTHED EMISSIONS (Gg/yr)

HFC-134a

0

10

20

30

40

MOLE FRACTION (ppt)

1980 1990 2000 2010

YEAR

(a) HFC-134a (b)

Figure 1-22. (a) Annual global

mean mole fractions for HFC134a derived from the AGAGE

(green line) and NOAA/ESRL

(red diamonds) networks of

ground-based sites (O’Doherty et

al., 2004; Montzka et al., 1996).

The early measurements (before

1995, in blue) are based on

analysis of the air archived in

Australia (Oram et al., 1996). (b)

Annual global emissions of HFC134a smoothed over three-year

intervals, calculated by inverting

the measurements in (a) using

the 12-box model. A lifetime of

14.0 years was used. The green

error bars, which are very small,

indicate the effect of a ±17%

uncertainty in the lifetime.
1997 has propagated to the stratosphere and is continuing

(Rinsland et al., 2006).

Using historical estimated emissions rates of CF4

for the aluminum industry and the semiconductor industry,

Khalil et al. (2003) estimated that the aluminum industry

and the semiconductor industry contributed 33 ppt and 1

ppt, respectively, to the 74 ppt observed in 1997. CF3CF3

is also emitted in aluminum processing. By assuming that

the CF3CF3 and CF4 emissions in 1997 associated with the

aluminum industry maintained a constant ratio through

the years, Khalil et al. concluded that the aluminum

industry emissions are responsible for 2.4 ppt of the

observed CF3CF3 concentration, with the balance assumed

to be from the semiconductor industry.

Harnisch and Höhne (2002) have compared the

estimates of global emissions of CF4 and CF3CF3 from

atmospheric observations (top-down) with a summation

of emissions in national inventories (bottom-up). For CF4,

atmospheric measurements indicate global emissions of

16 Gg/yr (late 1980s) decreasing to 12 Gg/yr (late 1990s),

compared with a decrease from 11to 8 Gg/yr over the same

period based on a summation of national inventories.

Khalil et al. (2003) estimated a decrease from 16 to 10

Gg/yr, respectively, based on global aluminum production

(using an industry-based, varying emission factor)

together with a small (of order 10%) contribution from the

semiconductor industry. Atmospheric measurements of

CF3CF3 indicating constant global emissions of about 2

Gg/yr (late 1980s to late 1990s) compare well with 1.4-

2.2 Gg/yr over the same period based on a summation of

national inventories. An estimate of the annual emissions

of CF3CF3 from Japan (0.19 Gg/yr, 2003) has been made,

based on interspecies ratios (with HCFC-22) observed in

aircraft profiles near Tokyo (Yokouchi et al., 2005).

1.4.5.2 CF3CF2CF3 AND C-C4F8

Long-term trend data (1978-1997) have been

reported for perfluoropropane (CF3CF2CF3) in the Cape

Meares, Oregon, air archive, and in flask data from Point

Barrow, Alaska, and Palmer Station, Antarctica (Khalil et

al., 2003; Culbertson et al., 2004). CF3CF2CF3 mixing

ratios at Cape Meares were <0.1 ppt in 1978, growing to

>0.2 ppt in 1997, corresponding to global emissions of

0.04 Gg/yr in 1977-1982, rising to 0.48 Gg/yr in 1992-

1997. An estimate of the annual emissions of CF3CF2CF3

LONG-LIVED COMPOUNDS

1.50

1980 1985 1990 1995 2000 2005

Year

 

 

0

2

3

4

5

mole fraction (ppt)

HFC-125

AGAGE (Mace Head)

AGAGE (Cape Grim)

UEA (Cape Grim)

1980 1985 1990 1995 2000 2005

Year

 

 

0

2

3

4

5

6

mole fraction (ppt

HFC-152a )

AGAGE (Mace Head)

AGAGE (Cape Grim)

1980 1985 1990 1995 2000 2005

Year

1980 1985 1990 1995 2000 2005

Year

 

0

4

8

12

16

20

24

mole fraction (ppt)

HFC-23

UEA (Cape Grim) Figure 1-23. Monthly means (ppt) and

standard deviations of the AGAGE

measurements of HFC-125, HFC-152a,

and HFC-23 at Mace Head, Ireland

(green) and Cape Grim, Tasmania (red).

Individual measurements of the Australian air archive for Cape Grim analyzed by Oram (1999) are shown in blue.
and perfluorocyclobutane (c-C4F8) from Japan (0.06 and

0.02 Gg/yr, respectively, 2003) has been made, based on

interspecies ratios (with HCFC-22) observed in aircraft

profiles near Tokyo (Yokouchi et al., 2005).

1.4.5.3 SF6

Sulfur hexafluoride (SF6), which is mostly used as

an insulating fluid in transformers, has been measured by a

number of laboratories at Cape Grim, Tasmania, since 1978

via a combination of in situ and flask measurements,

including measurements on the Cape Grim air archive

(Fraser et al., 2004) (Figure 1-24). The long-term growth

rate of SF6 as observed at Cape Grim has increased from

0.1 ppt/yr in the late 1970s to 0.24 ppt/yr in the mid-1990s.

Since then the growth rate has remained relatively constant

at 0.23 ± 0.02 ppt/yr, indicating relatively constant global

emissions (±10%) since 1995. The 2003 annual mean SF6

concentration at Cape Grim was 5.0 ± 0.1 ppt (Fraser et al.,

2004). The global average SF6 mixing ratio in 2003 was

5.2 ppt, growing at 0.2 ppt/yr (Thompson et al., 2004).

The burden of SF6 above the Jungfraujoch

(Switzerland) has continued to accumulate during the past

few years by 3.4 × 1012 molec cm2/yr, which corresponds

to an increase of 4.1%/yr in 2003-2004 (Zander et al.,

2005). Corroboration of the regular accumulation of

atmospheric SF6 at other sites, namely, Kitt Peak (32°N)

and Ny Ålesund (79°N), has been reported by Krieg et al.

(2005). These authors have also shown that the temporal

extrapolation of the best fit to the two decades-long

Jungfraujoch database of SF6 results in atmospheric SF6

loadings in 2050 and 2100 that are significantly lower than

those predicted by the most realistic emission scenarios

(see Figure 1-21 in WMO 2002, Montzka and Fraser et al.,

2003). Satellite observations of SF6 show a globally-averaged upper tropospheric mixing ratio of 4.3 ppt and stratospheric value of 3.5 ppt in September 2002 (Burgess et al.,

2004).

Harnisch and Höhne (2002) have compared the estimates of global emissions of SF6 from atmospheric observations (top-down) with a summation of emissions in

national inventories (bottom-up). The bottom-up estimates of emissions, which are critically dependent on the

accuracy of assumed leak rates from electrical transformers, are 60 ±10% of the top-down estimates. An estimate of the annual emissions of SF6 from Japan (0.15

Gg/yr, 2003) has been made, based on interspecies ratios

(with HCFC-22) observed in aircraft profiles near Tokyo,

which compare favorably to the national emissions for

2001 (0.20 Gg/yr) (Yokouchi et al., 2005).

1.4.5.4 SF5CF3

Trifluoromethylsulfurpentafluoride (SF5CF3) has an

exceptionally large radiative forcing efficiency and was

reported to have an atmospheric mixing ratio of about 0.12

ppt in 1999 (Sturges et al., 2000). A potential source of this

gas has recently been identified (Huang et al., 2005). Under

spark discharge conditions, SF6 does not react with PFCs,

but does so with HFC-23 (CHF3) and HFC-32 (CH2F2),

which may be emitted from the surface of fluoropolymers

in high-voltage equipment. The only other potential source

that has been identified is a byproduct during the manufacture of certain fluorochemicals (Santoro, 2000).

LONG-LIVED COMPOUNDS

1.51

0

2

3

4

5

6

1975 1980 1985 1990 1995 2000 2005

Year

sulfur hexafluoride (SF

 AGAGE

 NOAA-ESRL

 UEA

 UH

 SIO

mole fraction (ppt)

Figure 1-24. Ground-based

measurements of sulfur hexafluoride at Cape Grim, Tasmania.

The data are updates of the following time series: continuous

AGAGE measurements (red;

Fraser et al., 2004); NOAA/ESRL

flask measurements (blue;

Thompson et al., 2004); individual measurements of the

Cape Grim air archive analyzed

by Oram (orange; UEA, 1999)

and by Maiss and Brenninkmeijer

(purple, University of HeidelbergUH, 1998); and Scripps Institution of Oceanography flask

measurements by Vollmer and

Weiss (green; SIO, 2002).
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CHAPTER 2

Halogenated Very Short-Lived Substances

Lead Authors:

K.S. Law

W.T. Sturges 

Coauthors:

D.R. Blake

N.J. Blake

J.B. Burkholder

J.H. Butler

R.A. Cox

P.H. Haynes

M.K.W. Ko

K. Kreher

C. Mari

K. Pfeilsticker

J.M.C. Plane

R.J. Salawitch

C. Schiller

B.-M. Sinnhuber

R. von Glasow

N.J. Warwick

D.J. Wuebbles

S.A. Yvon-Lewis

Contributors:

A. Butz

D.B. Considine

M. Dorf

L. Froidevaux

L.J. Kovalenko

N.J. Livesey

R. Nassar

C.E. Sioris

D.K. Weisenstein

Final Release: February 2007

From Scientific Assessment of Ozone Depletion: 2006

CHAPTER 2

HALOGENATED VERY SHORT-LIVED SUBSTANCES

Contents

SCIENTIFIC SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.1

2.1 INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.5

2.2 SOURCES, DISTRIBUTIONS, AND TRENDS OF VSLS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.9

2.2.1 New Observations of the Distributions and Abundances of VSL Source Gases . . . . . . . . . . . . . . . . . 2.9

2.2.1.1 Bromine and Iodine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.12

2.2.1.2 Chlorine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.13

2.2.2 Terrestrial Emissions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.13

2.2.3 Oceanic Emissions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.15

2.2.3.1 Bromine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.15

2.2.3.2 Iodine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.17

2.2.3.3 Chlorine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.17

2.2.4 Industrial and Other Anthropogenic Emissions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.17

2.2.4.1 Bromine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.17

2.2.4.2 Chlorine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.18

2.2.4.3 New VSL Gases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.18

2.2.5 Sources and Distributions of Inorganic Halogens of Marine Origin . . . . . . . . . . . . . . . . . . . . . . . . . 2.18

2.3 ATMOSPHERIC CHEMISTRY OF VSLS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.19

2.3.1 Removal of Halogen Source Gases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.19

2.3.2 Tropospheric Lifetimes of Halocarbons . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.20

2.3.3 Production and Gas-Phase Removal of VSL Organic Product Gases . . . . . . . . . . . . . . . . . . . . . . . . 2.21

2.3.4 Multiphase Processes Involving Inorganic Bromine and Iodine Compounds . . . . . . . . . . . . . . . . . . 2.21

2.3.4.1 Wet Deposition Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.21

2.3.4.2 Heterogeneous Halogen Activation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.23

2.3.5 Iodine Chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.23

2.4 DYNAMICS AND TRANSPORT IN THE TROPOPAUSE REGION AND IMPLICATIONS FOR VSLS 2.24

2.4.1 Tropical Convection, the TTL, and Tropical Troposphere-to-Stratosphere Transport . . . . . . . . . . . . 2.25

2.4.2 Stratosphere-Troposphere Exchange in the Extratropics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.28

2.4.3 Predictive Modeling of the Tropopause Region . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.30

2.5 CONTRIBUTION OF HALOGENATED VSLS TO STRATOSPHERIC HALOGEN LOADING . . . . . . . . 2.31

2.5.1 Source Gas Injection (SGI) and Product Gas Injection (PGI) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.31

2.5.1.1 Observational Evidence for SGI . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.31

2.5.1.2 Observational Evidence for PGI . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.32

2.5.1.3 Model Estimates of SGI and PGI . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.34

2.5.2 Estimates of VSLS Contributions to Stratospheric Halogens Based on

Measurements of Inorganic Species . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.35

2.5.2.1 Bromine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.36

2.5.2.2 Chlorine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.41

2.5.2.3 Iodine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.42

2.6 POTENTIAL IMPACT OF VSLS ON OZONE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.43

2.6.1 Effects of Halogenated VSLS on Column Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.43

2.6.2 ODPs for Halogenated VSLS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.45
2.7 POTENTIAL FUTURE CONSIDERATIONS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.45

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.46
VERY SHORT-LIVED SUBSTANCES

2.1

SCIENTIFIC SUMMARY

Definition of Halogenated Very Short-Lived Substances (VSLS)

• Very short-lived substances (VSLS) are defined as trace gases whose local tropospheric lifetimes are comparable to,

or shorter than, tropospheric transport time scales, such that their tropospheric distributions are non-uniform. In

practice, VSLS are considered to have atmospheric lifetimes of less than 6 months. We consider only halogenated

VSLS, i.e., those that contain bromine, chlorine, or iodine. We consider VSLS to include very short-lived (VSL)

source gases (SGs), halogenated organic and inorganic product gases (PGs) arising from SG degradation, and other

sources of tropospheric inorganic halogens.

Importance of VSLS for Stratospheric Halogen and Ozone Trends 

BROMINE

• Our quantitative understanding of how halogenated very short-lived substances contribute to halogen levels

in the stratosphere has improved significantly since the last Assessment (WMO, 2003), with brominated

VSLS believed to make a significant contribution to total stratospheric bromine and its effect on stratospheric ozone. Various lines of evidence show that brominated VSLS contribute about 5 ppt (with estimates

ranging from 3 to 8 ppt) to total stratospheric inorganic bromine (Bry):

– Estimates of total stratospheric Bry derived from different observations of bromine monoxide (BrO) relevant

to the late 1990s are about 18 to 25 parts per trillion (ppt). This is greater than the 16 to 17 ppt of bromine

delivered to the stratosphere by “long-lived” brominated source gases (i.e., the halons and methyl bromide,

CH3Br) during this period.

– Measurements of organic brominated very short-lived source gases (VSL SGs) in the tropical upper troposphere amount to about 3.5 ppt. Product gases (from SG degradation) and other sources of tropospheric inorganic bromine may also contribute comparable amounts.

• The inclusion of additional stratospheric Bry from VSLS in models leads to larger ozone destruction at midlatitudes and polar regions compared with studies including only long-lived bromine source gases. In both

regions, the enhanced ozone loss occurs in the lower stratosphere via interactions of this bromine with anthropogenic chlorine. Midlatitude ozone loss is most enhanced during periods of high aerosol loading. Ozone loss

through cycles involving bromine and odd-hydrogen (HOx) is also enhanced at midlatitudes under all conditions,

becoming comparable with ozone loss by HOx cycles alone. This additional amount of stratospheric bromine has

not, to date, been routinely included in model calculations of midlatitude ozone depletion.

• Levels of stratospheric Bry continue to show evidence for a trend that is consistent with that of tropospheric

total bromine. Further studies are required to determine if the recent decline in tropospheric bromine is reflected in

stratospheric bromine abundance.

IODINE

• It is unlikely that iodine is important for stratospheric ozone loss in the present-day atmosphere. There is

little evidence for inorganic iodine, in the form of iodine monoxide (IO), in the lower stratosphere at concentrations

above about 0.1 ppt. This difference in behavior compared with bromine may be partly attributed to the short photochemical lifetime of iodine SGs, their lower abundance, and aerosol uptake of iodine oxides.

CHLORINE

• The sum of the chlorine content from VSL SGs in the tropical upper troposphere is currently estimated to be

about 50 ppt, on the basis of available observations. While a 50 ppt contribution to total inorganic stratospheric

chlorine (Cly) from VSLS represents only 1 to 2% of Cly from long-lived source gases (~3500 ppt), it would represent a significant contribution compared with the background Cly from methyl chloride of about 550 ppt (Chapter 1). 
VERY SHORT-LIVED SUBSTANCES

2.2

• Phosgene (COCl2), which can be produced from chlorinated VSL SGs as well as long-lived chlorine gases, has

been observed at levels of around 20 to 25 ppt in the tropical upper troposphere. The contribution of this gas to

stratospheric Cly (up to about 40 to 50 ppt) is not currently included in most models.

• Analyses of upper stratospheric hydrogen chloride (HCl) measurements from satellite instruments are generally consistent with a possible contribution of VSLS to stratospheric Cly. Further improvements in HCl measurement accuracy are required to quantify this contribution.

Sources and Trends of Halogenated Very Short-Lived Source Gases 

BROMINE

• The majority of known brominated VSL SGs are of natural origin. A few brominated VSLS have almost

exclusively anthropogenic sources, notably n-propyl bromide (n-PB). Some have small contributions from

anthropogenic sources, e.g., certain brominated trihalomethanes.

• Most brominated VSL SGs are emitted from the ocean, with higher emissions in tropical coastal regions. Sea surface supersaturations of some gases, due to production by marine microorganisms, have been found to be elevated in the

tropical open ocean. High concentrations in air have been observed near coasts over tropical and temperate waters.

• Inorganic tropospheric sources of bromine, such as sea salt and volcanoes, may also contribute to inorganic

bromine in the free troposphere. Some fraction of this contribution could be subsequently transported to the

stratosphere.

• There is no evidence for a trend in the sum of brominated VSL SGs during the latter half of the 20th century.

This is according to trend reconstructions from firn air studies, and reflects their predominantly natural origins. The

exceptions are some trihalomethane species, which have increased slightly in the Northern Hemisphere, possibly

due to small anthropogenic sources.

IODINE

• Iodinated VSLS are exclusively of natural origin in the present-day atmosphere, although the possibility of

deliberate future anthropogenic use of new SGs exists. A notable example is the proposed use of trifluoroiodomethane (CF3I).

• Iodinated VSL SGs are emitted from the ocean, with higher emissions in tropical ocean regions. Emissions in

the marine boundary layer are dominated by iodomethanes such as methyl iodide (CH3I) and chloroiodomethane

(CH2ClI), which are produced both by marine organisms and by abiotic photochemical processes in surface waters.

There is some evidence for enhanced marine emissions of iodinated VSL SGs in the tropics and subtropics, and also

in regions of high productivity in coastal zones.

CHLORINE

• Chlorinated VSL SGs originate largely from anthropogenic emissions, although natural sources also contribute. Their principal emissions are, therefore, more likely to be located at northern midlatitudes.

• There is evidence for significant recent declines in concentrations of some chlorinated VSL SGs, notably chloroform, dichloromethane, and tetrachloroethene.

Transport, Lifetimes, and Sinks of Halogenated VSLS 

• The predominant pathways for VSLS transport into the upper troposphere are likely to be in tropical convection regions, co-located with high emissions of VSLS over tropical oceans. The fraction of a halogenated VSL SG

that reaches the stratosphere via the source gas injection (SGI) pathway depends on its local chemical lifetime relative
VERY SHORT-LIVED SUBSTANCES

2.3

to local transport time scales. Local SG lifetimes depend strongly on distributions of photochemical sinks and emission patterns and, as such, this limits the value of using a single “global” lifetime, as has often been used previously. 

• The amount of VSL PGs entering the stratosphere depends on the locations of their tropospheric production

and loss processes. The fraction of halogenated organic and inorganic PGs (produced from the decomposition of

VSL SGs) that enters the stratosphere via the product gas injection (PGI) pathway depends on the location in the

troposphere where the “parent” SG decomposes, and on subsequent loss processes such as (possibly irreversible)

uptake of the PG on to aerosol/cirrus, or uptake of soluble species in precipitating clouds.

• Many more data have become available on Henry’s Law constants, uptake coefficients, and heterogeneous

reactivity for a number of relevant surfaces and for a wide range of PGs. Many uncertainties, however, still

remain. Furthermore, incorporation of this knowledge into large-scale models is often lacking, thus limiting assessment of the location-dependent lifetimes and distributions of VSL PGs. 

• The impact of halogens from VSL SGs on column ozone is likely to be greatest in the extratropics. The dominant transport pathway is likely to be via quasi-horizontal transport from the tropical upper troposphere

into the extratropical lowermost stratosphere. While transport time scales via this pathway, and back to the troposphere, are relatively short (a few months), the short chemical lifetime of VSL SGs means that they will be fully

degraded into inorganic halogen before they can be transported back into the troposphere. Quasi-horizontal transport from the tropical upper troposphere is highly seasonal, being much larger in the summer than winter. 

• Recent model simulations continue to suggest that, depending on lifetime and location of emissions, less than

1% of the VSL SG flux emitted at the surface enters the stratosphere via the SGI pathway, and a larger

amount (up to a few percent) enters via the PGI pathway. Studies are limited, however, by the ability of current

global models to accurately simulate deep convection and transport into the lower stratosphere.

Ozone Depletion Potentials (ODPs) of Halogenated VSL SGs 

• There are no new evaluations for the ODP of n-propyl bromide (0.1 for tropical emissions, and 0.02-0.03 for

emissions restricted to northern midlatitudes).

• New analyses suggest ODPs for CF3I ranging from 0.011 to 0.018 for surface emissions at midlatitudes and in the

tropics, respectively, which are higher than found previously (<0.008). The model studies do not include uptake of

iodine on aerosols, the inclusion of which could result in lower calculated values. If CF3I were to be used for fire

fighting/inhibition on aircraft, then it could be emitted at higher altitudes and the corresponding ODP would be larger. 

• New analyses confirm previous estimates that the ODP of a chlorinated VSL SG with a lifetime of ~25 days,

one chlorine atom, and similar molecular weight to CFC-11, is about 0.003.

Future Considerations for Halogenated VSLS 

• Possible future changes in anthropogenic VSLS. If anthropogenic emissions increased, or if presently unused

halogenated VSL SGs were to come into widespread commercial use, then halogenated VSLS would become of

increased importance in affecting the future behavior of stratospheric ozone.

• Delivery of VSLS to the stratosphere may change in the future in response to circulation changes. The impact

of natural halogenated VSLS might also be influenced by changes in the atmospheric circulation, which could, for

example, increase the rate of delivery of VSL SGs and PGs into the stratosphere.

• Natural VSLS emissions may respond to future changes in climate processes. Natural sources could respond to

changes in, for example, carbon dioxide (CO2), land use, wind speed, and temperature. Our knowledge about these

potential effects, and many other relevant feedbacks, is very limited at present.

2.1 INTRODUCTION

An important goal of previous Assessments was to

quantify the impact of halogen-containing source gases

on stratospheric ozone (e.g., WMO, 2003, 1999). In the

last Assessment (WMO, 2003), the possible contribution

of very short-lived substances (VSLS) was examined in

detail in Chapter 2 (Ko and Poulet et al., 2003) and an

approach for determining their Ozone Depletion Potentials

(ODPs) was presented. VSLS are defined as substances

that have atmospheric lifetimes comparable to, or less than,

average tropospheric transport time scales of about 6

months. We consider halogenated very short-lived (VSL)

organic source gases (SGs), and the halogenated organic

and inorganic product gases (PGs) arising from the

atmospheric degradation of VSL SGs, and also inorganic

tropospheric halogen sources such as sea salt. Note that,

in contrast to WMO (2003), sulfur-containing VSLS (e.g.,

dimethyl sulfide (DMS), and sulfur dioxide (SO2)) are

dealt with in Chapter 5 of this Assessment. Halogen atoms

bound in SGs and intermediate degradation products are

referred to as organic halogen, and the final products are

referred to as inorganic halogen (containing no carbon)

(e.g., hydrogen chloride (HCl), hydrogen bromide (HBr),

chlorine monoxide (ClO), bromine monoxide (BrO), and

iodine monoxide (IO)). Some important acronyms used

in this chapter are defined in Box 2-1 below.

This chapter assesses the contribution of halogenated VSLS to the halogen loading in the stratosphere,

as well as their possible contribution to past, present, and

future stratospheric ozone loss, and considers the ODPs of

substances that are, or might be, released as a result of

human activities. Table 2-1 provides a list of the VSL SGs

discussed in this chapter, together with an estimation of

their local photochemical lifetimes and an indication of

whether these gases have predominantly natural or anthropogenic origins and/or are proposed for new commercial

applications. The majority of currently observable VSL

SGs are thought to be partly or wholly of natural origin.

A few halogenated VSLS have predominantly anthropogenic origins (e.g., n-propyl bromide (n-C3H7Br), 1,2-

dibromoethane (CH2BrCH2Br), dichloromethane (CH2Cl2),

tetrachloroethene (C2Cl4), and, to a lesser extent, chloroform (CHCl3)). Currently, among the compounds listed in

Table 2-1, only bromochloromethane (CH2BrCl) is considered by the Montreal Protocol. In the light of observed

VERY SHORT-LIVED SUBSTANCES

2.5

Box 2-1. Definition of Acronyms in Chapter 2

VSLS Very Short-Lived Substances – organic and inorganic gases with lifetimes of less than 0.5 years

VSL Very Short-Lived 

SG/SGI Source Gas / Source Gas Injection – refers to a halogenated organic source gas and its injection into

the stratosphere

PG/PGI Product Gas / Product Gas Injection – refers to the sum of the halogenated organic and inorganic

degradation products for a particular source gas and their injection into the stratosphere

CFC Chlorofluorocarbon

HCFC Hydrochlorofluorocarbon

Brx Inorganic bromine oxides and radicals (e.g., BrO, atomic bromine (Br), bromine nitrate (BrONO2))

Bry Total inorganic bromine (e.g., HBr, Brx) resulting from degradation of bromine-containing organicsource gases (halons, methyl bromide, VSLS), and natural inorganic bromine sources (e.g., volcanoes, sea salt, and other aerosols)

Bry

VSLS The component of stratospheric Bry from the degradation of brominated organic VSL SGs and tropospheric inorganic bromine sources (also called “additional” stratospheric Bry)

Bry

CH3Br+Halons The component of stratospheric Bry from the degradation of methyl bromide (CH3Br) and halons

Cly Total inorganic stratospheric chlorine (e.g., HCl, ClO) resulting from degradation of chlorine-containing source gases (CFCs, HCFCs, VSLS), and natural inorganic chlorine sources (e.g., sea

salt and other aerosols) 

Cly

VSLS The component of stratospheric Cly from the degradation of chlorinated organic VSL SGs and tropospheric inorganic chlorine sources

Iy Total inorganic stratospheric iodine (e.g., IO, iodine dioxide (OIO), hypoiodous acid (HOI)) resulting

from degradation of iodine-containing source gases (VSLS), and natural inorganic sources (e.g.,

sea salt and other aerosols)

ODP Ozone Depletion Potential

EESC Equivalent Effective Stratospheric Chlorine
and predicted decreases in anthropogenic chlorofluorocarbons (CFCs), halons, etc., it is important to quantify the

contribution of these VSLS to the natural background

halogen loading in the stratosphere. Possible changes in

the emissions of naturally occurring compounds, for

example as a result of changing climatic factors, also need

to be assessed. 

In order to quantify the contribution of halogenated

VSLS to halogen loading and ozone loss in the stratosphere, it is necessary to quantify the amount of VSLS

entering the stratosphere directly via the SG injection

(SGI) pathway, the degradation of SGs in the troposphere,

transport of organic and inorganic PGs into the stratosphere via the PG injection (PGI) pathway, and also injection of inorganic tropospheric halogen. The contribution

to, for example, total inorganic bromine (Bry) in the stratosphere, which derives from the sum of VSL SGs (via the

SGI and PGI pathways) and tropospheric inorganic

halogen, is denoted Bry

VSLS. The impact of such VSLSderived inorganic halogen on stratospheric ozone loss will

depend on the abundance of different halogen radicals and

their efficiency at destroying ozone. Note that one atom

of iodine is much more efficient at removing ozone than

bromine, which in turn is more efficient than chlorine, as

expressed by so-called α-factors (see Chapter 8 and

Section 2.6.1 for further discussion).

Figure 2-1 shows a schematic of the processes influencing the transport, degradation, and loss of VSL SGs

and PGs in the troposphere and stratosphere. It is an

updated version of Figure 2-2 in Chapter 2, WMO (2003).

VERY SHORT-LIVED SUBSTANCES

2.6

Table 2-1. Halogenated VSL source gases considered in this Assessment. New additions since WMO

(2003) are shown in boldface; other compounds were assessed in WMO (2003).

Source Gas Formula Local Lifetime a Atmospheric

(days) Source b

Bromochloromethane CH2BrCl 150 N

Dibromomethane CH2Br2 120 N

Bromodichloromethane CHBrCl2 78 N(A)

Dibromochloromethane CHBr2Cl 69 N(A)

1,2-Dibromoethane

(ethylene dibromide, EDB) CH2BrCH2Br 55 A

Bromoethane (ethyl bromide) C2H5Br 34 A

Tribromomethane (bromoform) CHBr3 26 N(A)

1-Bromopropane (n-propyl bromide) n-C3H7Br 13 A

Phosphorous tribromide (“PhostrEx”) PBr3 <0.01 Ac

Iodomethane (methyl iodide) CH3I 7N

Trifluoroiodomethane CF3I 4Ac

Iodoethane (ethyl iodide) C2H5I 4N

2-Iodopropane (isopropyl iodide) i-C3H7I 1.2 N

1-Iodopropane (n-propyl iodide) n-C3H7I 0.5 N

Chloroiodomethane CH2ClI 0.1 N

Bromoiodomethane CH2BrI 0.04 N

Diiodomethane CH2I2 0.003 N

Trichloromethane (chloroform) CHCl3 150 AN

Dichloromethane (methylene chloride) CH2Cl2 140 A(N)

Tetrachloroethene (perchloroethylene, PCE) C2Cl4 99 A

1,2-Dichloroethane CH2ClCH2Cl 70 A

Chloroethane (ethyl chloride) C2H5Cl 30 AN

Trichloroethene (trichloroethylene; TCE) C2HCl3 4.6 A a Local lifetimes are calculated as (τlocal)

-1 = (τOH)-1 + (τJ)-1, with [OH] = 1 × 106 molec cm-3, T = 275 K, and a globally and seasonally averaged solar

flux for 5 km altitude. See Section 2.3.2. b N ; entirely natural atmospheric sources; A ; entirely anthropogenic atmospheric sources; AN ; combination of natural and anthropogenic atmospheric sources (a minor component shown in parentheses). c Proposed application.
In terms of chemical processes affecting VSLS, it is the

degradation and loss of intermediate products, and inorganic multiphase heterogeneous chemistry, that remain

the largest uncertainties in our knowledge about the fate

of VSLS. The most efficient transport pathway is in the

tropics, where there is rapid transport of air from the surface to the tropical tropopause layer (TTL) by deep convection, followed by transport of a small fraction into the

stratosphere. Quasi-horizontal transport of air from the

TTL into the lowermost stratosphere (LMS), as well as

direct transport of air from the surface at midlatitudes

either by frontal processes or deep convection, also need

to be considered.

Given that VSL SGs and PGs have, by definition,

shorter photochemical lifetimes than the large-scale transport time constant in the troposphere of roughly 6 months,

mixing ratios (molar fractions) of both SGs and PGs

(unless the latter are long-lived) are not uniform in the troposphere, and significant vertical as well as horizontal gradients can exist. Therefore, the amount of halogen entering

the stratosphere may be sensitive to, among other factors,

the location and timing of emission of VSL SGs, chemical

degradation to produce organic and inorganic PGs, tropospheric transport processes, and location/altitude of

removal of inorganic degradation products by wet deposition or other heterogeneous processes. Measured ratios of

boundary layer (BL) to upper tropospheric (UT) mixing

ratios of VSL SGs can provide an indication of the fraction available to be transported into the stratosphere. It

has been noted previously that brominated and iodinated

VSL SGs have been observed at mixing ratios of a few

parts per trillion (ppt or pmol mol-1) at the surface (WMO,

VERY SHORT-LIVED SUBSTANCES

2.7

5 km 

TTL

 SG

 SG

 SG SG PG 

cirrus

 Xy ozone

 loss

OH OH

OH

1-2 km 

11 km 

20 km 

28 km 

boundary layer

EQUATOR

 SG

washout

washout

Key

 VSLS - Very Short-Lived (VSL) Substances

 SG - organic VSL Source Gas

 PG - VSL Product Gas: halogenated 

 organic/ inorganic degradation products

 Xy - inorganic halogen, X=Cl, Br, I from VSL

 SG degradation or tropospheric inorganic

 halogen

 SGI - Source Gas Injection

 PGI - Product Gas Injection

 TTL - Tropical Tropopause Layer

 ExTL - Extra-tropical Tropopause Layer

 LMS - Lowermost Stratosphere

Chemical and Dynamical Processes Affecting VSLS

tropical 

convection

subsidence

crosstropopause

exchange clouds

large-scale 

transport

frontal

uplift

ExTL

G

 PG G

mid-latitude PG 

convection

 Xy

PG SG

 PG 

 PG 

 SG

 Xy

 Xy

PGI

PGI

PGI SGI

SGI

SGI

quasi-horizontal LMS transport

SGI

stratospheric overworld

troposphere

VSLS

VSLS VSLS

SLS

VSLS

 SG SG

Figure 2-1. Schematic showing principal chemical and dynamical pathways transporting VSL source gases

(SG) and organic/inorganic product gases (PG). Stratospheric halogen loading is maintained by transport of

source gases followed by their degradation in the stratosphere (the SGI pathway), and transport of intermediate products and inorganic halogens produced in the troposphere (the PGI pathway). Tropospheric inorganic halogens can derive from degradation of VSL SGs, or from inorganic halogen sources. This figure is an

update to Figure 2-2 in WMO (2003); courtesy of K.S. Law (Service d’Aéronomie/CNRS, France) and P.H.

Haynes and R.A. Cox (University of Cambridge, U.K.).
2003). An estimate was also given for the tropospheric

abundance of chlorinated VSL SGs ranging from 50 to 100

ppt. This value also included phosgene (carbonyl dichloride, COCl2), produced from the degradation of VSL SGs

and long-lived chlorine gases. Apart from phosgene, very

few measurements of organic VSL PGs are available in

the free troposphere or stratosphere. Measurements of

inorganic halogen are available and represent the final

degradation products of halogen-containing sources gases

(e.g., CFCs, halons, VSLS) as well as direct emission of

inorganic halogens (e.g., sea salt, volcanoes). 

The different radical families that are important for

ozone loss in the midlatitude stratosphere are shown in

Figure 2-2. At the levels of total inorganic iodine used in

these model calculations (Iy = 0.1 ppt), which are representative of upper limits based on rather limited observaVERY SHORT-LIVED SUBSTANCES

2.8

Figure 2-2. Calculated fractional contribution to odd oxygen (Ox) loss by catalytic cycles involving nitrogen

(NOx), hydrogen (HOx), chlorine (ClOx), bromine (BrOx), and iodine (Iy Fast and Slow; see below) radicals for

March 1995, 32°N. Jet Propulsion Laboratory (JPL) kinetics (Sander et al., 2006) were used, except for the rate

constant of IO + HO2, for which the expression given by Knight and Crowley (2001) representing all available

data, 1.4 × 10–11exp(554/T), was used (this is the same rate constant as given by the February 2006 International

Union of Pure and Applied Chemistry (IUPAC) evaluation, posted at http://www.iupac-kinetic.ch.cam.ac.uk/).

Total inorganic stratospheric iodine was set equal to 0.1 ppt, the upper limit reported by Bösch et al. (2003) and

Butz (2006). For the left-hand panel, other model inputs are the same as given in Table 4 of Wennberg et al.

(1997), which assumed stratospheric Bry derived only from the breakdown of CH3Br and halons. For the righthand panel, 5 ppt has been added to Bry at all altitudes, representing Bry

VSLS (see text and Section 2.5). The

ClOx curve represents loss from the ClO + O and ClO + HO2 cycles, plus other minor cycles that involve ClO, but

not BrO or IO. The BrOx curve represents loss from the BrO + ClO and BrO + HO2 cycles, plus other minor

cycles that involve BrO, but not IO. Two model curves are given to represent loss by all cycles that involve IO in

the rate-limiting step. The “Iy Fast” curve uses the rate constant for IO + BrO given by Rowley et al. (2001) and

assumes all products of this reaction lead to catalytic loss of ozone. The “Iy Slow” curve uses the JPL rate constant for IO + BrO and assumes that production of OIO results in a null cycle. These two curves are shown to

illustrate the sensitivity of iodine-induced ozone loss to the rate constant of IO + BrO and to the fate of OIO (e.g.,

photolysis of OIO to O + IO leads to a null cycle, whereas photolysis to I + O2 leads to O3 loss). If the temperature-independent rate constant for IO + HO2 given by JPL is used, the fractional contribution to ozone loss by

iodine in the LMS is about a factor of two less than illustrated. Adapted from Wennberg et al. (1997); see also

the extensive discussion of iodine kinetics in Bösch et al. (2003).
tions in the stratosphere (Bösch et al., 2003; Butz, 2006),

loss of ozone by iodine is unlikely to be a driving factor

for present-day midlatitude ozone loss. Note that the efficiency of ozone loss by iodine is sensitive to unresolved

details of the photochemical mechanism, regarding the

rate constant of the IO + BrO reaction and different

decomposition products resulting from iodine dioxide

(OIO) photolysis (“Iy Fast” versus “Iy Slow” in Figure 2-

2; see caption). The likely relatively small contribution

of iodine to stratospheric ozone loss is in contrast to

bromine, which has been observed regularly in the stratosphere at levels significantly higher than can be supplied

solely by the longer lived brominated gases (CH3Br and

halons) and at sufficient concentrations in the lower stratosphere to make a difference to present-day ozone.

A key question for this chapter, then, is to what

extent brominated VSLS contribute to stratospheric Bry.

The contributions of chlorinated and iodinated VSLS to

stratospheric Cly and Iy budgets are also considered.

Figure 2-3, which is an updated version of Figure 1-8,

Chapter 1, WMO (2003), shows Bry estimated from balloonborne BrO observations adjusted to the time when the

measured air masses were last in the troposphere (see

Figure 2-3 caption for further details). In the late 1990s,

this amounted to about 18 to 25 ppt of Bry. Also shown

are observed trends in surface mixing ratios of brominecontaining halons and methyl bromide (CH3Br) since 1980

(see Chapter 1 for details). Note that the trends of CH3Br

have not been corrected to account for vertical gradients

in the troposphere, so that this contribution to stratospheric

Bry may be overestimated, and therefore Bry

VSLS underestimated. It can be seen that there is an apparent discrepancy between Bry from halons plus CH3Br, and Bry

inferred from BrO data. In WMO (2003), this discrepancy, denoted “additional” stratospheric bromine, was

attributed to either direct injection into the stratosphere of

inorganic bromine from the troposphere as well as from

degradation of organic VSLS, or to possible calibration

errors in the measurement of either BrO or the organic

source gases. Since the last Assessment, there have been

many studies focused on quantification of this “additional”

bromine (Bry

VSLS). Figure 2-3 shows possible contributions of Bry

VSLS varying between 3, 5, and 7 ppt, as indicated by the thin blue lines. If 5 ppt of Bry

VSLS is incorporated into a photochemical model, then ozone loss in the

lower stratosphere from reactions involving BrO occurs at

a rate approaching that due to the traditionally considered

hydrogen oxide cycles alone (see Figure 2-2). 

The structure of this chapter is as follows. Section

2.2 discusses the emissions, observations, and trends of

VSL SGs, including vertical distributions at different locations, and their likely mixing ratios in the upper troposphere. Section 2.3 provides an update to our knowledge

about the atmospheric chemistry of halogenated VSL SGs

and PGs (organic and inorganic), wet removal of PGs, and

a discussion about the geographical variability of chemical lifetimes. The role of heterogeneous chemistry in the

inorganic halogen budget is also discussed. Section 2.4

revisits our understanding of transport processes that are

important for the transport of VSL SGs and PGs to the

stratosphere via the SGI and PGI pathways in both the

tropics and extratropics. The ability of current models to

simulate the distributions of short-lived gases is also examined, since these are the tools used to estimate the impact

of VSLS on stratospheric halogen loading, ozone loss, and

their ODPs. Section 2.5 presents an assessment of the contribution of VSL SGs and PGs to stratospheric halogen

budgets. Model estimates of the SGI and PGI fraction

entering the stratosphere are also discussed. The possible

contribution of VSLS (bromine) to stratospheric ozone

loss, and an update of ODPs for VSLS, are discussed in

Section 2.6. Finally Section 2.7 summarizes recent evidence for, and speculates on, possible future changes in

halogenated VSLS that could affect future stratospheric

halogen loadings and ozone loss.

2.2 SOURCES, DISTRIBUTIONS, AND TRENDS

OF VSLS

2.2.1 New Observations of the Distributions

and Abundances of VSL Source

Gases

Since the last Assessment (WMO, 2003), a number

of shipboard, atmospheric, firn air, and coastal observations have provided new information that either adds to or

constrains earlier estimates of atmospheric mixing ratios

of VSL SGs, which are known to exhibit significant geographical variation. Table 2-2 provides updated information since the last Assessment on VSL SG mixing ratios in

the atmosphere and, in particular, it assesses the subset of

relevant observations required to arrive at an estimated

range of mixing ratios in the tropical marine boundary

layer and tropical upper troposphere for each source gas.

It suggests mean tropical boundary layer mixing ratios of

8.4, 1.2, and 81 ppt for total bromine, iodine, and chlorine, respectively, in VSL SGs; and corresponding mixing

ratios in the tropical upper troposphere of 3.5, 0.08, and

55 ppt. It is clear from these observations that many of

the chlorinated and brominated compounds are delivered

efficiently from the boundary layer to the upper troposphere. Bromoform (CHBr3) and dibromomethane

(CH2Br2) contribute the most VSL organic bromine

VERY SHORT-LIVED SUBSTANCES

2.9
VERY SHORT-LIVED SUBSTANCES

2.10

Figure 2-3. Trends (solid lines) of expected total inorganic stratospheric bromine (Bry) due to measured tropospheric abundances of bromine from methyl bromide (CH3Br; thick blue lines), the sum of CH3Br plus halons

(thick purple lines), and this sum together with assumed additional bromine (3, 5, and 7 ppt) derived from

VSLS and/or inorganic tropospheric bromine sources (thin blue lines). This figure is an update of Figure 1-8

from the last Assessment (WMO, 2003). The smooth thick blue and purple lines prior to 1995 are based on a

combination of Antarctic firn air reconstructions and ambient air measurements from a number of studies,

adjusted to represent mean global abundances (references in WMO, 2003), and updated here with mean

ambient air measurements (non-smooth blue and purple lines post-1995) from NOAA global monitoring stations (Montzka et al., 2003; updated). There may be small calibration differences between the different studies

(see, for example, Fraser et al., 1999). The points (open and filled squares) are total inorganic bromine derived

from stratospheric balloon measurements of BrO and photochemical modeling to account for BrO/Bry partitioning. The filled black squares are from the slopes of Langley plot BrO observations above balloon float altitude, while the open squares are lowermost stratospheric BrO measurements (Dorf, 2005, and Dorf et al.,

2006a, and references therein; together with more recent data as described in Section 2.5.2.1). Bold/faint

error bars correspond to the precision/accuracy of the estimates, respectively (Dorf, 2005). For stratospheric

data, the date corresponds to the time when that air was last in the troposphere, i.e., sampling date minus

mean age of air in the stratosphere. No correction has been made for gradients of CH3Br in the troposphere

(see text for discussion). Preindustrial levels of CH3Br of 5.8 and 5.5 ppt have been inferred by Saltzman et al.

(2004) from Antarctic ice cores and by Trudinger et al. (2004) from Antarctic firn air, respectively, whereas

preindustrial levels of the halons have been reported to all be below detection limits in Arctic and Antarctic firn

air by Reeves et al. (2005). Adapted from Dorf (2005); see also Dorf et al. (2006b).
VERY SHORT-LIVED SUBSTANCES

2.11

Table 2-2. Reported mixing ratios of VSL source gases in the troposphere (global and hemispheric

values), at 10 km in the tropics, and best estimates from this Assessment of median mixing ratios in

the tropical marine boundary layer [MBL] (<1 km), and tropical upper troposphere [UT]. All mixing ratio

values are in parts per trillion (ppt). New information since WMO (2003) is indicated in boldface.

Species Reported Reported Estimated Tropical Estimated Estimated

Tropospheric 10-km [MBL] Mixing Ratio Tropical [UT] Ratio of

Mixing Ratio a Tropical and Range Mixing Ratio and Tropical

Mixing Ratio a Range [UT]/[MBL]

CH2BrCl 0.47 (0.38-0.59)c 0.32 (0.26-0.35)c 0.7

CH2Br2 0.8-3.4 0.6-0.9 1.1 (0.7-1.5)b,c,f 0.9 (0.7-1.0)b,c,e 0.8

CHBr2Cl 0.1-0.5 0.04-0.11 0.30 (0.06-0.76)b,g,i 0.08 (0.03-0.12)b,c 0.3

CHBrCl2 0.12-0.6 0.04-0.11 0.33 (0.14-0.91)b,c,g,j 0.12 (0.05-0.15)b,c 0.4

CHBr3 0.6-3.0 0.4-0.6 1.6 (0.5-2.4)b,c,f,g 0.37 (0.13-0.7)b,c,e 0.2

CH3I 0.1-2.0 0.05-0.2 0.8 (0.3-1.9)b,c,f 0.08 (0.02-0.18)b,c 0.1

CH2ClI 0.35g

CH2Cl2 17.5 (9-39)b,c,f 13.2 (9-19)b,c,e 0.75

CHCl3 NH, 10-15

12.4 (9.8-14.5)k

SH, 5-7

8.0 (6.5-9.1)k

C2HCl3 NH, 1-5

SH, 0.01-0.1

C2Cl4 NH, 5-15

5.3 (3.3-7.3)k

SH, 0.7-1.5

1.5 (1.1-1.6)k

C2H5Cl NH, 2.6l

SH, 1.6l

CH2ClCH2Cl NH, 20-40

SH, 5-7

COCl2 22.5 (20-25)m

Total Cl 81 (75-99)n 55 (52-60)n,o

Total Br 8.4 (6.9-9.6)n 3.5 (3.1-4.0)n

Total I 1.2 (0.7-2.3)n 0.08 (0.02-0.18)n

NH = Northern Hemisphere, SH = Southern Hemisphere.

a Kurylo and Rodríguez et al. (1999), except as noted. b Transport and Chemical Evolution Over the Pacific (TRACE-P), Pacific Exploratory Missions (PEM) “Tropics A”/“Tropics B,” Tropospheric Ozone

Production about the Spring Equinox (TOPSE); Blake et al. (1999a,b, 2001, 2003a,b). c 10-14 km tropical upper troposphere data from the Biomass Burning and Lightning Experiment-B (BIBLE-B) (Choi et al., 2003; Kondo et al., 2002). d TRACE-P data only. e 10-17.5 km non-stratospheric data from the Preliminary Aura Validation Experiment (preAVE; ozone < 75 ppt) (ftp://espoarchive.nasa.gov/archive/-

pre_ave/data/wb57). f Cruise data (Butler et al., 2006). g Chuck et al. (2005). h TRACE-P and PEM-Tropics A only (see above). i High value from Chuck et al. (2005) increased mean from 0.15 to 0.30. 

j High value from Chuck et al. (2005) increased mean from 0.19 to 0.33. k 2002-2004 surface (Simpson et al., 2004; Blake, 2005). l Low et al. (2003). m Toon et al. (2001). n Sum of medians and the root mean square (rms) range determined from data for individual compounds shown in the table. o COCl2 is not included in Total Cl, as part of it may be derived from long-lived species (see text).

3.1 ±0.7 7.8 (5.2-13.3)b,c,f,k 6.0 (4.8-7.5)b,c,e 0.78

1-3 1.8 (1.2-3.8)b,c,f,k 1.3 (0.9-1.6)b,c,e 0.7

0-0.1 0.5 (0.05-2)b,c,d 0.14 (0.02-0.3)b,c,d,e 0.3

5.0 (2.7-5.9)d 1.5 (1.0-1.8)d 0.3

14.9 ±1.1 3.7 (0.7-14.5)b,c,h 1.8 (0.7-3.3)b,c,e,h 0.5
(>80%) to both the marine boundary layer and the upper

troposphere. Similarly chloroform (CHCl3) and dichloromethane (CH2Cl2) contribute most of the VSL organic

chlorine (~80%) to the marine boundary layer and the

upper troposphere. There are few observations of measurable iodine source gases in the upper troposphere or

lower stratosphere.

As discussed later in this chapter, many of the compounds listed in Table 2-2 are converted to inorganic

forms, but few of these have been measured in the upper

troposphere. A notable exception is phosgene (COCl2)

which may account for an additional c. 45 ppt of chlorine

(Cl) in the tropical upper troposphere although, as discussed in Section 2.3.3, part of the phosgene might be associated with the decomposition of long-lived source gases

in either the troposphere or stratosphere.

The abundances, distributions, and sources of these

short-lived, potentially ozone-depleting gases are considered in more detail below.

2.2.1.1 BROMINE AND IODINE

Atmospheric mixing ratios of many brominated and

iodinated VSLS are typically higher where marine influence is significant. They tend to be highest near coastal

areas, oceanic fronts, and in the tropics or subtropics (e.g.,

Quack and Wallace, 2003; Carpenter, 2003; Chuck et al.,

2005; Yokouchi et al., 2005b; Butler et al., 2006). The

tropics, including coastal areas, are significant in that they

comprise the main region where deep convection occurs,

and are therefore the most likely location for transport of

gases emitted at the surface to be carried to the upper troposphere and lower stratosphere.

Carpenter et al. (2003) compared atmospheric

measurements of a range of reactive organic halogens at

Mace Head, Ireland (September, 1998) and Cape Grim,

Tasmania (January-February, 1999). Mixing ratios of

methyl iodide (CH3I), chloroiodomethane (CH2ClI),

dibromochloromethane (CHBr2Cl), CHBr3, and CH2Br2

at Cape Grim were on average only 25-50% of those at

Mace Head, Ireland, which is more strongly influenced by

emissions from local macroalgae (seaweeds). Average

total VSL bromine abundances at Cape Grim and Mace

Head during these campaigns were 18 ppt and 39 ppt as

bromine (Br), respectively. These amounts are much

higher than the global averages of 2.4-3.5 ppt reported in

the last Assessment, and reflect the intense emissions from

coastal waters. Bromoform contributed 45% and 54% of

the total VSL bromine at these two sites, respectively, and

accounted for most of the variability. Peters et al. (2005)

noted extremely high values of CH3I, CHBr3, and other

VSLS over seaweed beds in northern Europe. These

values, as high as 1830 ppt for CH3I and 393 ppt for

CHBr3, have not been reported by any other investigations,

the closest being 40 ppt of CHBr3 reported by Yokouchi

et al. (2005b). Other iodinated VSLS have been observed

in such coastal environments, including diiodomethane

(CH2I2), bromoiodomethane (CH2BrI), iodoethane

(C2H5I), and 1-iodopropane (n-C3H7I), but at smaller concentrations (Carpenter et al., 2003; Peters et al., 2005).

Quack and Wallace (2003), summarizing global air

and seawater mixing ratios of CHBr3 by oceanic region,

showed that background mixing ratios of CHBr3 in marine

boundary layer air are generally in the range 0.5-1.5 ppt.

This is consistent with data in Butler et al. (2006), but it

is about half of the range of 1.3-3.9 ppt reported by Chuck

et al. (2005) for the open ocean. This difference may have

resulted from natural temporal and spatial variability, but

it may also be due to differences in calibration scales,

which underscores the need for intercalibration of these

gases among laboratories. Quack and Wallace (2003) and

Butler et al. (2006) showed high and variable coastal

marine boundary layer CHBr3 mixing ratios ranging

between 1 and 8.3 ppt, and levels in the open ocean

ranging between 0.4 and 2.5 ppt. These levels are consistent with the previous Assessment (WMO, 2003).

Yokouchi et al. (2005b) recently reported over 40 ppt of

CHBr3 (i.e., >120 ppt as Br) in the first study of air over

tropical coastal waters, suggesting that such areas may be

particularly important in delivering halogenated VSLS to

the stratosphere.

Seasonality in mixing ratios of VSLS is evident

over most of the ocean (Butler et al., 2006) and is particularly marked in coastal areas. Cohan et al. (2003) and Cox

et al. (2005) reported high frequency, in situ observations

at Cape Grim. Average background levels of CH3I for

1998-2001 were 1.4 ppt (Cox et al., 2005). Mixing ratios

peaked during the summer (amplitude 0.47 ppt) despite

faster photolytic loss, suggesting that the local oceanic

emissions were a significant source of this gas (Cohan et

al., 2003). Carpenter et al. (2005) found a strong, broad

seasonal range of CHBr3 (1.8-5.3 ppt) at Mace Head,

noting higher values in the summer, and identifying both

terrestrial and marine sources.

Although organic VSL source gases predominate

in the marine boundary layer, it is possible that small

amounts of inorganic bromine or iodine in the boundary

layer could also be rapidly convected to the lower stratosphere. Recent observations suggest that bromine

monoxide (BrO) in the marine boundary layer, whether

from organic or inorganic sources, is on the order of 0-2

ppt (Leser et al., 2003; Saiz-Lopez et al., 2004a), with

higher values observed locally. High amounts of iodine

monoxide (IO) are routinely observed in the marine

VERY SHORT-LIVED SUBSTANCES

2.12
boundary layer near macroalgae, with values as high as 7-

8 ppt having been reported along the coasts of Ireland and

France (Saiz-Lopez and Plane, 2004; Saiz-Lopez et al.,

2004b; Peters et al., 2005). Saiz-Lopez and Plane (2004)

suggest that emission of molecular iodine (I2) from macroalgae may be the dominant source of IO. Although VSLS

produced at midlatitudes are not as likely to be convected

into the upper troposphere as are VSLS produced in tropical regions, this still underscores the potential contribution of VSLS from coastal regions. 

2.2.1.2 CHLORINE

Cox et al. (2005) reported background mixing ratios

of 6.3 ppt for CHCl3 at Cape Grim (1998-2000), and

Simmonds et al. (2006) reported 8.7 ppt for CH2Cl2 also

at Cape Grim (1998-2004). Observed elevated levels of

CHCl3 (55 ppt) were attributed to local natural sources,

while elevated levels of CH2Cl2 (up to 70 ppt above background) were attributed to anthropogenic sources. Data

referenced in Butler et al. (2006) show higher tropical

marine air mixing ratios of CH2Cl2 (18.8 ppt), but similar

tropical mixing ratios of CHCl3 (7.4 ppt) to those observed

at Cape Grim. Numerous studies (e.g., Prinn et al., 2000;

Thompson et al., 2004; Simmonds et al., 2006) show

higher Northern Hemispheric mixing ratios of both of

these gases, which suggest greater emissions in the

Northern Hemisphere and are consistent with a large

anthropogenic source. 

Low et al. (2003) reported the first calibrated measurements of atmospheric chloroethane (C2H5Cl). The

median mixing ratios observed at a clean California coastal

site were 3.3 and 0.3 ppt for C2H5Cl and C2H5Br, respectively. No significant correlation was found between the

atmospheric concentrations of the methyl halides and the

ethyl halides, suggesting different source-sink relationships. These results, combined with data from the

Transport and Chemical Evolution over the Pacific experiment (TRACE-P), suggest an average global mixing ratio

for C2H5Cl of 2.6 ppt. Using these data with an average

hydroxyl radical (OH) lifetime of ~1 month, the global

burden and required annual source of C2H5Cl to the atmosphere are 25 Gg (1 Gg = 109 g), and 300 Gg yr-1, respectively (Redeker et al., 2003). 

2.2.2 Terrestrial Emissions

Terrestrial emissions of brominated or iodinated

VSLS are generally thought to be smaller than from the

ocean (Table 2-3). Emissions from many terrestrial

sources, such as rice cultivation, fires, and wastewater

effluents, are anthropogenically influenced. Several evaluations of fluxes from rice paddies, using both new data

and modeling techniques, indicate that rice paddies remain

a significant source of CH3I to the atmosphere. Lee-Taylor

and Redeker (2005) re-evaluated the global fluxes of CH3I

from rice cultivation, including the effects of temperature,

seasonality, and soil moisture. Their best estimate was 16

to 29 Gg yr-1 of CH3I. This range is similar to the previous

estimate of Muramatsu and Yoshida (1995), but is considerably lower than the earlier estimates of Redeker et al.

(2000) and Redeker and Cicerone (2004). Redeker and

Cicerone (2004) did note that CH3I emissions from paddies were positively correlated with temperature, essentially doubling with an increase of 10 degrees (25-35°C).

The first measurements of emissions of C2H5Cl,

CHBr3, and bromodichloromethane (CHBrCl2) from rice

paddies, or any other terrestrial ecosystem, were reported

by Redeker et al. (2003). Integrated seasonal emissions

of C2H5Cl were 0.60 ± 0.50 milligrams (1 mg = 10-3 g)

per meter squared (mg m-2), about 25 micrograms (1 μg =

10-6 g) m-2 for CHBr3, and 2.9 μg m-2 for CHBrCl2.

Redeker et al. (2003) extrapolated their findings to the

global coverage of rice and calculated an annual emission

of about 1 Gg of C2H5Cl, which is insignificant compared

with annual global emissions of c. 300 Gg yr-1.

Evidence for direct terrestrial input of inorganic and

organic halogens into the lower troposphere comes from

studies of volcanoes and salt flats. Bureau et al. (2000)

estimated the bromine yield of recent volcanic eruptions

based on laboratory examination of ejected gases, and

Bobrowski et al. (2003) measured enhanced BrO in a volcanic plume. Both authors suggest that these emissions

could contribute to stratospheric bromine. Bobrowski et

al. (2003) further estimated that volcanoes might emit c.

30 Gg BrO yr-1 into the lower atmosphere. Afe et al.

(2004), however, showed that space-based measurements

of column BrO are not correlated with enhancements of

column sulfur dioxide in volcanic plumes, casting some

doubt on the latter hypothesis. Volcanic perturbation to

stratospheric bromine is not commonly considered in

ozone assessment simulations.

Stutz et al. (2002) recorded elevated levels of BrO

(6 ± 0.4 ppt) and ClO (15 ± 2 ppt) associated with mobilization of halogens from salt near the Great Salt Lake in

the U.S. However, no flux estimates were provided, and

it is not clear how widespread this phenomenon may be.

Similarly, Zingler and Platt (2005) reported IO abundances

of 0.3-2 ppt, and occasionally more than 10 ppt, over Dead

Sea salt flats, apparently from inorganic sources. It is

unlikely, however, that these emissions could contribute

significantly to stratospheric halogen loading.

VERY SHORT-LIVED SUBSTANCES

2.13
VERY SHORT-LIVED SUBSTANCES

2.14

Table 2-3. Estimated local lifetimes, burdens, removal rates, and sources for some halogenated VSLS.

New information since WMO (2003) is indicated in boldface.

Compound Local Estimated Burden Estimated Estimated Source from

Lifetime (Gg) Removal Rate Inventory or Estimate of

(days) (Gg yr-1) Biogeochemical Cycle (Gg yr-1)

CH2BrCl 150a 1.2 (Br), 0.5 (Cl)1 2.9 (Br), 1.3 (Cl)c

CH2Br2 120a 18-22 (Br)b 55-67 (Br)c

19 (Br)1 58 (Br)c

CHBrCl2 78a 1.3-1.5 (Br)b 6.1-7.0 (Br)c 17 (Cl) open ocean15,h

1.2-1.3 (Cl) 5.4-6.2 (Cl) 19 (Br) open ocean15,h

1.2 (Br), 1.1 (Cl)1 5.5 (Br), 4.9 (Cl)c

CHBr2Cl 69a 0.8-2.2 (Br), 0.2-0.5 4.2-12 (Br), 0.9- 5 (Cl) open ocean15,h

(Cl)b 2.7 (Cl)c 23 (Br) open ocean15,h

2.3 (Br), 0.5 (Cl)1 12 (Br), 2.7 (Cl)c

CHBr3 26a 11-18 (Br)b 150-250 (Br)c 209 (47-370) (Br)3

14 (Br)1 200 (Br)c 800 (240-1760) (Br)9

2852, d 28 (20-112) (Br) industry9

CH3I 5e 1.7-2.2 (I)b 120-160 (I)c 90-450 (I)4

6f 4.8 (I)f 214 (I)f 272 (I) total10

191 (I) net ocean (incl. 66 

anthropogenic)10

180 (I) open ocean15,h

C2H5I 4a 0.5 (I) 46 (I)c

CH2ClI 0.1a 95 (I) open ocean15,h

27 (Cl) open ocean15,h

CH2Cl2 140a 83-250 (Cl)b 220-650 (Cl)c 487 (Cl) industrial5

1806, e 250 (Cl)6 650 (Cl)c 160 (Cl) ocean

500 (Cl)6,d 49 (Cl) biomass burning16,i

CHCl3 150a 66-130 (Cl)b 160-320 (Cl)c 564 (Cl)6

1837,e 210 (Cl)7 511 (Cl)c 588 (392-784) (Cl)11

470 (350-600)7,e 320 (240-400) (Cl) seawater11

412 (Cl)6,d 196 (107-285) (Cl) soil11

C2HCl3 4.6a 3.1 (Cl)b 260 (Cl)c 95 (Cl) industry5

5.56,e 5.30 (Cl)6 440 (Cl)c 20 (Cl) ocean6

350 (Cl)6,e 3 (Cl) fossil fuel6

C2Cl4 99a 17-85 (Cl)b 63-310 (Cl)c 313 (Cl) industry5

1336,e 160 (Cl)6 590 (Cl)c 16 (Cl) ocean6

440 (Cl)6,e 2 (Cl) fossil fuel6

222 (210-235) (Cl)12,g

C2H5Cl ~ 3013 14 (Cl)13 165 (Cl)c

~ 2414 4.6-7.3 (Cl)14 70-110 (Cl)c

CH2ClCH2Cl 70a 5-26 (Cl)b 26-130 (Cl)c 7008

References: 1Kurylo and Rodríguez et al. (1999); 2Dvortsov et al. (1999); 3Carpenter and Liss (2000); 4Singh et al. (1983), Liss and Slater (1974),

Moore and Groszko (1999); 5McCulloch et al. (1999); 6Keene et al. (1999); 7Khalil and Rasmussen (1999); 8Khalil (1999); 9Quack and Wallace

(2003); 10Bell et al. (2002); 11McCulloch (2003); 12Simpson et al. (2003); 13Redeker et al. (2003); 14Low et al. (2003); 15Chuck et al. (2005); 16Lobert et al., (1999). a From Table 2-4 of Ko and Poulet et al. (2003). b Burden estimated using Equation 2.3 and the median BL mixing ratios from Table 2-7 in Ko and Poulet et al. (2003), and the estimated scale height

from Table 2-10 in Ko and Poulet et al. (2003). c From estimated lifetime and estimated burden, i.e., (column 3)/(column 2). The range reflects the range in the estimated burden. d From two-dimensional (2-D) model. 
2.2.3 Oceanic Emissions

Numerous studies evaluating the oceanic production, degradation, and emission of VSLS suggest high

spatial and temporal variability, with emissions generally high in the tropics and near ocean fronts, and highest

in coastal regions (Figure 2-4). These studies continue

to suggest that oceanic emissions constitute by far the

largest source of brominated and iodinated VSLS to the

atmosphere, making up 90-95% of the total global flux

of these gases. The potential for spatially and seasonally

dependent variability is shown clearly in the work of

Cohan et al. (2003), where model results indicate that the

mean oceanic flux of CH3I from the Southern Ocean into

the marine boundary layer at 40°S-50°S is 3.1 (1.9-5.2)

μg m-2 day-1 during the summer. For comparison, the

global annual oceanic flux estimate reported by Moore

and Groszko (1999) corresponds to 1.0 to 2.7 μg m-2 day-1,

and the open ocean range given in Butler et al. (2006) is

0.6 to 5.7 μg m-2 day-1.

2.2.3.1 BROMINE

Both Butler et al. (2006) and Quack et al. (2004)

show the highest open ocean fluxes of CHBr3 in areas of

upwelling, such as near the equator and at ocean fronts,

and their combined data suggest that tropical fluxes are

highest in the Pacific, where equatorial upwelling is most

intense. Quack and Wallace (2003) noted that subtropical coastal and shelf oceans of both hemispheres contribute significantly to global emissions, with 180 (72-

270) Gg Br (CHBr3) yr-1 emitted from their “shore”

regime and 380 (240-940) Gg Br (CHBr3) yr-1 emitted

from their “shelf” regime, representing 23 and 48% of

the entire oceanic CHBr3 flux, respectively. Extrapolating Atlantic open ocean fluxes to the global ocean

yielded 240 (32-800) Gg Br (CHBr3) yr-1, representing

about one-third of the global oceanic CHBr3 emission of

~800 (240-1800) Gg Br (CHBr3) yr-1. These estimates

are similar to the 160 and 890 Gg Br yr-1 reported by

Butler et al. (2006) for open ocean and global sources

respectively. Quack and Wallace (2003) attribute much

of the source to CHBr3 production by phytoplankton,

which they calculated to be ~240 Gg Br yr-1.

The large differences between estimated open

ocean fluxes and global fluxes underscore the importance of coastal waters as a substantial source of CHBr3.

High fluxes of many polyhalogenated VSLS are attributed mainly to attached macroalgae (e.g., Laturnus et

al., 1996; Carpenter et al., 2000), and numerous investigators have reported very high mixing ratios in and

above coastal waters (e.g., Carpenter and Liss, 2000).

There have been, however, no systematic investigations

of these gases in tropical coastal regions until recently,

when Yokouchi et al. (2005b) measured atmospheric

concentrations of CHBr3, CH2Br2, and CHBr2Cl over

several areas, including both Arctic and tropical coastal

waters, and the open ocean, finding by far the highest

amounts of these gases in association with tropical

coasts. Their global estimate of the oceanic CHBr3

source was 820 (± 310) Gg Br yr-1, which is consistent

with the independent estimates of Quack and Wallace

(2003) and Butler et al. (2006). 

Although coastal macroalgae are undoubtedly an

important global source, Quack et al. (2004) and Butler et

al. (2006) also provide evidence that the sources of CHBr3

and CH2Br2 throughout the tropical open ocean are associated with the deep chlorophyll maximum near the thermocline. Transport of these gases to the ocean surface in

association with upwelling and mixing is consistent with

the higher concentrations and fluxes observed at the

equator and near ocean fronts (Chuck et al., 2005; Butler

et al., 2006). This may be significant because equatorial

upwelling, which carries these gases to the surface where

they are emitted, is influenced by wind speed and direction, which in turn are affected by climate, weather, and

El Niño-Southern Oscillation (ENSO) events. This suggests a link between climate, wind-driven upwelling, and

the supply of bromine to the tropical upper troposphere,

where it can be transported to the lower stratosphere (see

Section 2.4). 

Several studies indicate that CHBr3 sources contribute more organic and reactive bromine to the lower

atmosphere than other organobromines. Carpenter et al.

VERY SHORT-LIVED SUBSTANCES

2.15

Table 2-3, continued.

e Burden estimated from information given in Ko and Poulet et al. (2003): Equation 2.3, Table 2-7 (median BL mixing ratios), Table 2-10 (estimated

scale heights). f From three-dimensional (3-D) model simulation of Bell et al. (2002). g Estimate for 2002. h Estimated using data from Chuck et al. (2005). The surface ocean during the Chuck et al. study was generally a sink for CHBr3. i May be a less significant and more uncertain source than estimated here, although no new global emission estimates are available (Simmonds et al.,

2006; Scheeren et al., 2002).
VERY SHORT-LIVED SUBSTANCES

2.16

Figure 2-4. Plots of the oceanic fluxes of CHBr3 and CH2Br2 into the marine boundary layer based upon data

from seven research cruises between 1994 and 2004 (Butler et al., 2006). Cruise data are color-coded to

match those in the map. All measurements are on the same calibration scale.
(2003) extrapolated their data to estimate molar source

strengths of CHBr2Cl and CH2Br2, suggesting that the flux

of CHBr2Cl is about 3-6% of that of CHBr3, and that

CH2Br2 emissions would be about 15-25% of the global

CHBr3 flux. Yokouchi et al. (2005b) suggest somewhat

similar molar emission ratios, with CHBr2Cl amounting

to ~7% and CH2Br2 to ~11% of the flux in coastal waters.

Butler et al. (2006) noted that CH2Br2 fluxes were about

25-35% of CHBr3 fluxes for data collected mostly in the

open ocean. These values all fall within ranges estimated

independently from concentration and lifetime data.

2.2.3.2 IODINE

Fluxes of iodinated VSLS are driven in part

through photochemistry in the surface waters. Results of

a study of the production of CH3I in the tropical Atlantic

Ocean support prior inferences that CH3I is produced in

the open ocean by a light-dependent production pathway

that is not directly dependent on biological activity

(Richter and Wallace, 2004). This assertion is supported

by data from Smythe-Wright et al. (2005) and Chuck et

al. (2005). Also, photo- or biodegradation of some polyhalogenated methanes can lead to the production of

others. For example, Martino et al. (2005) carried out

laboratory studies of aqueous degradation kinetics of

the photolabile compounds CH2I2, CH2BrI, and CH2ClI

in different types of natural and artificial seawater.

Photolysis of CH2I2 in artificial and natural seawater generated CH2ClI with a yield of 25-30%, suggesting that

this reaction is an important source of marine CH2ClI.

Dark-incubations indicated that photolysis is the main

abiotic degradation mechanism of CH2I2 in seawater and

that CH2I2 originating at depth and transported to the surface would be photolyzed before escaping into the atmosphere. The longer aquatic photolytic lifetime of CH2ClI

allows this compound to vent to the atmosphere and

indeed, although CH2ClI has been detected at very low

levels in the remote marine boundary layer, its sea-air

fluxes appear to be of the same order as those of CH3I

(Carpenter et al., 2003; Chuck et al., 2005).

2.2.3.3 CHLORINE

Although previous studies (e.g., Khalil et al., 1999)

led Keene et al. (1999) to conclude that oceanic emissions

of dichloromethane (CH2Cl2) could account for about 25%

of the total emissions to the atmosphere, recent work by

Moore (2004) suggests that the apparent supersaturation

of this gas in higher latitude waters is an artifact of circulation and mixing. Moore (2004) interpreted the measurements of CH2Cl2 in waters of the North Atlantic and

Labrador Sea to show that the compound at depth in the

ocean is imprinted with an older atmospheric source, and

appears to persist for years to decades in deeper waters.

In this sense, the observed surface supersaturation of the

gas would likely have resulted from the recent decline of

this gas in the atmosphere (e.g., Simmonds et al., 2006;

Thompson et al., 2004) with the ocean gradually re-equilibrating, and not necessarily the result of production in the

ocean. Thus, the ocean “source” of 200 Gg Cl yr-1 would

likewise be a matter of re-equilibration, not production. 

2.2.4 Industrial and Other Anthropogenic

Emissions

2.2.4.1 BROMINE

Except for the chlorinated VSLS, anthropogenic

emissions of most halogenated VSLS are presently

dwarfed by natural emissions (Table 2-3). Emissions of

CHBr3 through water treatment, however, can be locally

significant. Quack and Wallace (2003) reassessed the

anthropogenic CHBr3 source from water chlorination,

including reactions on saltwater effluents, and suggested

that it would amount to about 28 Gg Br yr-1 (0.20-110 Gg

Br yr-1), which is minor compared with natural sources

(<5% of biogenic CHBr3 emissions). Zhou et al. (2005)

reported contributions from effluents from the chlorination of drinking, waste, and recreational waters of ~1-10%

of the CHBr3 emitted from the Great Bay area of New

England, U.S. In the same region, they determined that

the contribution from coastal power plant cooling effluents amounted to 15-50% of the CHBr3 measured in the

Great Bay, constituting most, if not all, of the anthropogenic source to the bay.

Because polar firn air provides a natural archive

of “old” (typically from early 20th century) air, long-term

trends of various gaseous species can, with some caveats,

be reasonably reconstructed. Significant 20th century

trends are often indicative of increasing anthropogenic

emissions. For example, Worton et al. (2006) reported

firn air measurements of the brominated trihalomethanes

(THMs: CHBr3, CHCl2Br, CHClBr2) and dihalomethanes

(DHMs: CH2Br2 and CH2ClBr) in Greenland. The brominated DHMs showed no annual trend over time, suggesting that their sources are almost entirely natural. The

THMs, however, appeared to increase by 20% from

1950-1990. This translates to increases of ~16 ± 9, 0.7 ±

0.4, and 0.8 ± 0.3 Gg Br yr-1 for CHBr3, CHBr2Cl, and

CHBrCl2, respectively. Worton et al. (2006) noted that

similarities in the trends in the deepest sections of the

firn (oldest air) indicate that the three THMs may have

similar sources. The authors further suggested that

VERY SHORT-LIVED SUBSTANCES

2.17
chlorination of seawater for cooling coastal power plants

may be the largest anthropogenic source of the brominated THMs. Emissions of THMs from water chlorination could potentially increase in the future in response

to increasing demand for water treatment (potable and

waste water) and power generation.

2.2.4.2 CHLORINE

Antarctic firn air measurements from Trudinger et

al. (2004) show CHCl3 and CH2Cl2 increasing in atmospheric mixing ratios from 3.9 ppt to 6.35 ppt and 1.4 ppt

to 8.9 ppt, respectively, between 1940 and 1990. They

noted that records of anthropogenic emissions of CH2Cl2

were consistent with their results. Northern Hemisphere

firn air measurements from Worton et al. (2006) also support an anthropogenic contribution to CHCl3. Their

model calculations suggest anthropogenic emissions

were ~14-20% of total CHCl3 emissions (270-335 Gg Cl

yr-1) in 1950, and increased to 41-50% of the total (417-

506 Gg Cl yr-1) at the peak in global mixing ratio in about

1990. Anthropogenic emissions subsequently decreased

to ~19% of total emissions (310 Gg Cl yr-1) in 2001. This

anthropogenic component was attributed mostly to paper

and pulp manufacture, and the declining emissions since

~1990 to changes in working practices in this industry.

Recent ambient atmospheric data also suggest that

emissions of these two gases are declining. Prinn et al.

(2000) report data for CHCl3 from 1983-1998 with a trend

ranging from −0.1 to −0.4 ppt yr-1. Simmonds et al. (2006)

noted that measurements from Mace Head, Ireland, and

from the National Oceanic and Atmospheric Administration (NOAA) U.S. global monitoring sites (Thompson

et al., 2004) show a decrease in CH2Cl2 mixing ratios

from 1995-2004 with trends of −0.7 ppt yr-1 (1995-2004)

and −0.3 ppt yr-1 (1998-2004), while an increase of 0.05

ppt yr-1 was observed at Cape Grim, Australia.

Emissions of some chlorinated gases such as tetrachloroethene (C2Cl4) have been targeted for reduction

owing to public health concerns. Levels of tropospheric

C2Cl4, for which sources were estimated in WMO (2003)

to be 95% anthropogenic, declined substantially between

1989 and 2002 (Simpson et al., 2004). During this period,

annual mean C2Cl4 mixing ratios for the extratropical

Northern Hemisphere dropped from 13.9 ± 0.5 ppt to less

than half this value (6.5 ± 0.2 ppt), and global averages

declined from 6.3 ± 0.6 ppt to 3.5 ± 0.2 ppt (Simpson et

al., 2004). These values suggest that the global C2Cl4

burden decreased by roughly 820 Gg Cl between 1989-

2002, and is consistent with decreasing anthropogenic

Northern Hemispheric emissions. Simmonds et al. (2006)

showed continued declines in C2Cl4 from 2000-2004.

They report means (trends in parentheses) for Mace Head,

Ireland, and Cape Grim, Australia, of 4.9 ppt (−0.18 ppt

yr-1) and 0.75 ppt (−0.01 ppt yr-1), respectively.

High regional variability in emissions of anthropogenic VSLS is underscored by the study of Barnes et

al. (2003) who, in contrast to the above, found no uniform trend of C2Cl4 between 1996 and 1999 at Harvard

Forest, Massachusetts, U.S., but suggested that urban

and industrial emissions of C2Cl4 for the New York CityWashington, D.C. corridor were increasing. Similarly,

estimates from an aircraft study over Sagami Bay, Japan,

suggest high emission rates of CH2Cl2 (61.8 Gg Cl yr-1)

and C2Cl4 (28.0 Gg Cl yr-1) (Yokouchi et al., 2005a). These

estimates, as well as those for CHCl3 (6.2 Gg Cl yr-1) and

trichloroethene C2HCl3 (39.8 Gg Cl yr-1), are consistent

with the Japanese Pollutant Release and Transfer Register

for 2002 (PRTR, 2004). 

2.2.4.3 NEW VSL GASES

Future emissions of VSLS in the upper troposphere

are potentially problematic. Trifluoroiodomethane (CF3I)

is being considered as a direct replacement for halon use

in aircraft, and as a replacement for the potent greenhouse

gas, hexafluoroethane (C2F6), currently used in the plasma

etching industry. No information on its emission is currently available, although its use on aircraft could potentially lead to emission at altitude. This species is rapidly

photolyzed (Table 2-1), but emission at altitude could

provide a pathway for future injection of iodine to the

stratosphere (Li et al., 2006). A new product, “PhostrEx”

(PBr3), has recently won approval from the U.S.

Environmental Protection Agency for in-flight aircraft

engine fire protection, and has passed all Federal Aviation

Administration certification fire testing; future widespread use is possible. Although this compound is

highly soluble and consequently has a very short lifetime

(<0.1 s at 50% humidity), its release during flight in the

upper troposphere could be of concern because of the

low humidity and hence slower removal rates of the gas

and its degradation products at these altitudes.

2.2.5 Sources and Distributions of

Inorganic Halogens of Marine Origin

Any substance present in the upper troposphere can

potentially be exported to the stratosphere; therefore inorganic halogens in the upper troposphere have to be considered to quantify the total import of halogens into the

stratosphere. In the tropics, rapid convective transport

means that VSLS may also be transported from the

VERY SHORT-LIVED SUBSTANCES

2.18
boundary layer into the tropical upper troposphere (see

Section 2.4.1).

BrO has been observed directly in the midlatitude

marine boundary layer (Leser et al., 2003; Saiz-Lopez et

al., 2004a). The presence of inorganic bromine generally

in the marine boundary layer has also been inferred from

the routinely observed large bromide deficit in sea salt

aerosol (e.g., Sander et al., 2003). These occurrences are

almost certainly due to halogen activation from sea salt by

heterogeneous reaction chemistry.

Combinations of space-, ground-, and balloonbased measurements indicate that BrO is widely present

in the free troposphere, with mixing ratios of 0.2-2 ppt

(Harder et al., 1998; Wagner and Platt, 1998; Pundt et al.,

2000; McElroy et al., 1999; Fitzenberger et al., 2000;

Wagner et al., 2001; Van Roozendael et al., 2002; Richter

et al., 2002; Schofield et al., 2004). The study by Yang et

al. (2005), incorporating a detailed bromine chemistry

scheme of gas-phase and heterogeneous reactions, showed

that monthly mean mixing ratios of 0.1-1.0 ppt of BrO in

the free troposphere can be explained by a combination of

bromine release from sea salt aerosol and the breakdown

of bromomethanes. Other bromine sources were not considered in this study, but might contribute (e.g., volcanic

BrO, as noted above, and tropospheric decomposition of

other organic bromine gases). They suggest that sea salt

might contribute as much as 10% of the total Br in the

upper troposphere.

In contrast to bromine, there is limited evidence for

the presence of inorganic chlorine radicals in the free troposphere.

McFiggans et al. (2004) and Saiz-Lopez et al.

(2006) reported direct coastal boundary layer observations

of molecular iodine, ultrafine particle production, and

iodocarbons. They demonstrated for the first time that

ultrafine iodine-containing particles are produced by intertidal macroalgae exposed to ambient levels of ozone.

However, the lifetime of inorganic iodine species released

in the boundary layer is short, and they are generally

unlikely to be transported to the stratosphere.

2.3 ATMOSPHERIC CHEMISTRY OF VSLS

The halogenated source compounds considered

here are listed in Table 2-1. The Table gives the estimated

local lifetimes (τlocal) for these compounds, defined as

(τlocal)

-1 = (τOH)-1 + (τJ)-1 where (τOH) is the lifetime due to

reactions with hydroxyl radical (OH), and (τJ) is the lifetime due to ultraviolet (UV) photolysis. It is worth noting

that even within these halogenated VSLS there is a wide

range of local lifetimes, with values ranging from a few

minutes to 150 days. The atmospheric loss processes for

the source gases (organic halogens) occurs primarily in

the gas phase via reaction with OH and UV photolysis.

Multiphase processing of source gases is not a significant

loss process and is not considered in this chapter.

The removal of inorganic halogenated reservoir

species (e.g., hydrogen chloride, HCl) by wet and dry deposition is the main atmospheric loss process for inorganic

halogens. Heterogeneous processes that recycle inorganic

bromine and (most probably) iodine to insoluble reactive

forms, may significantly increase the effective lifetime of

Brx and Ix (see boxed text in Section 2.1 for definition of

Brx; Ix is analogously defined) in the troposphere and the

lower stratosphere and, as a result, enhance the efficiency

of halogen transport via the PG pathway into the lower

stratosphere, as discussed in Section 2.5. 

Processing of both organic and inorganic halogenated degradation products on atmospheric aerosol

and cloud particles is currently not well represented in

atmospheric models, although the knowledge of these

processes continues to improve. For example, a clearer

picture is evolving regarding the partitioning of trace

gases onto ice surfaces in cirrus clouds (Abbatt, 2003),

and the mechanisms of halogen reactions on ice surfaces

under tropospheric conditions (Fernandez et al., 2005).

In this section, we assess the data needed to represent the

atmospheric processing in both the gas and condensed

phases of halogenated VSLS in atmospheric models.

2.3.1 Removal of Halogen Source Gases

The halogenated VSL SGs listed in Table 2-1,

with the exception of CF3I, C2H5Br, 1,2-dibromoethane

(CH2BrCH2Br), and PBr3 were evaluated in WMO (2003).

The kinetic and photochemical parameters for the previously assessed VSL source gases have been updated in the

recent evaluations of the International Union of Pure and

Applied Chemistry (IUPAC, Atkinson et al., 2005; http://-

www.iupac-kinetic.ch.cam.ac.uk), and the National Aeronautics and Space Administration (NASA) Jet Propulsion

Laboratory (JPL), U.S. (Sander et al., 2006; http://-

jpldataeval.jpl.nasa.gov), with only minor revisions.

The atmospheric loss of the VSL SGs can be summarized as follows: chlorinated VSLcompounds have long

photolysis lifetimes and are predominantly removed in the

troposphere through reaction with OH; the brominated

VSLS are removed by a combination of UV photolysis

and OH reaction depending on the degree of halogen substitution (for example, increased importance of photolysis

with higher bromine substitution); and iodine-containing

VSLS are removed almost exclusively by photolysis.

Bayes et al. (2003) reported bromine atom quantum yields

for the photolysis of CHBr3 at several wavelengths beVERY SHORT-LIVED SUBSTANCES

2.19
tween 303 and 324 nanometers (nm). Their results imply

that the UV absorption cross section data currently recommended for use in the determination of the atmospheric

photolysis rate of CHBr3 may be systematically high by

as much as 20% over this wavelength region.

The OH rate coefficients and atmospheric photolysis rates for C2H5Br and CH2BrCH2Br are given in Table

2-4. The photolysis lifetimes are significantly longer than

their respective OH reaction lifetimes, and therefore photolysis is not an important atmospheric loss process for

these compounds.

2.3.2 Tropospheric Lifetimes of

Halocarbons

The lifetimes given in Tables 2-1 and 2-3 provide

only an approximate estimate of the global mean lifetimes,

because of significant regional variations in the OH radical concentration, solar flux, and the spatial and seasonal

distributions of the halogen source gases. The geographic

distribution of VSL SG surface fluxes with respect to these

regional variations in the corresponding sinks is an important factor in determining the global mean atmospheric

lifetimes of the VSL SGs. Warwick et al. (2006) have provided an estimate of the importance of the geographic flux

distribution on the atmospheric lifetime of CHBr3 by using

a global atmospheric three-dimensional (3-D) model to

calculate the CHBr3 global atmospheric lifetime for a set

of prescribed oceanic CHBr3 emission data. In their calculations, the CHBr3 atmospheric lifetime, defined as the

global burden divided by the annual global loss, varied

from 37 days in a scenario containing emissions distributed over the entire open ocean, to 15 days in a scenario in

which CHBr3 was emitted from tropical coastline regions

only (see Table 2-5). The shorter lifetime in the tropical

emission scenario is due to higher OH concentrations and

UV levels in this region. This represents an uncertainty of

approximately ±40% in the CHBr3 atmospheric lifetime

that is solely due to changes in the emission distribution.

For halogen source gases with longer atmospheric lifetimes and better constrained flux distributions, this uncertainty will be reduced.

The different CHBr3 emission scenarios in the

Warwick et al. (2006) study also showed a significant

variation in the amount of CHBr3 reaching the tropical

upper troposphere. A scenario based on a uniform ocean

emission distribution of 210 Gg CHBr3 yr-1 (Kurylo and

Rodríguez et al., 1999; Carpenter and Liss, 2000)

showed CHBr3 peak mixing ratios of 0.2 ppt at 110 hPa,

whereas a scenario with larger emissions (587 Gg CHBr3

yr-1) predominantly situated in the oceanic tropics,

which showed better agreement to tropospheric observations, contained peak CHBr3 mixing ratios of ~1 ppt

at the same altitude.

VERY SHORT-LIVED SUBSTANCES

2.20

Table 2-4. OH rate coefficient and photochemical parameters for bromoethanes added since the WMO

(2003) Assessment. Temperatures (T) in Kelvin.

Compound k(OH) k300(OH) Photolysis Rate References

(cm3 molec–1 s

–1) (cm3 molec–1 s–1) Constant, J a (s–1)

Bromoethane 2.9 × 10–12exp(−640/T) 3.43 × 10–13 <1 × 10–9 Sander et al. (2006)

(C2H5Br)

1,2-dibromoethane 1.44 × 10–11exp(−1270/T) 2.1 × 10–13 <1 × 10–9 Qiu et al. (1992);

(CH2BrCH2Br) Sander et al. (2006)

a 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).

Table 2-5. Bromoform emission scenarios used

by Warwick et al. (2006).

Global Flux Geographic Modeled

(Gg CHBr3 Distribution of Atmospheric

yr-1) Emissions Lifetime (days)

210 uniform ocean a 30

235 open ocean b 37

587 global coastlines b 26

400 tropical ocean c 21

587 tropical coastlines c 15

595 tropical ocean and 18

tropical coastlines c

a Kurylo and Rodríguez et al. (1999). b Quack and Wallace (2003). c Warwick et al. (2006).
2.3.3 Production and Gas-Phase Removal

of VSL Organic Product Gases

The first generation of stable reaction products

formed following the gas-phase degradation of the halogenated VSL SGs was reviewed in WMO (2003). In general, the stable products consist of a variety of organic

carbonyl (RC(O)R′) and peroxide (ROOR′) compounds,

and inorganic halogen species. Since the last Assessment,

little new evidence for the atmospheric reactivity and

photolysis of these compounds has been reported. The

recent IUPAC and JPL evaluations include the updates

for these compounds where available. Detailed degradation mechanisms for halomethanes in general, and

n-propyl bromide (n-C3H7Br) specifically, were given in

WMO (2003). A major uncertainty in the evaluation of

the impact of the degradation products is a lack of experimental data for their reactivity and photolysis.

The atmospheric degradation mechanisms for the

VSL SGs are reasonably well understood, although quantification of the various possible reaction channels and

degradation products is lacking in some cases. In general,

the atmospheric degradation of the halogenated VSL

source gases leads to the formation of compounds with

atmospheric lifetimes shorter than the lifetime of the

source gas. A notable exception is the formation of phosgene, COCl2, which is a product formed in the degradation of the VSLS chloroform (CHCl3) and tetrachloroethene (C2Cl4). Phosgene is also formed in the

degradation of longer lived chlorinated compounds such

as carbon tetrachloride (CCl4) and methyl chloroform

(CH3CCl3). In the stratosphere, COCl2 is removed mainly

by photolysis, since the rate constant for reaction with OH

is relatively small (<5 × 10-15 centimeters cubed per molecule per second). On the basis of recently published cross

sections (Atkinson et al., 2005; Sander et al., 2006), the

local photolysis lifetime of COCl2 is estimated to be similar to that of CFC-11 (trichlorofluoromethane; CCl3F),

and about twice that of CCl4, at altitudes above 18 km at

30°N; thus significant stratospheric loss of COCl2 is to be

expected. Kindler et al. (1995), using a one-dimensional

(1-D) model, reported a stratospheric lifetime for COCl2

of ~2 years (for COCl2 produced in the stratosphere); less

than lifetimes calculated for long-lived SGs such as CCl4.

Kindler et al. calculated a much longer stratospheric lifetime (>8 years) for COCl2 produced in the troposphere,

which they explain as being due to the requirement for

COCl2 produced in the troposphere to be transported to

higher stratospheric altitudes, where photolysis rates are

greater. In the troposphere, removal of COCl2 is dominated by aerosol washout (total wet removal lifetime of

~60 days from Kindler et al. (1995) and ~70 days from

WMO, 2003). Therefore, wet removal of COCl2 may further reduce the contribution of chlorinated source gases

(both long-lived and VSLS) to stratospheric chlorine (see

Section 2.3.4.1).

When loss of CHBr3 is initiated by reaction with

OH, the final organic decomposition product is likely to

be carbonyl dibromide (COBr2). When loss is initiated by

photolysis, then formyl bromide (COHBr) is thought to be

produced (Ko and Poulet et al., 2003). Present kinetics

(Sander et al., 2006) indicate about two-thirds of the loss

of CHBr3 occurs by photolysis. There are few kinetic

studies of COBr2 and COHBr and no atmospheric observations of these species. Measurements of the UV cross

sections (Libuda, 1992) indicate, however, that these compounds have a shorter lifetime due to photolysis loss than

does CHBr3. Many other organic intermediates are involved with the decomposition of other VSL SGs (e.g.,

Ko and Poulet et al., 2003).

2.3.4 Multiphase Processes Involving

Inorganic Bromine and Iodine

Compounds

Models used to predict tropospheric amounts of

halocarbons and their degradation products currently use

very simplified algorithms for heterogeneous processes

and physical removal. Although the multiphase chemistry for inorganic bromine, and especially iodine compounds, is still poorly known, there are significant new

data and understanding since the last Assessment, which

are summarized below. This opens the possibility for a

more realistic treatment of multiphase chemistry and

more quantitative representation of the heterogeneous

recycling of reactive halogens, the potential importance

of which was highlighted in WMO (2003). In future

model studies, the partitioning of gases on ice should be

described using the suggested Langmuir description.

Uptake on aqueous particles should, especially for very

soluble species (see Yin et al., 2001), be calculated kinetically, i.e., not assuming instant equilibrium between gas

and aqueous phase. The relevant Henry’s law coefficients are provided in this section.

2.3.4.1 WET DEPOSITION PROCESSES

The parameters needed to evaluate the rates of the

processes leading to wet deposition are the partition coefficients for gas-phase degradation products into liquid

water (the effective Henry’s constants, Heff), and on to ice

surfaces (the Langmuir equilibrium constants, Keq).

Henry’s constants were given for HOX and XONO2 (X =

Br and I) in the last Assessment. Subsequently, an extenVERY SHORT-LIVED SUBSTANCES

2.21
sive database that covers these and other compounds

including HCl, HBr, HI, ClONO2, HOCl, BrCl, Br2, I2,

ICl, and IBr is available from the National Institute of

Standards and Technology (NIST) webbook (http://-

webbook.nist.gov/chemistry/). Heff values for these

species and for the organic chlorine reservoir COCl2 are

given in Table 2-6. The tropospheric removal of this and

similar halogenated carbonyls was considered in depth in

WMO (1995), and the conclusion that these compounds

are relatively rapidly removed by precipitation remains

valid.

A substantial amount of new data for trace gas-ice

interaction at temperatures of the upper troposphere/lower

stratosphere (UTLS; 200-230 K) has been reported from

laboratory and field studies (Abbatt, 2003; Popp et al.,

2004). Partitioning of HX (X = Cl or nitrate, NO3) to the

ice phase occurs by Langmuir-type adsorption, in which

the amount of material adsorbed is normally limited to

sub-monolayer surface coverage, depending on the partial

pressure of the gas and the Langmuir equilibrium constant,

Keq (= kadsorption/kdesorption). The surface coverage, u (molec

cm-2), is given in terms of Ns (the maximum coverage in

VERY SHORT-LIVED SUBSTANCES

2.22

πTable 2-6. Henry’s Law constants (H, in units of moles per liter per atmosphere) and uptake coefficients (γ

het) for heterogeneous processing of inorganic halogens.

H298 γhet

(M atm-1)

H2O(l) H2O ice H2O ice Sulfuric acid, H2SO4 (aq)

HCl/HBr H2SO4 (aq) HCl/HBr doped

doped (X = Cl, Br)

HOCl 260 – γ0 > 0.2 0.3 time 0.02-0.2 a

saturates a dependent dependent on

HHOX and HHX

HOBr 93-1900 >0.01 0.3 0.004-0.3 >0.01 b 0.1-1

HOI >45 – – 0.05 0.015-0.07 b >0.02

ClONO2 hydrolysis 0.024 c – 0.3-0.6 dependent on dependent on

[H2SO4], HHX a

H, and T

BrONO2 hydrolysis 0.06 c >0.3 >0.2 b >0.2 c 0.9

(H2SO4(s))

IONO2 hydrolysis >0.1 c – >0.05 b >0.3 c –

Cl2 0.093 – >0.2 – – –

(+HBr(g))

Br2 0.76 – –– – –

I2 3.3 5 × 10-4 – 0.02 – –

(HI doped)

BrCl 0.94 <0.03 – – – –

ICl 110 – 0.02 0.3 – –

IBr 24 – 0.03 0.3 – –

HCl 1.1*

(Ka = 107 d) physical uptake only

HBr 0.71* (Ka = 109 d) physical uptake only

HI 2.5* e (Ka = 3.2 × 109 d) physical uptake only

COCl2 0.07 f

* Heff = H298 × exp(−¢Hs/R × (1/T − 1/298)) × (1 + Ka/[H+

]) (Ka = acidity constant) a http://www.iupac-kinetic.ch.cam.ac.uk/ b Protonation followed by reaction to form dihalogens. c Reaction leading to hydrolysis to form HOX, where X = Br, I. d The dissociation constants are from Arnaud (1966), or see Schweitzer et al. (2000). e The physical constant is from Marsh and McElroy (1985). f Slow hydrolysis occurs; see WMO (1995).
molec cm-2 at saturation), Keq, and the gas partial pressure

of the gas (PHX):

For uptake on ice, the value of Ns is approximately 3 ×

1014 molec cm-2 (Abbatt, 2003). At UTLS temperatures,

nitric acid (HNO3) and HX (X = Cl, Br, I) are strongly

adsorbed and the process is important. Weak acids and

organic species are generally weakly adsorbed and do not

partition efficiently to ice. When formation of long-lived

compounds (e.g., nitric acid trihydrate: NAT) is thermodynamically favored, there is a greater capacity for partitioning, since adsorption is not confined to the surface

layer. 

2.3.4.2 HETEROGENEOUS HALOGEN ACTIVATION

Conversion of stable inorganic halogen reservoirs,

which are removed by deposition, into reactive halogen

radicals potentially lengthens the lifetime of inorganic

halogen because the reactive products remain in the

gaseous phase. The key heterogeneous reactions (where

X and Y are atoms of Br, Cl, or I) leading to release of

reactive halogens are:

HOX + Y– + H+ → XY + H2O

XONO2 + Y– + H+ → XY + HNO3

XONO2 + H2O (aq) → HOX + HNO3 (aq)

These reactions can occur in aqueous aerosols

(water, sulfuric acid, sea salt) or on ice or solid aerosol,

and XY is partitioned into the gas phase where it is photolyzed. 

Heterogeneous reaction kinetics are expressed in

terms of dimensionless reactive uptake coefficients, γ

het,

which allow the rate to be evaluated on a per-collision

basis for a prescribed surface area of the condensed phase.

Table 2-6 provides a summary, extended since the last

Assessment, of representative γ

het values for surfaces relevant for the troposphere and the stratosphere. These are

based on the updated IUPAC and JPL evaluations, and

other recent literature sources.

The large uptake coefficients of halogen nitrates

(XONO2, where X = Cl or Br) into aqueous droplets

(Deiber et al., 2004) means that uptake in liquid water

clouds will be efficient, and since concentrations of bromide and iodide in precipitating clouds are likely to be

very small, wet deposition will predominate over activation. Evaporating cloud droplets, however, will lead to

reactive halogen recycling. Model calculations have suggested that uptake into cloud droplets can extend reactive

halogen lifetime and lead to a vertical redistribution of

inorganic bromine compounds (von Glasow et al., 2002).

In their global model runs, von Glasow et al. (2004)

showed that heterogeneous reactions can extend the lifetime of Bry from about 6-9 days to about 9-15 days below

500 hPa, with smaller changes in the free troposphere. A

laboratory study showed the release of labile dibromine

monoxide (Br2O) and molecular bromine (Br2) from sulfuric acid solutions following exposure to the soluble

species hypobromous acid (HOBr) and hydrogen bromide

(HBr) (Iraci et al., 2005). These observations suggest that

the PGI of Bry species should be evaluated for various

aerosol and cloud washout lifetimes.

Activation of halogens can also occur on ice surfaces, such as those in cirrus clouds. Laboratory studies

have shown that Cl and Br activation by reaction of chlorine nitrate (ClONO2) and HOBr with hydrogen chloride

(HCl) on ice surfaces at UTLS temperatures scale linearly

with the surface coverage of HCl (Mössinger et al., 2002;

Fernandez et al., 2005). At typical tropospheric HCl concentrations, the rate of activation of bromine on ice clouds

is probably fast, but chlorine activation is less efficient.

Nevertheless recent field observations (Thornton et al.,

2003) suggest that Cl activation is occurring on cirrus

clouds, leading to the local production of ClO. 

2.3.5 Iodine Chemistry

All iodine source gases with significant emissions

fall into the category of halogenated VSLS. However,

ozone destruction in the lower stratosphere due to catalytic

cycles involving iodine remains poorly understood. New

information on the chemistry of iodine species relevant

for the atmosphere has been obtained since the WMO

(2003) Assessment. The new chemical kinetics information may have significance for the efficiency of transport

of iodine species to the UTLS and the ozone destruction

efficiency factor for stratospheric Iy. Relevant multiphase

iodine chemistry is presented in Section 2.3.4 above.

The IUPAC and JPL kinetics evaluations have critically reviewed all but the most recent kinetic and photochemical studies. Rate coefficients for the reactions of

atomic iodine (I), IO, and nitric oxide (NO) with OIO

(Joseph et al., 2005; Plane et al., 2006), and IO with methyl

peroxy radicals (CH3O2) (Bale et al., 2005) have since

been reported (in units of cm3 molec-1 s

-1) to be 1.1 × 10-10,

1.2 × 10-10, 7.6 × 10-13 exp(607/T), and 6.0 × 10-11

, respectively (rates are for room temperature only, except where

noted). New photolysis quantum yield studies for OIO

have also been reported (Joseph et al., 2005; Tucceri et al.,

VERY SHORT-LIVED SUBSTANCES

2.23

θ = +N

K P

K P s

eq HX

1 eq HX
2006). The OIO quantum yield (Φ)measurements indicate Φ(O(3

P)) < 0.007 at 532 nm and between 558 and 578

nm, and Φ(I) < 0.24 at 532 nm and <0.05 between 560 and

580 nm. Photolysis quantum yields at longer wavelengths

have not yet been reported.

Reactions between halogen monoxide radicals have

several product channels. The efficiency of ozone (O3)

loss depends on the halogen species and the product yields.

For example, the IO self-reaction produces OIO (~40%)

and iodine oxide dimer IOIO (~55%) at 1000 hPa (Bloss

et al., 2001). The IOIO product is unstable (Misra and

Marshall, 1998) and most likely rapidly thermally decomposes to yield OIO. Therefore, the effective OIO yield

from the IO self-reaction is high. The efficiency of O3 loss

is reduced by photolysis of the halogen dioxides formed

in the IO + IO and IO + BrO reactions, if an oxygen atom

is produced.

OIO has a series of strong diffuse absorption bands

in the region between 470 and 610 nm, but only the I atomproducing channel is energetically possible: 

OIO + hν → I + O2

The alternate channel producing IO + O leading to no net

O3 loss requires wavelengths <470 nm. Recent laboratory

studies (Joseph et al., 2005) show that OIO absorption at

562 nm leads to internal conversion from the excited state

(OIO*) to the vibrationally excited ground state (OIO#),

rather than dissociation:

OIO + hν → OIO* → OIO#

OIO# + M → OIO + M

Nevertheless, a small probability of photolysis, integrated

over the entire OIO absorption band, could still lead to

significant photochemical conversion of OIO to I and consequent ozone loss.

Studies have shown that higher iodine oxides can

nucleate to form new particles (Hoffmann et al., 2001;

O’Dowd and Hoffmann, 2005). The formation of OIO in

the gas phase appears to initiate this process. In the atmosphere, the oxidation of OIO by reaction with IO or OIO

may lead to the formation of higher oxides (e.g., I2O3 and

I2O4) that are of low volatility and may condense on preexisting particles, providing a sink for reactive iodine.

This is different from the oxidation mechanisms of chlorine and bromine oxides that do not lead to the formation

of unreactive low-volatility oxides. This mechanism provides a possible explanation for the absence of detectable

IO concentrations in the UTLS from degradation of

iodocarbons, while the BrO radical derived from the oxidation of bromine source gases is clearly present in this

region (see Sections 2.2.5 and 2.5).

2.4 DYNAMICS AND TRANSPORT IN THE

TROPOPAUSE REGION AND IMPLICATIONS FOR VSLS

The distribution of VSLS is governed by the distribution of surface sources, and by competition between

vertical transport (boundary layer to free troposphere to

stratosphere) and chemical destruction or removal via

washout. The implication for VSLS-related stratospheric

ozone loss therefore depends crucially on the strength and

spatial distribution of transport processes, and of the chemical, microphysical, or moist processes that remove the

degradation product gases (PGs). Note in particular (see

following sections) that rapid vertical transport, which

potentially takes VSL SGs and PGs from the boundary

layer to the tropopause region, is usually associated with

moist processes (warm conveyor belts, convective clouds)

that can potentially remove soluble PGs through washout.

Nevertheless, such processes are potentially effective in

the vertical transport of total bromine (or other halogens),

since the time scale for such transport in an individual convection event (hours) is short compared with the time scale

for chemical breakdown of the VSLS.

Given that VSL SGs and PGs reach the upper troposphere through potentially rapid transport mechanisms, it is important to establish how rapidly they are

subsequently transported into the stratosphere via the SGI

and PGI pathways. A schematic depiction of transport

processes in the upper troposphere and lower stratosphere

relevant to VSLS is shown in Figure 2-1. There is net

transport across the tropopause from troposphere to stratosphere in the tropics and net transport across the

tropopause from stratosphere to troposphere in the extratropics, with this net transport controlled by global-scale

processes as part of the Brewer-Dobson circulation

(Holton et al., 1995). However, when considering VSLS,

it is particularly important to note first, that transport

across the tropopause is in fact two-way, especially at midlatitudes, with both troposphere-to-stratosphere transport

(TST) and stratosphere-to-troposphere transport (STT),

and second, that the transition from tropospheric chemical

characteristics to stratospheric chemical characteristics

takes place not in a sharp jump across a uniquely defined

tropopause, but across a finite tropopause transition layer,

most notably in the tropics. In the tropics, it is now conventional to describe the transition layer as the “tropical

tropopause layer” (TTL). It is now clear that in the extratropics there is a corresponding “extratropical tropopause

VERY SHORT-LIVED SUBSTANCES

2.24
layer” (ExTL). In both the tropics and the extratropics,

large-scale and small-scale processes together play important roles in determining the structure of the tropopause

layers. Representing these processes correctly poses particular difficulties for global chemical models. In the following, the tropics, including the tropical tropopause layer,

and both vertical and quasi-horizontal exchange to the

stratosphere, are discussed in Section 2.4.1. The extratropics and the extratropical tropopause layer are discussed

in Section 2.4.2, and the representation of relevant

processes in global models is discussed in Section 2.4.3. 

2.4.1 Tropical Convection, the TTL, and

Tropical Troposphere-to-Stratosphere

Transport

The lower part of the tropical troposphere is dominated by ascent in convective clouds. The outflow from

these clouds is concentrated in two distinct layers, a lower

layer from about 2 to 5 km (800-550 hPa), and an upper

layer from about 10 to 15 km (300-150 hPa). The compensating descent in the regions surrounding the clouds is

associated with evaporative cooling in the lower layer, and

radiative cooling in the upper layer (e.g., Folkins and

Martin, 2005). In the lower part of the troposphere, moist

convection can efficiently transport chemical species vertically via strong convective updrafts, but at the same time,

the scavenging associated with convective precipitation

can efficiently remove the soluble species (Crutzen and

Lawrence, 2000; Mari et al., 2000; Barth et al., 2001).

Studies using explicit modeling of convection have shown

that scavenging in warm clouds should be considered as a

kinetic rather than an equilibrium process (Yin et al., 2001;

Barth et al., 2001). Such treatment is not included in the

vast majority of global chemical models.

So far, the observational data on soluble species in

the entrainment and detrainment zones of isolated or

organized convective systems in the tropics that might

better constrain these model results are still very limited.

In cold clouds, as liquid water freezes and riming occurs

around ice, chemical species are also absorbed into ice

crystals by various processes, such as bubble inclusions

within crystals. A limited number of field laboratory and

theoretical studies have proposed trapping coefficients of

chemicals in ice (e.g., Stuart and Jacobson, 2003; Krämer

et al., 2006; see also Section 2.3.4.1). The current experimental data are apparently inconsistent, poorly understood, and only available for a limited set of chemical

species (e.g., HCl, HBr). Parameterization of scavenging

by ice crystals is only now being implemented in global

models (Kärcher and Basko, 2004). For example, it has

been found that if the gas retention coefficient (the fraction of dissolved gas that is retained as a liquid droplet

freezes) of a highly soluble gas is allowed to vary in a relatively narrow range (0.5-1.0), the corresponding bulk gas

transport within the cloud varies, for maritime clouds, by

as much as a factor of 4 (Yin et al., 2002). See also Section

2.3 for a discussion of washout and solublities of VSL SGs

and PGs. In addition, the microphysical structure of convective clouds and ice clouds is highly sensitive to the

boundary layer aerosol concentration and shows variations between, for example, maritime and continental

clouds (Yin et al., 2002; Kärcher and Lohmann, 2003;

Nober et al., 2003; Ekman et al., 2004). This has implications for the concentration of trace gases inside anvil ice

crystals.

The part of the tropical atmosphere above the level

of maximum convective outflow (at about 12 km altitude,

345K potential temperature), and below the cold point

tropopause (at about 17 km altitude, 380K potential temperature), is now commonly described as the TTL (see,

for example, WMO, 2003; Gettelman and Forster, 2002).

This definition of the TTL, equivalent to the “sub-stratosphere” of Thuburn and Craig (2002), has it as the region

of the troposphere in which there is a transition between

radiative-convective domination and radiative domination

of the thermal balance. The TTL acts as a source region

both for the stratospheric overworld and for the extratropical lowermost stratosphere (see Figure 2-1). Net

exchange from the TTL to the stratospheric overworld in

the tropics, say across the 100 hPa level as an approximation to the cold point tropopause, is regarded as dominated

by large-scale ascent in the Brewer-Dobson circulation,

with the strength of this circulation controlled by largescale dynamical processes (e.g., Holton et al., 1995).

The overall picture of TTL processes is somewhat

updated from WMO (2003), though estimates of transport

time scales between the boundary layer and the lower part

of the TTL are largely unchanged. A useful recent description of the relevant processes is given, for example, in

Folkins and Martin (2005). In the “lower TTL” (12-15

km, 345-360K), convective penetration is sufficiently frequent so that most of the air detraining from convection

descends to the lower troposphere. The upward mass flux

in convective clouds, the magnitude of the radiative

cooling outside the clouds, and hence the magnitude of

the implied descent decays with height. Küpper et al.

(2004) give estimates of the upward cumulus mass flux as

a function of height from various modeling and data

studies, but note that there are significant difficulties in

estimating this flux. At a level z0 (the level of zero radiative heating) of about 15 km or 360K, there is a transition

from clear-sky radiative cooling to clear-sky radiative

VERY SHORT-LIVED SUBSTANCES

2.25
heating. Only above z0 in the “upper TTL” (15-17 km,

360-380K), where the upward cumulus mass flux is comparable to or less than the large-scale upward mass flux in

the Brewer-Dobson circulation, does most air detraining

from clouds ascend into the stratosphere. 

Note at this stage that the TTL has been defined by

some authors (e.g., Folkins et al., 1999; Sherwood and

Dessler, 2001; Füeglistaler et al., 2004; Sinnhuber and

Folkins, 2006) to be only this “upper TTL,” i.e., only as

the region above z0. This definition is based on sound reasoning, but for consistency with WMO (2003) it is not

adopted here, and the terms “lower TTL” and “upper TTL”

are used as introduced previously. 

Gettelman et al. (2004) noted that clouds reach the

typical level z0 (in their case identified by cloud brightness

temperature <200K) only in relatively localized geographical regions, implying that air that is likely to ascend into

the stratosphere is drawn from the surface in similarly localized regions. An equivalent statement is that the transport

time scale from the surface to the upper TTL is likely to be

significantly less in these localized regions than on average

over the tropics. Such regions include the Western

Equatorial Pacific during October to March (with a continuing but lesser role in other months) and the Bay of Bengal,

Southeast Asia, and Panama regions during April to

September (e.g., see Gettelman et al., 2004; Figure 9). 

Somewhat similar conclusions to the above follow

from recent trajectory calculations based on European

Centre for Medium-Range Weather Forecasts (ECMWF)

three-dimensional wind fields. In particular, the trajectory calculations of Füeglistaler et al. (2004) show that air

that ultimately ascends into the tropical lower stratosphere

crosses the lower boundary of the TTL preferentially in

the West Pacific region in Northern Hemisphere winter

and in a region centered on southeast Asia, but extending

toward India and into the West Pacific, in Northern

Hemisphere summer. The longitudinal variation of transport across the lower boundary of the TTL, according to

the Füeglistaler et al. (2004) calculations, is shown in

Figure 2-5. It should be noted that the wind datasets on

which such trajectory calculations are based do not resolve

individual convective cells and, therefore, that any results

need cautious interpretation. Nonetheless, the overall

behavior of the calculated trajectories is broadly as

expected from the schematic picture of the TTL outlined

previously. For example, the probability of ascent of forward trajectories into the stratosphere rapidly increases

with the height of trajectory initialization.

The Füeglistaler et al. (2004) trajectory calculations

predict a residence time (the time for forward trajectories

to move up or down by 10K in terms of potential temperature) that is maximum (about 12 days) for trajectories

starting at 360K (corresponding to z0) and decreases to

about 9 days for a layer centered at 370K. This implies a

residence time for the 360K-380K layer of about 20 days.

This is expected to be an underestimate, since it is expected

that such transport calculations, through their use of vertical winds from meteorological datasets, overestimate

vertical transport rates (both net transport and dispersion).

On the other hand, the time scale of about 80 days to cross

the 360-380K layer that has been inferred from a onedimensional tropical-mean model (Folkins and Martin,

2005) is almost certainly an overestimate, since it neglects

the dispersive effects of geographical variation in vertical

velocity. On the evidence currently available, therefore,

the residence time for air parcels in the 360K-380K layer

is likely to be somewhat greater than 20 days. New calculations based on improved meteorological datasets, or

using radiative transfer codes to calculate vertical velocities, are likely to refine this estimate of residence time in

the near future. Transport time scales for the upper TTL

(20 to 80 days on the basis of the Füeglistaler et al. (2004)

and Folkins and Martin (2005) results mentioned above)

are expected to be significantly longer than those for the

lower TTL, i.e., about 10 days on the basis of typical

cumulus mass fluxes estimated by Küpper et al. (2004). It

is therefore inappropriate to assign a single transport time

scale to the whole TTL extending from the maximum in

convective outflow to the cold point. This point was noted

in WMO (2003), but was confused by also giving estimates for a single turnover or replacement time of 10 to

13 days for the 250 hPa-100 hPa layer (i.e., the whole TTL

by the more common definition).

The consequence of the longer timescale in the

upper TTL for injection of VSLS is not yet clear; there

might be more time for conversion from SG to PG, but the

net effect of this on total halogen entering the stratosphere

is likely to be small. A more important issue is whether

there is significant scavenging of PGs in the upper TTL

associated with uptake on cirrus clouds (see Section 2.3.4),

thus limiting the transport of halogen into the stratosphere

via the PGI pathway. 

As noted previously, upward transport out of the

TTL into the stratosphere is understood to be primarily

controlled by large-scale dynamical processes, and there

has been no reason to revise significantly net flux estimates over those given previously (e.g., Rosenlof, 1995).

However, there have been recent changes in temperatures

of the tropical lower stratosphere (e.g., Randel et al.,

2004), with a significant reduction in temperatures more

recently, i.e., 2000-2004 appears to be 1K cooler than the

period 1995-1999. Implications for stratospheric water

vapor are discussed in Chapter 5 of this Assessment.

Simple dynamical arguments imply that this may be

VERY SHORT-LIVED SUBSTANCES

2.26
explained by a 10% increase in the upward mass flux out

of the TTL, and that this is driven by changes in largescale dynamical processes (Randel et al., 2006). It is,

however, impossible to predict how long the current

anomalously large upward mass flux will persist.

Exchange between the TTL and the extratropical lowermost stratosphere is achieved through quasihorizontal transport associated with synoptic-scale eddies.

This has been apparent for some time on the basis of observations of subtropical filaments in midlatitudes (e.g.,

Vaughan and Timmis, 1998; O’Connor et al., 1999), water

vapor observations (e.g., Dessler et al., 1995) and, more

recently, chemical observations in the extratropics (see

Section 2.4.2). Trajectory and similar transport calculations based on meteorological data also show the possibility of such transport, particularly above the level of the

subtropical jet (Chen, 1995; Haynes and Shuckburgh,

2000). Both direct observations of the chemical composition of the extratropical lowermost stratosphere (e.g.,

Randel et al., 2001; Prados et al., 2003) and transport calculations imply that transport from the TTL to the lowermost stratosphere is, at least in the Northern Hemisphere,

highly seasonal, with much greater transport in summer

than in winter. The air in the extratropical lowermost stratosphere may be regarded as a mixture of air that has been

transported from the TTL, and air that has descended from

the stratospheric overworld. The fraction of air from the

TTL has been estimated by Hoor et al. (2005), on the basis

of observed concentrations of carbon monoxide (CO) and

of the carbon dioxide (CO2) seasonal cycle, to be about

one-third in winter, perhaps increasing to about one-half

in summer. Similar conclusions have been obtained from

other datasets. 

Recent trajectory calculations of Levine et al.

(2006) for Northern Hemisphere winter show that more

than 75% of the air that enters the TTL, and subsequently

reaches the stratosphere, does so through quasi-horizontal

transport into the extratropical stratosphere. It is likely

that this fraction will be larger in summer than in winter.

It is interesting that, according to the Levine et al. (2006)

calculations, the distribution of the entry locations of such

air through the base of the TTL is spread relatively uniformly in longitude, in contrast to the corresponding contribution for air that subsequently reaches the tropical

lower stratosphere. Note also that while the strength of

this quasi-horizontal exchange with the extratropics is not

VERY SHORT-LIVED SUBSTANCES

2.27

Figure 2-5. Results of 2-month

backward trajectory calculations,

using three-dimensional winds from

ECMWF operational analysis data,

reported by Füeglistaler et al. (2004).

The backward trajectories were to

points distributed evenly over the

400K surface in the tropics. The contour lines indicate the density of intersections of the backward trajectories

with the 340K surface (corresponding roughly to the lower boundary

of the TTL), relative to the density of

such intersections if spread uniformly

over the globe. The upper and lower

panels show the density calculated

from trajectories followed over the

periods January-February 2001

(Northern Hemisphere (NH) winter)

and July-August 2001 (NH summer),

respectively. The conclusion is that air parcels reaching the tropical lower stratosphere are in NH winter drawn

preferentially from the lower troposphere in the Western Pacific and, to a lesser extent, the Indian Ocean

region. In the NH summer, the corresponding regions include the West Pacific, but extending significantly over

Southeast Asia. (The color scale indicates the average mixing ratio of water vapor determined by the minimum saturation mixing ratio along each trajectory and assigned to the point of intersection of that trajectory

with the 340K surface. See Chapter 5, Sections 5.2.5 and 5.3.5 for further discussion of water vapor and its

implications for ozone.)

January-February 2001

July-August 2001

[ [
uniform with height, there is little reason to believe that

convective penetration to the level z0 has the same significance for such exchange as it does for exchange with the

stratospheric overworld, since the quasi-horizontal

exchange transport is not strongly affected by the direction of slow vertical motion. 

A final issue concerns the spatial structure of chemical fields within the TTL. There is no reason to expect

chemical fields to be homogeneous. Recent in situ observations have shown complex spatial structure in this region

in chemical species with a whole range of lifetimes (e.g.,

Tuck et al., 2004) apparently resulting from several different mechanisms including different surface source

regions, and transport out of the extratropical stratosphere.

The spatial structure of VSL SGs and PGs are expected to

be correspondingly variable. 

2.4.2 Stratosphere-Troposphere Exchange

in the Extratropics

Transport from the tropical upper troposphere to

the extratropical lower stratosphere has been discussed in

the previous section. There is also significant transport

directly from the extratropical troposphere to the extratropical lower stratosphere, particularly in summer (e.g.,

Hintsa et al., 1998; Pan et al., 1997). However, the high

static stability in the lower stratosphere, and the fact that

the average circulation is downward, limit the vertical penetration of tropospheric air. The result is a finite transition

layer (the ExTL) in which chemical characteristics vary

continuously between the extratropical upper troposphere

and lower stratosphere (see Figure 2-1). In principle, this

transition layer is a region in which concentrations of VSL

SGs and PGs may be relatively high, and in which substantial in situ ozone loss due to VSLS may take place.

However, typical ozone mixing ratios in the ExTL are 300

parts per billion (ppb) or less, which limits the possible

impacts on the ozone column. 

Chemical measurements in the ExTL have

increased significantly since WMO (2003). The location

of the ExTL can be deduced, for example, from ozone,

carbon monoxide (CO), total reactive nitrogen (NOy),

and water vapor (H2O) data collected as part of the

Stratosphere-Troposphere Experiments by Aircraft Measurements (STREAM) and the German Spurenstofftransport in der Tropopausenregion (SPURT) programs

(Hoor et al., 2002; Fischer et al., 2000; Hoor et al., 2004;

Krebsbach et al., 2006; Engel et al., 2006), which also

show that the layer is deeper in summer, with a stronger

tropospheric influence, than in winter. Note that CO has a

short chemical lifetime of 2 to 3 months, making it a useful

proxy for VSLS transport into the lower stratosphere. The

upper limit of the ExTL has been estimated from enhanced

CO and H2O mixing ratios to be about 25K in potential

temperature above the local tropopause defined by the 2

PVU (potential vorticity unit) surface, corresponding to

about 2 km in height. Pan et al. (2004) identified the transition layer using aircraft data and characterized it as centered on the thermal tropopause, and about 2-3 km thick.

Note that the thermal tropopause is usually above the

tropopause defined by the 2 PVU surface. 

Alongside the new observations of chemical

species in the extratropical tropopause region, many

recent studies have used trajectory calculations based on

winds from large-scale meteorological datasets to investigate vertical transport processes, such as in the large

body of work coordinated in the STACCATO (Influence

of Stratosphere-Troposphere Exchange in a Changing

Climate on Atmospheric Transport and Oxidation Capacity) project as summarized in Stohl et al. (2003). An

important conclusion relevant to VSLS has been that

warm conveyor belts associated with synoptic-scale baroclinic eddies, and associated frontal zones, play an important role in rapid transport from the boundary layer to the

tropopause region. One consequence is that rapid upward

transport tends to originate from the boundary layer in

preferential longitudinal regions, such as the midlatitude

storm track regions east of Japan and east of the U.S. (e.g.,

Sprenger and Wernli, 2003; see Figure 2-6 for further

details). Surface emissions (e.g., anthropogenic or coastal

VSLS) in these regions may therefore be more likely to

contribute to the chemical composition of the extratropical tropopause region. The importance of warm conveyor belt transport, and the usefulness of trajectory calculations for studying such transport, have been confirmed in various case studies. For example, Nedelec et

al. (2005) showed very high CO concentrations arising

from Siberian fires observed by MOZAIC (Measurement

of Ozone, Water Vapor, Carbon Monoxide and Nitrogen

Oxides by In-service Airbus Aircraft) aircraft in the lower

stratosphere. 

Trajectory studies have confirmed that many air

parcels moving from the boundary layer to the tropopause

region remain in that region for several days (e.g., Wernli

and Bourqui, 2002; Sprenger and Wernli, 2003), thus

implying that VSL SGs or PGs transported into the very

low stratosphere could have an impact on ozone loss.

Since stretching rates of material surfaces in the extratropical tropopause region (which give an indication of

the rate at which an anomalous air mass is stirred with,

and eventually mixed with, its environment) are of the

order of a few days (e.g., Stohl, 2001), it seems very

unlikely that air masses arriving in the tropopause region

VERY SHORT-LIVED SUBSTANCES

2.28
are predisposed to descend rapidly, but instead are likely

to mix into the background.

Trajectory studies for the SPURT dataset (Hoor et

al., 2004) imply that the composition of the ExTL is influenced by recent transport from the troposphere on time

scales of days to a few weeks. While trajectory studies

might be expected to capture the effects of warm conveyor

belts, they almost certainly miss the effect of smallerscale processes such as deep convection or convection

embedded in frontal systems. The quantitative effect of

deep convective injection of tropospheric chemical species

into the extratropical lower stratosphere remains relatively

uncertain. There are some in situ chemical measurements

that point very clearly to such injection (e.g., Poulida et

al., 1996; Fischer et al., 2003), while other measurements

have given no direct evidence of penetration of extratropical convection into the stratosphere (Ridley et al.,

2004). There are also studies that combine chemical measurements with meteorological or trajectory studies that

argue for such penetration (e.g., Hegglin et al., 2004).

Dessler and Sherwood (2004) argue that convective penetration to the 380K surface in midlatitudes, while much

less frequent than in the tropics, is frequent enough to have

a significant effect on water vapor at these altitudes,

though not on ozone since the relative difference in ozone

concentrations between troposphere and stratosphere is

much smaller than the relative difference in water vapor

concentrations.

Convection might contribute to troposphere to

stratosphere transport (TST) directly by overshooting

the tropopause at the anvil outflow (Mullendore et al.,

2005), or indirectly by exciting gravity waves that break

significantly above the convective clouds (Wang, 2003).

Further, recent evidence for convective transport of

low-altitude air into the lowermost stratosphere in the

extratropics is provided by in situ observations in the

stratosphere that show clear evidence of biomass

burning products such as CO and aerosols, apparently

arising from forest fires over Canada (e.g., Ray et al.,

2004). Fromm and Servranckx (2003) have suggested

that enhanced convective cells associated with the fires

themselves may be very likely to penetrate the lower

stratosphere; terms such as “pyro-convection” or “pyrocumulonimbus” are becoming widely used to describe

this phenomenon (e.g., Fromm et al., 2005). While the

occasional penetration of pyro-convection events into

the lower stratosphere has been observed, it is not clear

at present that the pyro-convection has a significant

effect on the overall chemical balance of the extratropical lowermost stratosphere.

VERY SHORT-LIVED SUBSTANCES

2.29

GM

a)

deep STT

destinations

winter

GM

b)

deep TST

origins

winter

Figure 2-6. Geographical distribution of the (a) low tropospheric “destinations” of deep STT and (b) low tropospheric “origins” of deep TST events during winter. Values are in % and correspond to the probability that a

low tropospheric air parcel (pressure > 700 hPa) was transported downward from the lowermost stratosphere

during the previous four days (deep STT) or will be transported upwards into the lowermost stratosphere

during the next four days (deep TST). The implication of (b) is that the probability of rapid transport from the

surface to the tropopause region is strongly dependent on geographical location (at the surface). From

Sprenger and Wernli (2003).
2.4.3 Predictive Modeling of the

Tropopause Region

Predictive modeling of the impact of VSLS on ozone

depends on model representation of the processes discussed

in Sections 2.4.1 and 2.4.2. In the tropics, the greatest challenge for models is representing convective processes,

although in the TTL the subtle interplay between different

processes, and the fact that significant changes in dynamical and chemical characteristics take place over 2-3 km,

also present a challenge. In the extratropics, the transport

characteristics of synoptic-scale eddies, and their interplay

with smaller-scale processes, must be captured. These

smaller-scale processes include gravity wave breaking,

turbulence in the vicinity of jet streams, and radiative

processes associated with upper level clouds. From the

chemistry point of view, the representation of soluble

species, as well as aerosols and microphysical processes, is

still poor in global models. Scaling up the heterogeneous

reactions on ice and aerosols to realistic parameterizations

is still a major challenge for prediction of the tropopause

region. Uncertainties in surface and lightning emissions

may cause significant biases in simulating the chemical

composition of the upper troposphere.

The parameterization of tropical convective processes in global-scale models remains a conceptual challenge, with different theoretical approaches possible.

Basic thermodynamics are not well characterized, while

transport and scavenging processes are even less well

understood. One widely used method for parameterizing convective transport processes is based on the

mass-flux formulation (Arakawa and Schubert, 1974),

in which transport processes are represented by an

ensemble of vertical mass fluxes at subgrid scales. This

formulation however was originally developed for representing the subgrid-scale thermodynamic effects

(phase changes and fluxes of heat and moisture) of convection, and its generalization to chemical transport

processes is not straightforward (Donner et al., 2001).

Lawrence and Rasch (2005) have shown that, for a single

convection parameterization, differences in the implementation of detrainment and entrainment fluxes can

result in significant differences in the upper tropospheric

mixing ratios of short-lived tracers (lifetimes of a few

days or less). In particular, the bulk formulation of convection, widely used in global models, in which entrainment and detrainment are each assumed to occur at

a single level, may lead to underestimation of up to 30%

in zonal and monthly mean concentrations of such

tracers, relative to, for example, a more sophisticated

formulation in which a distribution of entrainment and

detrainment levels is assumed. One implication of the

above is that VSLS could play an even more important

role in the upper troposphere than previously thought.

Detailed evaluation of cumulus parameterizations,

in comparison with observations and cloud resolving

models, have recently taken place both with respect to

thermodynamics (Yano et al., 2004) and with respect to

tracers (Bell et al., 2002; Collins et al., 2002). More accurate estimates of mass fluxes, based on high-resolution

data from either observations or modeling, could lead to

improved parameterizations of these convective transports. Evaluation of the mass flux formulation against

high-resolution spatial data obtained from explicit numerical models reveals that, in its standard form, it substantially underestimates vertical transport of heat, moisture,

and momentum by deep convection (Yano et al., 2004),

with improvement possible through nonlinear averaging

of the large-scale variables. Similar issues will almost certainly arise in the convective transport of chemical tracers. 

Notwithstanding these recent results on convective

transport and scavenging effects (see Section 2.4.1), current global models necessarily make significant simplifications in their representation of convective transport and

scavenging processes. In some models the two are linked

(e.g., Considine et al., 2005), whereas in others the two are

completely separate. Lack of relevant observations and

model simulations means that there has been no systematic

validation of model representation of convective processes

acting on halogenated VSLS against observations. Historically, model representations of such processes have

been validated against observations of radon and lead

mixing ratios. Lead-210 (210Pb) is produced in the atmosphere by radioactive decay of radon-222 (222Rn), which

has a radioactive half-life of 3.8 days. Lead-210, which

has a radioactive half-life of about 22 years and can therefore be considered radioactively stable for this purpose, is

removed by attachment to aerosols and washout. Most

recently, building on the large amount of previous work on

this subject, Considine et al. (2005) have presented an intercomparison of 222Rn and 210Pb calculated by the NASA

Goddard GMI (Global Modeling Initiative) model using

meteorological fields from three different general circulation models (GCMs)/data assimilation systems, and comparison with observations. The three different simulations

show some common discrepancies relative to observations

and the authors deduce, in particular, that midlatitude convective transport may represent detrainment at too great an

altitude, and that cloud scavenging frequency in midlatitudes may be too great. Convective transport is confirmed

to play a very important role in setting upper troposphere

and lower stratosphere concentrations that are in good

agreement with observations, though part of this role may

VERY SHORT-LIVED SUBSTANCES

2.30
be indirect, for example, in bringing species from the

boundary layer to the upper troposphere, from which they

can be transported effectively by the larger-scale circulations. These findings are all potentially relevant to model

simulations of halogenated VSLS, though it is difficult to

extrapolate results directly to model skill for predicting

VSLS, since the chemical properties and the location and

time of emissions of the VSLS will be different from these

radioactive tracers. 

In the extratropics, the role of convection is not as

great as in the tropics, and the fact that many of the relevant processes are relatively large scale might seem to

make this region more accessible to global chemical transport models and global chemistry-climate models. The

success of trajectory studies based on meteorological wind

fields in reproducing chemical distributions observed in

field experiments is encouraging in this regard. However,

“operational” resolutions used for global chemical models

are still relatively coarse. Comparison between global

models and chemical data (e.g., Brunner et al., 2003) has

revealed the limitations of such models that, for example,

overestimate the two-way transport in the tropopause

region, resulting in excessive ozone in the upper troposphere and excessive CO in the lower stratosphere. This

is an indication that, for current models, useful quantitative prediction of the effect of halogens from VSLS penetrating the lower stratosphere in the extratropics remains

difficult.

2.5 CONTRIBUTION OF HALOGENATED VSLS

TO STRATOSPHERIC HALOGEN LOADING

2.5.1 Source Gas Injection (SGI) and

Product Gas Injection (PGI)

The concept of source gas injection (SGI) and

product gas injection (PGI) into the stratosphere was

described in Chapter 2 of the last Assessment (Ko and

Poulet et al., 2003) as well as the introduction to this

chapter. For SGI, the VSL source gas (SG) is transported

to the stratosphere and then reacts to release halogen

atoms. Supply to the stratosphere depends on emission

location and time, as well as details of the tropospheric

transport (e.g., convection) and troposphere-to-stratosphere exchange. For PGI, either intermediate (i.e.,

organic) products or final (i.e., inorganic) products are

produced in the troposphere, and these species are transported to the stratosphere. The details of PGI depend on

the same factors as SGI, plus factors such as the photochemistry of the source species, intermediate products,

and final products. The gas-phase partitioning between

radical and longer lived inorganic reservoirs, as well as

the solubility and heterogeneous reactivity of all species,

are critical for assessing the efficiency of PGI (see Section

2.3). Indeed, differences in the efficiency of stratospheric

supply of chlorine, bromine, and iodine from VSL organics

might be due to differences in the photochemical properties of the respective inorganic products. Another important consequence of PGI is that measurements of the

organic VSL SGs alone are not sufficient to determine the

efficiency of stratospheric injection.

Injection of tropospheric particles into the stratosphere is another mechanism for cross-tropopause transport of halogens, depending on the chemical composition

of the particles. The substantial organic content of many

aerosol particles just above the tropopause suggests injection of tropospheric particles into the stratosphere, and the

presence of Br on these particles suggests that crosstropopause transport of bromine, in both directions, can

occur by aerosols (Murphy et al., 1998; Murphy and

Thomson, 2000). However, observations of this process

are limited and, as a consequence, quantification of crosstropopause halogen transport by aerosols is a considerable gap in our present state of knowledge. Most model

simulations do not include cross-tropopause transport of

halogens by aerosols in their estimates of PGI.

2.5.1.1 OBSERVATIONAL EVIDENCE FOR SGI 

Bromine

Bromoform (CHBr3) is thought to be the most

abundant bromine-bearing VSL compound near the surface (Table 2-2). As noted previously in this chapter,

CHBr3 levels can reach or exceed 40 ppt near sea level

(Yokouchi et al., 2005b). Observations reveal between

~0.2 ppt (Schauffler et al., 1998) and 1 ppt (Sturges et al.,

2000) of CHBr3 in the tropical free troposphere (see also

Table 2-2). Measurements in the upper troposphere and

tropopause region by Sturges et al. (2000) indicated

CHBr3 abundances of 0.5 ± 0.2 ppt for data collected

during three flights (June and August 1998 and March

1999). Schauffler et al. (1998) measured 0.1 ppt of CHBr3

at the tropical tropopause during February 1996. However, CHBr3 was only detected at the tropical tropopause

for one of three flight series that sampled this region.

Bromoform exhibits similar variability near the midlatitude tropopause, ranging from near zero to a maximum

abundance of ~0.7 ppt (Schauffler et al., 1999). The crosstropopause transport of CHBr3 is difficult to quantify from

observations because of atmospheric variability and limited sampling. However, there are indications that, on

VERY SHORT-LIVED SUBSTANCES

2.31
occasions, values of CHBr3 are high enough near the

tropical tropopause for SGI of CHBr3 to be important for

the bromine budget of the lowermost stratosphere.

Specifically, if the approximately 0.5 ppt of CHBr3 measured near the tropical tropopause by Sturges et al. (2000)

turns out to be representative of the average mixing ratio

in the TTL, it would provide an SGI flux that could maintain an additional 1.5 ppt of Bry.

The VSL species CH2Br2 and CH2BrCl are less

abundant near the surface than CHBr3, but are longer lived

and therefore more likely to enter the stratosphere via SGI

(Tables 2-1 and 2-3). Wamsley et al. (1998) showed that

CH2Br2 and CH2BrCl cross the tropopause, supplying ~2.3

ppt of Bry by the SGI pathway. The total bromine content

of the VSL species CH2Br2, CH2ClBr, CHCl2Br, CHClBr2,

and CHBr3 was measured to be 2.2 ppt at the tropical

tropopause during February 1996 (Schauffler et al., 1998)

(see also the compilation in Table 2-2, which concludes a

higher value of 3.5 ppt of bromine from these species).

Plots of CH2BrCl versus CFC-11, and CH2Br2 versus CFC11, in the Wamsley study show that bromine from these

compounds is released in the lowermost stratosphere

(LMS), prior to the release of the majority of bromine from

CH3Br and halons. Bromine release in the LMS has important consequences for the photochemistry of this region.

Significant supply of inorganic bromine (Bry) leads to

much faster ozone loss by two catalytic cycles represented

by the rate-limiting steps defined by the BrO plus ClO,

and BrO plus hydroperoxy radical (HO2) reactions, compared with calculations that consider supply of bromine

only from CH3Br plus halons (Salawitch et al., 2005). A

significant role for SGI injection of CH2BrCl and CH2Br2

at all times, and lack thereof for CHBr3 under some conditions, is consistent with their lifetimes for photochemical

removal in the troposphere of 150, 120, and 26 days,

respectively (Table 2-1).

Chlorine

Contributions to stratospheric inorganic chlorine

(Cly) from VSL SGs are dominated by the species CHCl3,

CH2ClCH2Cl, CH2Cl2, C2H5Cl, and C2Cl4 and can potentially supply about 55 ppt of chlorine to the stratosphere

(Table 2-2). These gases have natural and anthropogenic

sources, as described in Section 2.2. The tropospheric lifetimes of these gases (30 to 150 days) are comparable to

those for some of the brominated VSL SGs (see Table 2-

1), so if SGI is important for bromine, it may also play a

role for chlorine. Although the overall perturbation of Cly

by VSLS is small (because the chlorine content of longer

lived gases is much larger than that of the VSLS), this perturbation could be significant for the LMS, where the VSL

SGs are more likely to release their chlorine content. This

could lead to significantly higher levels of ClO in the LMS

than would be present in models that only consider supply

of Cly from long-lived gases.

Iodine

The supply of stratospheric iodine presently appears

to be dominated by methyl iodide (CH3I) (e.g., Davis et

al., 1996; Table 2-2). Abundances of CH3I range from

~0.8 ppt in the marine boundary layer to ~0.08 ppt in the

tropical upper troposphere (Table 2-2). Higher values of

CH3I, occasionally reaching ~1.0 ppt, were reported in the

subtropical upper troposphere (Davis et al., 1996). The

much shorter lifetimes of iodinated VSLS compared with

most brominated VSLS (Table 2-1) suggests, in general,

less efficient SGI of iodine to the stratosphere.

2.5.1.2 OBSERVATIONAL EVIDENCE FOR PGI

Bromine

PGI refers to cross-tropopause transport of degradation products of VSL source gases, as well as final

inorganic products. For bromine, the discussion of

observational evidence for PGI begins with the suggestion of a global, ubiquitous background abundance of 1

to 2 ppt of bromine monoxide (BrO) distributed

throughout the troposphere (Harder et al., 1998; Müller

et al., 2002; Richter et al., 2002; Van Roozendael et al.,

2002). Model simulations indicate that Bry levels of 4

to 8 ppt are needed in the middle troposphere to sustain

this much BrO (Yang et al., 2005). If 4 to 8 ppt of Bry is

present in the global troposphere, then cross-tropopause

transport of these species could dominate the bromine

budget of the LMS (Ko et al., 1997; Pfeilsticker et al.,

2000). As explained below, it is unclear whether BrO is

present at this level in the whole troposphere. There is

no evidence of enough inorganic chlorine or iodine in

the global troposphere to perturb stratospheric Cly or Iy

in a manner comparable to the potential perturbation to

stratospheric Bry by tropospheric, inorganic bromine.

The presence of 1 to 2 ppt of BrO in the troposphere

was first suggested based on comparison of total column

BrO from the Global Ozone Monitoring Experiment

(GOME) to stratospheric columns inferred from balloon

measurements of BrO profiles (Harder et al., 1998).

Observations of the solar zenith angle (SZA) variation of

differential slant column density of GOME BrO also indicated that much of the column measured by GOME lies

below the tropopause, perhaps due to the presence of 2 to

3 ppt of BrO between the tropopause and the surface

VERY SHORT-LIVED SUBSTANCES

2.32
(Müller et al., 2002; Van Roozendael et al., 2002).

Measurements of BrO and O4 (the O2-dimer) columns over

the Pacific suggest 0.5 to 2.0 ppt of BrO uniformly distributed throughout the troposphere (Richter et al., 2002).

A global, ubiquitous distribution of 2 to 3 ppt of BrO would

account nearly completely for the “additional” stratospheric bromine reported by GOME (Van Roozendael et

al., 2002; Salawitch et al., 2005). The fact that excess BrO

is seen year round by GOME, at all latitudes (e.g., Chance,

1998; Hegels et al., 1998), argues against the springtime

release of bromine from sea-ice and frost flowers (the socalled bromine explosion; see, for example McElroy et

al., 1999) being the primary source of tropospheric background BrO.

Other observations, however, contradict the notion

of a global, background level of BrO at an abundance of 2

to 3 ppt. Ground-based observations of direct and scattered sunlight reveal mean abundances of BrO in the

troposphere over Lauder, New Zealand (45°S), of ~0.2

ppt and an upper limit of 0.9 ppt over this site (Schofield

et al., 2004). Low values of tropospheric BrO are also

observed over Arrival Heights, Antarctica (78°S), except

for bromine explosion events (Schofield et al., 2006).

Balloon profiles of tropospheric BrO indicate levels that

peak at ~2.3 ppt, but more commonly lie between ~0.4 and

0.6 ppt (Fitzenberger et al., 2000). If BrO levels in the

global free troposphere are less than 1 ppt, then the GOME

observations of total BrO columns imply larger stratospheric BrO columns than can be accounted for by that

supplied by CH3Br and halons (Salawitch et al., 2005).

The role of VSLS in maintaining the global tropospheric background levels of BrO and total Bry is a subject of active research. Von Glasow et al. (2004) showed

that the stratosphere-to-troposphere flux of inorganic

bromine combined with the tropospheric decomposition

of CH3Br was insufficient to maintain tropospheric levels

of BrO approaching ~2 ppt. They examined the sensitivity

of tropospheric BrO to various latitudinal distributions of

surface emissions of VSL SGs. Although the resulting distributions of BrO and Bry differed, the available data were

too sparse and uncertain to place strong constraints on the

geographic distribution of surface emissions. Yang et al.

(2005) presented model results indicating that sea salt

sources dominate the supply of bromine to the lower troposphere, and that VSL SGs are responsible for the

majority of bromine in the upper troposphere. Warwick et

al. (2006) suggested, based on a comparison of modeled

and measured CHBr3, that a large portion of CHBr3 emissions occurs in the tropics. They also showed that emissions of VSL SGs (i.e., CHBr3, CH2Br2, CH2BrCl,

CHBrCl2), together with CH3Br, could maintain ~2 ppt of

Bry in the tropical free troposphere, and between 0 and 2

ppt of Bry in the rest of the troposphere. Levels of Bry

were lower in the rest of the troposphere in this simulation, which involved only VSL SGs, because of aerosol

washout of product gas and the co-location of surface

emissions and convection in the tropics.

There have been no atmospheric observations of

organic degradation products of VSL SG decomposition

to date, such as COBr2 and COHBr from decomposition

of CHBr3 (Section 2.3). Most modeling studies have

assumed direct conversion to inorganic bromine upon the

decomposition of CHBr3 and other VSL SGs. Field observations to assess the accuracy of this assumption represent another gap in our present knowledge.

Chlorine

In contrast to the situation for bromine and iodine,

the tropospheric abundance of Cly is likely dominated by

hydrogen chloride (HCl). The reason for the different

behavior of the halogen families is that production of HCl

from Cl plus methane (CH4) is exothermic, while production of hydrogen bromide and hydrogen iodide (HBr and

HI) by reactions involving CH4 is endothermic. As a result,

PGI for chlorine should be defined by the troposphere-tostratosphere transport of HCl. Efficient scavenging of HCl

by clouds and aerosols is generally considered to lead to

negligible PGI of inorganic chlorine species. Measurements of HCl in the subtropical upper troposphere reveal

abundances between 0 and 80 ppt, but simultaneous

observations of other species indicate these levels of HCl

are maintained by stratosphere-to-troposphere transport

of Cly, rather than tropospheric degradation of VSL SGs

(Marcy et al., 2004).

There are many organic decomposition products

of chlorinated VSLS, most of which have not been measured in the atmosphere to date. An important exception

is phosgene (COCl2), produced by the decomposition of

chlorinated organic gases (see Section 2.3.3). COCl2

has a significant abundance in the upper tropical troposphere (on average 22.5 ppt, or 45 ppt as chlorine,

Table 2-2), and a longer photochemical lifetime than

some of its VSLS parent compounds (Section 2.3.3).

There are no known surface sources of COCl2. If COCl2

were present in the TTL, where its removal processes

are inefficient, it could potentially play a role in the PGI

of chlorinated VSLS.

In this Assessment we have not, however, included

phosgene in our evaluation of the contribution of VSLS

to stratospheric Cly. One reason for COCl2 not being

included is that tropospheric COCl2 is evidently produced

by the decomposition of both long-lived and VSL source

gases (Section 2.3.3). In the case of degradation of longVERY SHORT-LIVED SUBSTANCES

2.33
lived gases that decompose primarily in the stratosphere,

the majority of the tropospheric COCl2 contribution is

transported from the stratosphere (Kindler et al., 1995).

Consequently, upper tropospheric measurements of COCl2

do not necessarily define the contribution from degradation of VSL source gases alone to its abundance. The possibility exists for some degree of “double-counting” of Cly

if COCl2 is included in a summation of chlorine derived

from the tropospheric abundances of all source gases. In

addition, Kindler et al. (1995), based on a 1-D model,

stated that only about 0.4% of COCl2 produced in the troposphere is transported to, and decomposes in, the stratosphere. Measured stratospheric profiles of COCl2 presented by Toon et al. (2001), however, appear to contradict this conclusion, since they show a much larger decline

in mixing ratio with altitude than the model profiles of

Kindler et al. (1995), and are instead consistent with the

expected stratospheric loss of this compound discussed in

Section 2.3.3. More detailed modeling and measurement

studies of COCl2 are needed to quantify whether or not it

represents a significant route for PGI of stratospheric chlorine from VSL source gases. For the present we conclude

that the contribution of COCl2 to “additional” stratospheric

Cly may range from a negligible amount (based on Kindler

et al., 1995) to an upper limit of 40-50 ppt based on

observed upper tropospheric abundances (Table 2-2 and

Section 2.2.1).

Iodine

Both I2 and IO have been observed in the marine

boundary layer (Section 2.2.1). The tropospheric photochemical lifetimes of I2 and IO are very short (Saiz-Lopez

and Plane, 2004), and iodine oxides may self-nucleate,

leading to efficient atmospheric removal by aerosols

(Hoffmann et al., 2001). As a result, ozone depletion

models do not typically consider transfer of inorganic

iodine species across the tropopause, even though these

species are abundant in parts of the boundary layer.

Product gas injection of decomposition products of

CH3I could be important, particularly if rapid convection

lofts air parcels containing CH3I into the TTL, where

aerosol and cloud washout may be inefficient. In the

future, decomposition products of CF3I may also have to

be considered (Section 2.2.4.3). However, no field observations of CH3I and CF3I decomposition products are

available. The present state of knowledge for iodine is

focused on defining stratospheric levels of Iy, rather than

quantification of PGI sources.

2.5.1.3 MODEL ESTIMATES OF SGI AND PGI

The SGI and PGI fluxes can be calculated from

global model simulations of a single source gas and their

products by tagging and keeping separate the degradation

products produced in the stratosphere from those produced

in the troposphere. The lack of atmospheric observations

of the organic intermediate species, combined with limited knowledge of the kinetics of these intermediates and

the recycling of inorganic halogen compounds, means estimates of PGI injection presently rely on parameterizations

of these processes that are untested against observations.

The studies described below, updating progress since the

last Assessment, consider SGI and PGI of bromine. 

Prior to this Assessment, calculations involving

two-dimensional (2-D) (Dvortsov et al., 1999) and 3-D

(Nielsen and Douglass, 2001) models showed that oceanic

release of CHBr3 could maintain about 1 ppt of bromine

to the stratosphere. These calculations assumed a 10-day

lifetime due to “washout” by aerosols and clouds for the

inorganic products produced following the decomposition

of CHBr3. 

Sinnhuber and Folkins (2006) used a onedimensional mechanistic model of the tropical troposphere to quantify stratospheric supply of Bry from

CHBr3 as a function of aerosol and cloud washout lifetime. In the model, Bry reaching the stratosphere is

determined by the detrainment from convection of

CHBr3 above the level of zero radiative heating (z0, see

Section 2.4.1). Convective mass fluxes are constrained

by observed temperature and humidity. Rapid washout

of Bry in the convective cloud was assumed, resulting in

no detrainment of Bry (i.e., product gas). Product gas is

therefore formed from degradation of source gas following detrainment. A first order, uniform washout lifetime for Bry in the clear-sky atmosphere below 17 km

was treated as an adjustable parameter. By assuming

rapid washout for Bry in a particular simulation, they

showed that injection of Bry from the SGI pathway maintains about 0.5 ppt of Bry per ppt of surface CHBr3, generally consistent with the results of Dvortsov et al.

(1999) and Nielsen and Douglass (2001). They also

examined the sensitivity of PGI to washout lifetime. The

PGI of Bry was 0.3 ppt, 0.8 ppt, and 1.6 ppt, per ppt of

surface CHBr3, for lifetimes of 10, 30, and 100 days,

respectively (Figure 2-7). They discussed how the

washout of soluble species from the TTL depends on

details of ice microphysics and could, in principle, be

inefficient if air detraining from deep convection is dry.

VERY SHORT-LIVED SUBSTANCES

2.34
The delivery of total Bry via both SGI and PGI routes

varied from 0.8 to 2.1 ppt, per ppt of surface CHBr3, as

the washout lifetime increased from 10 to 100 days. 

It is also useful to use model results to diagnose the

SGI and PGI flux as a percentage of the flux of VSLS

emitted at the surface, as opposed to relating it to mixing

ratios as in the above study. Warwick et al. (2006) conducted a 3-D model simulation considering surface fluxes

of CHBr3 and other VSL SGs. Convection in this model

was parameterized using the mass flux scheme of Tiedtke

(1989) and included convective updrafts and large-scale

subsidence associated with deep and shallow convection.

An average tropospheric lifetime of 10 days relative to

rainout was assumed for inorganic bromine. Among the

scenarios were two cases where the CHBr3 emission flux

was either uniform in the open ocean or was limited to the

tropical coastlines. As discussed in Section 2.3.2, the calculated lifetimes for the open ocean case are 30-37 days,

compared with 15 days for the tropical coastline case. The

SGI fluxes in the two simulations were practically the

same at about 0.7% of the total surface emission flux. The

PGI fluxes were 4% and 5% respectively. The Warwick

et al. study also showed that CHBr3, together with the

VSLS CH2Br2, CH2BrCl, CHBr2Cl, and CHBrCl2, could

together maintain 3 to 4 ppt Bry in the tropical lower

stratosphere.

The Atmospheric and Environmental Research, Inc.

(AER) 2-D model (Rinsland et al., 2003) was also used to

simulate CHBr3. The model used the convection parameterization from Dvortsov et al. (1998). Washout in the

model was parameterized by a first-order removal rate that

was constant in time. For a fixed uniform surface mixing

ratio boundary condition, the SGI flux was 0.3% of the

surface emission flux and the residence time of SGI Bry in

the stratosphere was 1.4 years. Assuming the same residence time, the PGI flux was 0.6%, about twice the SGI

value. Note that in the Dvortsov et al. (1999) and the

Nielsen and Douglass (2001) studies, the PGI flux was

three times and twice the SGI flux, respectively. These

studies confirm the suggestion that the sum of the two (PGI

+ SGI fluxes) could be as large as a few percent of surface

emissions. As discussed in Section 2.4.3, the calculated

mixing ratio of the product gas in the stratosphere is related

to the SGI and PGI fluxes through the residence time in

the model stratosphere. It is important to note that the

larger fluxes calculated in Warwick et al. (2006) (a few

percent versus less than 1%) were accompanied by smaller

residence times for the stratosphere (0.2 years versus about

1.4 years). This may be related to whether diagnosed

fluxes represent the net fluxes across the tropopause. 

Finally it should be noted that the aforementioned

model studies do not include possible heterogeneous reactions (Platt and Hönninger, 2003) that liberate inorganic

bromine back to the gas phase (Section 2.3.4), which might

result in more efficient PGI of Bry from CHBr3 than was

considered in past Assessments.

2.5.2 Estimates of VSLS Contributions to

Stratospheric Halogens Based on

Measurements of Inorganic Species

This section focuses on the contributions of VSL

halocarbons to the stratospheric budgets of bromine, chlorine, and iodine, based on measurements of inorganic

species. For chlorine and bromine, the material is focused

mainly on studies that compare the total halogen content

of inorganic species with the halogen content of organics.

Typically the difference between the inorganic and organic

budgets is computed for an organic budget containing just

long-lived SGs. The resulting difference is then interpreted as being the contribution to the stratospheric

VERY SHORT-LIVED SUBSTANCES

2.35

0.5 1 1.5 2 2.5 3 3.5 14

15

16

17

18

19

0 10 30 100 ∞

Total bromine [ppt]

Altitude [km]

z

0

cold point

Figure 2-7. Calculated profile of total bromine from

bromoform present as both product gas (Bry) and

source gas (CHBr3), defined as Bry + 3 × CHBr3, for

different values of the lifetime of Bry in days (numbers given in graph), based on a one-dimensional

model of the tropical atmosphere. The calculations

assume 1 ppt of CHBr3 in the boundary layer and that

no Bry is detrained from convection (i.e., model

assumes efficient uptake and removal of Bry in

clouds). Locations of the level of zero radiative

heating and the cold point tropopause in the model

are indicated. From Sinnhuber and Folkins (2006).
halogen loading by VSLS. This method is unable to distinguish between SGI and PGI, and would also include

contributions from product gases that do not originate from

VSLS. For bromine the long-lived SG budget is defined

by the total bromine content of CH3Br and the three or four

halons that are commonly used to define baseline halogen

scenarios for this and previous Assessments. For chlorine

species the long-lived SG budget is defined by the 12

species (CFC-11, CFC-12, CFC-113, CFC-114, CFC-115,

CCl4, CH3CCl3, HCFC-22, HCFC-141b, HCFC-142b,

halon-1211, and CH3Cl; see Chapter 1) used to define the

baseline halogen loading scenarios in this and prior

Assessments. As defined in the introduction, the terms

Bry

VSLS and ClyVSLS refer to the component of stratospheric

Bry and Cly, respectively, due to SGI and PGI of sources

other than these long-lived SGs. For iodine, the discussion is focused on quantification of total stratospheric inorganic iodine (Iy) based on measurements of IO, since

supply is thought to be exclusively from VSL SGs. These

approaches represent the present state of knowledge

regarding quantification of the contribution of VSLS on

the stratospheric halogen loadings.

2.5.2.1 BROMINE

Prior to this Assessment, balloonborne observations

of BrO in the middle stratosphere at midlatitudes were

shown to be consistent with stratospheric Bry of 21.5 ± 3

ppt (Pfeilsticker et al., 2000). The total amount of organic

bromine in the upper troposphere reported in this study

was 18.4 (+1.8 / −1.5) ppt, based on measurements of

CH3Br, four halons, and a variety of VSL SGs. The higher

levels of Bry inferred from BrO, compared with Bry

inferred from the suite of organics, suggested either direct

transport of VSL PGs or inorganic bromine across the

tropopause, or perhaps larger abundances of VSL SGs in

the tropics compared with midlatitudes where the balloon

observations were obtained. Ground-based observations

of column BrO from many stations were shown to be consistent with a total stratospheric bromine abundance of

20 ± 4 ppt, reflecting a ~5 ppt contribution from VSLS

(Sinnhuber et al., 2002). Also in the last Assessment, a

time series of Bry inferred from BrO measurements in the

middle stratosphere over several years was shown to

exceed Bry from CH3Br + halons by an amount consistent

with supply of 4 to 5 ppt of Bry from VSLS (Figure 1-8,

WMO, 2003). However, the interpretation of these data

in the last Assessment was tentative, with the possibility

of “calibration errors either in the measurements of BrO

or the organic source gases” being noted as a possible

explanation of the observations.

Since the last Assessment, there have been many

studies, described below, that quantify Bry

VSLS. Most

of these studies, based on measurements of BrO by

many techniques, conclude that stratospheric Bry exceeds

the amount of bromine that can be supplied solely by

CH3Br and halons. Therefore, since the last Assessment,

we have gained greater confidence that the stratospheric

bromine budget is affected by cross-tropopause transport

of brominated VSLS, by either the SGI and/or PGI routes,

and that the difference noted in the last Assessment is not

due to “calibration errors.”

Figure 2-3 updates the time series of stratospheric

and tropospheric bromine shown in the last Assessment

(Figure 1-8, WMO, 2003). Observations of BrO using

balloonborne differential optical absorption spectroscopy

(DOAS) have been used to estimate stratospheric Bry

(Dorf, 2005; Dorf et al., 2006a). The most accurate

analysis is based on estimating the mixing ratio of BrO for

the middle stratosphere from the slope of BrO slant column

versus air column above the balloon float altitude (the

“Langley” method) and then using a photochemical model

to deduce Bry from measured BrO. Results for five flights,

from 1996 to 2005, are illustrated (points labeled “DOAS

BrO Langley obs.”) in Figure 2-3. Numerical values are

also given in Table 2-7. Estimates of Bry

VSLS are based on

the difference between Bry inferred from BrO and the

bromine content of CH3Br and halons (Bry

CH3Br+Halons).

The estimate of Bry

CH3Br+Halons considers age of air, inferred

from measured tracers, and neglects tropospheric gradients (i.e., it assumes that the sum of bromine from source

and associated product gases at the tropical tropopause,

notably in the case of CH3Br, is represented by the mean

of observed global surface mixing ratios). The difference

between Bry inferred from BrO and Bry

CH3Br+Halons yields

the estimated value for Bry

VSLS (bottom row, Table 2-7).

The mean and standard deviation for Bry

VSLS, considering

all five flights, is 4.1 ± 1 ppt. The overall uncertainty of

this estimate, which is dominated by the accuracy of the

BrO measurements and by the kinetics used to estimate

Bry from BrO, is ±2.5 ppt (see Table 2-7 for discussion of

uncertainties). Values of Bry

VSLS are illustrated in Figure

2-3 by the difference between the time series of Bry

inferred from BrO and Bry

CH3Br+Halons.

Other DOAS-based estimates of Bry are also

shown in Figure 2-3. These estimates, labeled “DOAS

BrO profiles,” are derived from model interpretation of

BrO profiles in the lower stratosphere. The DOAS BrO

profile estimate is considered to be a less reliable method

for inferring Bry

VSLS compared with the Langley method

(Dorf, 2005). It is reassuring that the two methods give

good agreement, since this shows an overall level of conVERY SHORT-LIVED SUBSTANCES

2.36
sistency between the two methods that sample different

parts of the atmosphere. DOAS BrO profile estimates are

also available for more fights than the estimates based on

the Langley method, resulting in a more complete time

series for stratospheric Bry. The fact that values of Bry in

the lower and middle stratosphere are similar lends credence to the notion that VSLS contribute significantly to

stratospheric Bry.

The values of Bry inferred from the balloon observations of BrO shown in Figure 2-3 imply, by an amount

that is larger than the overall uncertainty, a significant contribution to stratospheric Bry from VSLS. In addition, the

values of Bry observed during the past few years indicate

a “leveling off” that appears to be consistent with the

behavior of Bry

CH3Br+Halons (Dorf et al., 2006b). This suggests that Bry

VSLS has not changed dramatically between

1996 and 2005, the time period of the stratospheric observations, and that changes in the surface emissions of

CH3Br and halons are being reflected in the stratospheric

abundance of Bry.

As noted above, Figure 2-3 neglects tropospheric

gradients of CH3Br that may, at least partly, result from its

tropospheric degradation. The mixing ratios of CH3Br in

the TTL may be about 7% smaller than the globally averaged surface value (Schauffler et al., 1999; Montzka et al.,

2003). If the bromine compounds resulting from tropospheric decomposition of CH3Br do not reach the stratosphere, then estimates of Bry

VSLS that rely on surface

mixing ratios of CH3Br may consequently be too low by

up to 0.6 ppt.

Salawitch et al. (2005) examined GOME column

BrO as well as measurements of BrO from aircraft and

balloons. Their study suggested Bry

VSLS between 4 and 10

ppt; best agreement with aircraft data was achieved for a

value of 6.9 ppt, while a balloon profile of BrO in the

summer tropics (Pundt et al., 2002) suggested a value

closer to 10 ppt. Sinnhuber et al. (2002, 2005) showed

that the gas-phase reaction of bromine nitrate (BrONO2)

with oxygen atom triplet (O(3P)) proceeds quickly enough

(Soller et al., 2001) to affect the partitioning of Bry species.

Nearly all of the studies presented here consider this reaction, even though it is not included in any of the past JPL

or IUPAC evaluations. The exception is the study of

Salawitch et al. (2005), who relied solely on JPL kinetics.

Had they included this reaction, their values of Bry

VSLS

from the aircraft and balloon data would have been ~6 ppt

and ~8 ppt, respectively. The Salawitch et al. (2005) values

of Bry

VSLS are a little higher than those derived from the

DOAS measurements, but are generally consistent with

the picture presented in Figure 2-3.

Since the last Assessment, attention has also

focused on the interpretation of satellite observations of

BrO. As noted previously, observations of column BrO

from GOME reveal abundances of total column BrO,

throughout the global atmosphere, that are much larger

than expected based on simulations that assume supply of

atmospheric bromine from only CH3Br and halons (Figure

2-8, top panel). Measurements of column BrO from the

Scanning Imaging Absorption Spectrometer for Atmospheric Cartography (SCIAMACHY) satellite instrument

confirm the accuracy of the GOME observations of

column BrO (Sinnhuber et al., 2005; Sioris et al., 2006).

Recent scientific studies have focused on quantifying

whether the much higher than expected levels of BrO

reported by GOME reside in the stratosphere, the troposphere, or perhaps in both layers at levels important for

the photochemical environment. Evidence for and against

a global, background tropospheric level of BrO large

VERY SHORT-LIVED SUBSTANCES

2.37

Table 2-7. Contribution to stratospheric Bry

VSLS based on DOAS balloon BrO profiles.

Leon, Kiruna, Aire sur Kiruna, Teresina,

Spain Sweden l’Adour, France Sweden Brazil

1996 1999 2003 2003 2005

Bry from BrO (ppt)a 17.8 19.9 20.1 20.4 21.5

Age (yrs) 5 6 5 5 5

CH3Br (ppt)b 9.2 9.3 9.6 9.5 8.8

Halons (ppt)b 5.2 5.8 7.2 7.2 7.5

Bry

VSLS (ppt)c 3.4 4.8 3.3 3.7 5.2

a The 1σ precision of Bry estimated using the Langley method is about 0.7 ppt. The total uncertainty is 2.5 ppt, which represents the precision as well

as uncertainties in the BrO cross section (±8%) and the photochemical correction used to convert BrO to Bry (±8%). b Estimates of Bry

CH3Br+Halons are obtained from Montzka et al. (2003), assuming no tropospheric gradient of CH3Br. c The mean and standard deviation of the 5 estimates of Bry

VSLS are 4.1 ± 1 ppt. The standard deviation is consistent with the estimated precision in Bry.

Considering the overall uncertainty in Bry from BrO, and assuming that Bry

CH3Br+Halons is well known, leads to a range for BryVSLS of 4.1 ± 2.5 ppt,

which is the value for DOAS given in Table 2-8. From Dorf (2005).
enough to account for excess BrO revealed by GOME and

SCIAMACHY are described in Section 2.5.1.2. Clearly, a

better definition of tropospheric BrO would clarify the

stratospheric implications of the column BrO measurements, from both satellite and ground-based instruments.

Measurements of BrO profiles from SCIAMACHY

provide new insight on the global stratospheric distribution of BrO. However, retrievals of BrO have been conducted by two groups and the scientific interpretations

differ considerably. Retrievals by Sioris et al. (2006) show

large abundances in the tropical tropopause layer and lowermost stratosphere, consistent with large abundances of

stratospheric Bry, ~24 to 25 ppt, and Bry

VSLS = 8.4 ± 2 ppt

(Sioris et al., 2006) (Figure 2-8, middle and bottom

panels). Retrievals of SCIAMACHY BrO by Sinnhuber

et al. (2005) suggest stratospheric Bry = 18 ± 3 ppt,

implying a smaller value for Bry

VSLS of 3 ± 3 ppt (Figure

2-9). The BrO retrievals of Sinnhuber et al. (2005) show

a deficit in the tropical lower stratosphere compared with

calculations, particularly for model simulations that

include Bry from VSLS (Figure 2-9, panel labeled “18.5

km”). This deficit is contrary to the expectation of a significant perturbation to stratospheric Bry from VSLS.

Conversely, partial BrO columns retrieved by Sioris et al.

(2006) show large abundances in the tropical LMS, consistent with supply of Bry from VSLS (Figure 2-8, middle).

The retrievals of BrO in these two studies also offer

different views on tropospheric BrO. Comparison of the

Sinnhuber et al. (2005) stratospheric BrO profiles (based

on limb radiances) and total column BrO measurements

(based on nadir radiances) suggests the presence of 1 ±

0.5 ppt of tropospheric BrO (Figure 2-9, bottom). On the

other hand, the consistency of BrO columns integrated

above 12 km (limb) from SCIAMACHY and total column

BrO from GOME (nadir) shown by Sioris et al. (2006)

suggests a much smaller level of global tropospheric BrO

(Figure 2-8, middle). The reason for these alternate interpretations of SCIAMACHY radiances is due to differences

in the BrO retrievals computed by the two groups and does

not appear to be related to either the methods used to infer

VERY SHORT-LIVED SUBSTANCES

2.38

Figure 2-8. Top panel: Comparison of estimated stratospheric column BrO from GOME, October 1997,

assuming a uniform 1 ppt distribution of BrO in the troposphere (close to the upper limit of 0.9 reported by

Schofield et al., 2004) with stratospheric column BrO

from the AER 2-D model, for the WMO (2003) “best

guess” scenario Ab (labeled WMO Bry) and for simulations assuming Bry

VSLS of 4 and 8 ppt, respectively.

Plotted points show mean, standard deviation (thick

error bars) and extreme values (thin error bars) of

GOME data in 5°-wide latitude bins. The reasonably

good agreement between the 8 ppt curve and the data

suggest the BrO signal measured by GOME may be

due to the presence of higher amounts of BrO in both

the stratosphere and troposphere than is assumed in

most models. After Chance (1998) and Salawitch et

al. (2005). Middle panel: Comparison of total column

BrO measured by GOME (March, 2003) vs. latitude

and partial column BrO above various altitudes, based

on BrO profiles from SCIAMACHY limb radiances.

The comparison suggests a large component of the

GOME total column resides between ~12 and 21 km.

Bottom panel: Comparison of total column BrO above

15 km from SCIAMACHY with BrO partial columns

from the AER 2-D model, found for the WMO Bry and

for model simulations assuming Bry

VSLS of 4 and 8 ppt.

Model simulations from Salawitch et al. (2005);

SCIAMACHY BrO from Sioris et al. (2006).
VERY SHORT-LIVED SUBSTANCES

2.39

0

2

4

6

8

10

12

4

BrO VMR [ppt]

28.3 km

A

Zonal Mean 18−27 Sep 2002

SCIAMACHY

Model

Model + 3ppt

Model + 6ppt

0

2

4

6

8

10

12

14

BrO VMR [ppt]

25.0 km

B

0

2

4

6

8

10

12

14

BrO VMR [ppt]

21.7 km

C

0

2

4

6

8

10

12

14

BrO VMR [ppt]

18.5 km

D

0

2

4

6

8

10

12

14

BrO VMR [ppt]

15.2 km

E

−90 −60 −30 0 30 60 90

0

2

4

6

Limb above 15 km

Nadir

Latitude

BrO columns

[1013 molec / cm2

]

F

Figure 2-9. SCIAMACHY BrO observations. Panels A-E: Comparison between BrO volume mixing ratios (VMR)

retrieved from SCIAMACHY limb radiances and model calculations for different altitudes. SCIAMACHY error

bars are derived from the standard deviation of BrO measurements within each latitude bin. The modeled BrO

was calculated from an estimate of Bry derived from Michelson Interferometer for Passive Atmospheric Sounding

(MIPAS) measurements of CFC-11 and the Wamsley et al. (1998) relation between Bry and CFC-11 (scaled to the

time of observation) and a photochemical model estimate of BrO/Bry. Three model runs are shown: Bry from only

CH3Br and halons (i.e., WMO Bry) (red line), and using Bry

VSLS of 3 ppt of Bry (green line) and 6 ppt (blue line).

Panel F: Total BrO column from SCIAMACHY nadir measurements compared with the integrated vertical BrO

column above 15 km from SCIAMACHY limb observations. The difference between the integrated limb and the

nadir column implies the presence of a significant amount of BrO below 15 km, corresponding to an average BrO

mixing ratio of about 1.0 ± 0.5 ppt between the surface and 15 km. Adapted from Sinnhuber et al. (2005).
Bry from BrO, or to assumptions regarding Bry

CH3Br+Halons.

The SCIAMACHY BrO retrievals are sensitive to the

determination of the tangent height of the limb radiances

and methods used to account for spectral residuals (Sioris

et al., 2006). Efforts are underway to compare balloon

profiles of BrO with the SCIAMACHY retrievals (Dorf et

al., 2006a). However, much work remains to understand

the differences between the various retrievals of BrO and

to reduce the uncertainty in the global distribution of BrO

retrieved from SCIAMACHY radiances, which is perhaps

best reflected by the contrasting results of Sinnhuber et

al. (2005) and Sioris et al. (2006).

Measurements of BrO from the Microwave Limb

Sounder (MLS) have also been used to quantify Bry

VSLS

(Livesey et al., 2006). These observations, shown in Figure

2-10, imply a value for Bry of 18.6 ± 5.5 ppt for the upper

stratosphere (~30 to 42 km), between 55°S and 55°N, from

September 2004 to August 2005. The uncertainty reflects

the accuracy of the MLS measurements of BrO and possible errors in kinetic parameters used to infer Bry from

BrO. Livesey et al. (2006) inferred Bry from BrO using

two modeling approaches, but place greater confidence in

the approach that uses a model constrained by MLS measurements of O3 and nitrous oxide (N2O) (PSS results in

Figure 2-10). Their value of Bry

VSLS is 3.0 ± 5.5 ppt.

Observations of BrO in the polar vortices provide

an opportunity to examine the bromine budget in an environment where all of the organic precursors have decomposed, and where the low levels of reactive nitrogen oxides

(NO + NO2 = NOx) imply that nearly all of Bry is contained

in the two species BrO and bromine chloride (BrCl). These

observations, discussed in Chapter 4, also support a significant, non-zero value for Bry

VSLS. As discussed in

Chapter 4, the 6 to 8 ppt enhancement of BrOx (BrO+BrCl)

in the Arctic vortex (Frieler et al., 2006), presumably from

VSLS, relative to standard models (see Figure 4-16 in

Chapter 4) may account for the discrepancy between measured and modeled chemical loss rates for Arctic ozone that

was noted in the last Assessment.

Table 2-8 shows estimates of Bry

VSLS from nine

studies. The central value from each study is given, as

well as a range. For most studies, the uncertainty of

Bry

VSLS was explicitly calculated while, for a few, it has

been estimated by this Assessment based on information

in each paper (see Table 2-8). The range was estimated

using different methods by the various studies; the original papers should be consulted for details. Estimates of

Bry

VSLS from ground-based studies are included. The

Sinnhuber et al. (2002) study, discussed in the last

Assessment, has been described above. Schofield et al.

(2004, 2006) examined direct and scattered sunlight to

VERY SHORT-LIVED SUBSTANCES

2.40

Figure 2-10. Measurement of BrO from Aura

MLS versus pressure, for

limb scans covering 55°S

to 55°N, September 2004

to August 2005, and

values of Bry inferred from

MLS BrO using estimates

of BrO/Bry from the

SLIMCAT model (blue)

and from a photochemical

steady state (PSS) model

(red) constrained by MLS

BrO, O3, and H2O, and

values of NOy and CH4

derived from MLS N2O.

The value of Bry due to

CH3Br and halons during

the time of measurement

is shown by the dotted

line. Error bars reflect the accuracy of MLS BrO and uncertainties in the model estimate of Bry. The analysis

using the PSS simulation was considered to be a more reliable estimate for stratospheric Bry because this

model is constrained by measured O3 and exhibits a profile for NO2 that agrees well with Halogen Occultation

Experiment (HALOE) satellite measurements. The stratospheric bromine loading is estimated to be 18.6 ± 5.5

ppt, implying Bry

VSLS = 3.0 ± 5.5 ppt. From Livesey et al. (2006).

5 10 15 20 25 30

Bromine abundance / ppt

10.0

6.8

4.6

3.2

2.2

Pressure / hPa

MLS BrO

MLS Br (SLIMCAT)

MLS Br

(PSS)

 Br from CH Br and halons

32

34

36

38

40

42

Approximate altitude / km
retrieve BrO profiles over Lauder, New Zealand (45°S),

and Arrival Heights, Antarctica (78°S). They concluded that stratospheric column BrO was consistent with

Bry

VSLS of 6 ± 3 ppt at both locations. 

Estimates of the Bry

VSLS given in Table 2-8 range

from lows of 3.0 ± 5.5 ppt (MLS data: Livesey et al., 2006)

and 3 ± 3 ppt (SCIAMACHY retrievals: Sinnhuber et al.,

2005) to highs of 4 to 10 ppt (aircraft and balloon data:

Salawitch et al., 2005) and 8.4 ± 2 ppt (SCIAMACHY

retrievals: Sioris et al., 2006). Dorf (2005) report a value

of 4.1 ± 2.5 ppt from many years of balloon observations.

Sheode et al. (2006) report a value of 3.5 ± 4 ppt, based on

analysis of two years of SCIAMACHY data. This study

is an update to Sinnhuber et al. (2005), who considered

SCIAMACHY data over a 10-day period; the slightly

larger uncertainty given by Sheode et al. (2006) is due to

increased atmospheric variability over the longer time

period.

The mean of the central values of Bry

VSLS given in

Table 2-8 is 5 ppt. This is the “ensemble” value reported

at the bottom of the table. It is difficult to ascribe a true

uncertainty to this ensemble value, given that uncertainties in Bry

VSLS were determined using different methods,

and in some cases were estimated by this Assessment. As

a result, we place greater confidence in the range of central values in the Table (3 to 8 ppt) as an estimate of the

uncertainty in Bry

VSLS. While three of the studies could

not rule out a zero contribution from VSLS, the majority

of studies pointed to a positive value for Bry

VSLS large

enough to affect ozone photochemistry in the LMS

(Section 2.6).

2.5.2.2 CHLORINE

Measurements of HCl in the subtropics typically

exceed 50 ppt (Marcy et al., 2004). As noted above, these

VERY SHORT-LIVED SUBSTANCES

2.41

Table 2-8. Contribution to stratospheric Bry

VSLS.

Bry

VSLS BryVSLS

Data Source Central Value (ppt) Range (ppt) Reference

Ground-based BrO 5 1-9 a Sinnhuber et al. (2002)

11 sites, 78°S-79°N

Ground-based BrO 6 3-9 Schofield et al. (2004)

Lauder, New Zealand, 45°S

Ground-based BrO 6 3-9 Schofield et al. (2006)

Arrival Heights, Antarctica, 78°S

DOAS Balloon BrO Profiles 4.1 1.6-6.6 Dorf, 2005

5°S-68°N, 0-35 km

Aircraft & Balloon BrO Profiles, 7 4-10 a Salawitch et al. (2005)

22°S-35°N, 17-32 km and

GOME Satellite Column BrO,

60°S-60°N

SCIAMACHY Satellite BrO Profiles 8.4 6.4-10.4 Sioris et al. (2006)

80°S-80°N, 15-28 km

SCIAMACHY Satellite BrO Profiles 3 0-6 Sinnhuber et al. (2005)

80°S-80°N, 15-28 km 

SCIAMACHY Satellite BrO Profiles 3.5 0-7.5 Sheode et al. (2006)

80°S-80°N, 15-28 km 

MLS Satellite BrO Profiles 3.0 0-8.5 Livesey et al. (2006)

55°S-55°N, 30-42 km

Ensemble 5 (3-8) b

DOAS, Differential Optical Absorption Spectroscopy; GOME, Global Ozone Monitoring Experiment; SCIAMACHY, Scanning Imaging Absorption

Spectrometer for Atmospheric Cartography; MLS, Microwave Limb Sounder. a Range estimated by this Assessment, based on the uncertainty in stratospheric Bry inferred from BrO that was stated in the reference. b Average and range of the central values of the 9 estimates of Bry

VSLS.
data have been interpreted as being evidence for supply

of HCl from the stratosphere, rather than being due to

the decomposition of VSL SGs. There have been no

studies, since the last Assessment, quantifying the influence of VSL SGs on measured HCl in the tropical TTL

or LMS. Such a study would require accurate and precise measurement of HCl and a suite of organic chlorine

species, including long-lived SGs and VSL SGs, plus

either detailed model analysis and/or observations of

dynamical tracers to be able to isolate the effects on chlorine of transport from above (where the longer lived SGs

decompose) and from below (where VSL SGs potentially

contribute).

Another avenue for assessing the role of VSL compounds on stratospheric chlorine, which has received some

attention since the last Assessment, is quantification of

stratospheric chlorine either by measurements of a suite

of compounds in the middle stratosphere or by measurement of HCl in the upper stratosphere. Nassar et al. (2006)

quantified total stratospheric chlorine (ClTOT, the sum of

simultaneous measurements of organic and inorganic

species; see Figure 1-12 of Chapter 1) based primarily on

data from the Atmospheric Chemistry Experiment Fourier

Transform Spectrometer (ACE-FTS). They report a value

of ClTOT of 3.65 ppb over the period February 2004 to

January 2005, for both Southern and Northern Hemisphere

midlatitudes. The 1σ precision of ClTOT is 0.09 ppb and

the estimated accuracy is 0.13 ppb. The estimated contribution of just long-lived species to Cly during this time

period is 3.41 ppb (this estimate is based on a mean age

for stratospheric air of 5 years and an age of air spectrum

width of 3 years, applied to the WMO 2002 baseline scenario Ab (see Chapters 1 and 3 of this Assessment) for the

12 long-lived chlorine species). The fact that the ACEFTS value of ClTOT exceeds this value of Cly could be

indicative of a contribution from VSLS. On the other hand,

the combined uncertainty of the ACE-FTS value of ClTOT

(accuracy + precision) nearly overlaps with the expected

value of Cly from just long-lived species.

Time series of HCl near 55 km from the ATMOS

(Atmospheric Trace Molecule Spectroscopy experiment),

HALOE (Halogen Occultation Experiment), MLS, and

ACE-FTS instruments are shown in Figure 1-12 of Chapter

1, where it is discussed in terms of trends of Cly in the stratosphere. In the Figure these measurements are compared

with modeled stratospheric abundances of HCl due to

long-lived chlorinated gases alone (dashed line), and

enhanced by the addition of 100 ppt of chlorine to represent a possible contribution from VSLS and phosgene

(solid lines). Because of the level of uncertainty in the

respective satellite measurements (note the error bars in

Figure 1-12, which represent 2-sigma uncertainty), it is

difficult to discriminate between the model runs with and

without enhanced chlorine. The model run based on longlived source gases alone appears closer to the HALOE

measurements of HCl, which are lower than those obtained

by the other instruments (although within their reported

error limits). The MLS, ATMOS, and ACE-FTS measurements are suggestive of a contribution of chlorine from

VSLS. The contribution of VSLS to stratospheric Cly

cannot be quantified definitively given the current level of

accuracy of the satellite HCl measurements.

2.5.2.3 IODINE

Solomon et al. (1994) suggested that if stratospheric

inorganic iodine (Iy) levels were close to 1.0 ppt, catalytic

cycles involving IO + BrO and IO + ClO could be responsible for a significant component of midlatitude ozone

trends. That suggestion has led to considerable effort to

better understand the photochemistry of inorganic iodine

species (e.g., Bösch et al., 2003, and references therein)

and to define the stratospheric iodine budget.

Wennberg et al. (1997), based on analysis of

ground-based spectra, reported an upper limit for Iy of 0.2

ppt. Pundt et al. (1998) reported a similar upper limit,

based on analysis of balloonborne spectra. Both studies

were summarized in the last Assessment. Since the last

Assessment, several additional studies of the Iy budget

have appeared. Bösch et al. (2003), based on balloonborne

solar occultation spectra in the UV/visible region from

flights at middle and high latitudes, report an upper limit

for Iy of 0.1 ppt. Recent measurements of IO in the tropics,

using the same technique, also show very low levels of IO

(Butz, 2006) (Figure 2-11). These observations suggest

that either much less iodine enters the stratosphere than

expected based on levels of CH3I near the surface, or that

stratospheric iodine resides either in gaseous species

besides IO or perhaps in particulate form (see Section

2.3.5). As noted above, the tropospheric lifetime of the

dominant iodine VSL SGs is much smaller (i.e., ~5 days)

than the tropospheric lifetime of many brominated VSLS

(>100 days). This could account for SGI of bromine and

little or no SGI of iodine. Furthermore, observations of

self-nucleation of iodine oxides (Hoffmann et al., 2001)

and iodine uptake on tropospheric aerosols (Murphy and

Thomson, 2000) could account for inefficient transfer of

Iy to the stratosphere via PGI. 

The only evidence of significant levels of stratospheric iodine is provided by ground-based observations

of spectra recorded at 79°N that show, on some days, signs

of elevated column IO (Wittrock et al., 2000). The SZA

variation of slant column density of IO was used to suggest that the elevated IO lies in the stratosphere, with an

VERY SHORT-LIVED SUBSTANCES

2.42
abundance of Iy as high as 0.8 ppt. The considerable variation in Iy for different days is difficult to reconcile in terms

of a large stratospheric contribution, particularly for highlatitude air masses that should lack any residual CH3I.

Higher levels of IO were seen in February 1997 compared

with March 1997, associated with lower levels of NOx in

February. Similar observations of the SZA variation of

slant column IO in the Antarctic have been used to argue

for a burden located primarily in the troposphere (Frieß et

al., 2001). The observation of highly elevated IO on some

days by Wittrock et al. (2000) is presently the only piece of

observational evidence supporting a large source of Iy from

VSLS. It should be recalled, however, that the ozone

destruction efficiency of iodine is large compared with that

of both bromine and chlorine, and IO abundances in the

stratosphere of several tenths of 1 ppt could be potentially

significant (Figure 2-2).

2.6 POTENTIAL IMPACT OF VSLS ON OZONE

In WMO (2003) the focus of this topic was on (1)

the theoretical basis for estimating the change in column

ozone from emissions of organic VSL SGs, and (2) the theoretical basis for estimating Ozone Depletion Potentials

(ODPs), including the differences (Wuebbles and Ko, 1999;

Wuebbles et al., 2001) from the standard ODP definition,

and initial analyses of location-dependent ODPs for one

VSL SG, n-propyl bromide. WMO (2003) gave a comprehensive evaluation of the theoretical basis and procedures

for estimating ozone column effects and ODPs of VSL SGs;

the discussion here is intended to update that evaluation.

2.6.1 Effects of Halogenated VSLS on

Column Ozone

Halogens supplied by the decomposition of VSLS

have the potential to significantly alter the photochemistry

of ozone in the upper troposphere (UT) and lowermost

stratosphere (LMS). Within models that consider supply

of bromine solely from CH3Br plus halons, pure oddhydrogen radical (HOx) photochemistry is the dominant

loss process for odd oxygen in the LMS (Salawitch et al.,

2005; see Figure 2-2). As discussed in Section 2.5, Bry

VSLS

in the LMS, maintained by the SGI and PGI pathways, can

potentially cause significant local perturbation to the Bry

budget. Inorganic tropospheric bromine may also be

transported in the LMS and contribute to additional

Bry

VSLS (as discussed in Section 2.5). Enhanced levels of

bromine lead to greater efficiency for ozone loss by the

BrO + ClO catalytic cycle, particularly during times of

high aerosol loading following major volcanic eruptions

(Salawitch et al., 2005; Sinnhuber et al., 2006). Because

of this cycle, Bry

VSLS enhances ozone depletion due to

CFCs and other chlorocarbons. If Bry

VSLS is 5 ppt, as suggested by the assessment presented in Section 2.5, and

if the majority of this bromine is present in the LMS, then

ozone loss by the BrO + HO2 cycle becomes greatly

enhanced (Salawitch et al., 2005). Indeed, at Bry

VSLS

between 4 to 8 ppt, ozone loss by bromine cycles throughVERY SHORT-LIVED SUBSTANCES

2.43

Figure 2-11. Comparison of

measured and expected IO

absorption by balloonborne

UV/visible solar occultation spectroscopy at high and low latitudes

(black line: measured IO +

residual absorption; red line:

inferred IO absorption: blue line:

the expected IO absorption

assuming [IO] = 0.5 ppt for the 0

to 20 km altitude range). Panels

(a) and (b) are for observations

within the lowermost stratosphere, and panels (c) and (d) for

observations within the tropical

tropopause layer. TH = tangent

height. Panel (a) is adopted from

Bösch et al. (2003), and panels

(b) to (d) are from Butz (2006).
out the lowermost stratosphere can approach ozone loss

by HOx cycles alone (Figure 2-2). 

In the last Assessment, a value of 45 was given for

the relative effectiveness of stratospheric ozone depletion

by bromine; the so-called α parameter (Section 1.4.4,

WMO, 2003). This value has been reassessed in Chapter

8 of this Assessment (Section 8.2.2), and the recommendation made that the value of α be increased to 60 (65 for

the Antarctic stratosphere). As a consequence, the semiempirically determined values of ODPs for all brominated

source gases have now increased. Some models that also

include VSLS in their calculations suggest even larger

values for α. For example, Sinnhuber et al. (2006), using

the University of Bremen two-dimensional model

(Chipperfield and Feng, 2003), derive values for α of 66

without Bry

VSLS and 69 including BryVSLS, suggesting an

impact on α from halogenated VSLS. Updated versions

of the Atmospheric and Environmental Research, Inc.,

(AER) (U.S.) and University of Illinois at UrbanaChampaign (UIUC) (U.S.) models that include Bry

VSLS

from gases discussed earlier in this chapter (i.e., CH2Cl2,

CHCl3, C2HCl3, C2Cl4, CH2BrCl, CH2Br2, CHBr2Cl,

CHBrCl2, and CHBr3) both derive an α of 61. More work

is needed to assess the importance of Bry

VSLS on α.

There was no attempt in WMO (2003) to quantify

the effects of halogenated VSLS on past trends, or on

future changes in stratospheric ozone or ozone columns.

Since then, there has been increasing awareness that VSLS

have a potentially important effect on chlorine and

bromine in the lower stratosphere. As discussed in Section

2.5 of this report, halogenated VSLS appear to add about

5 ppt, on average, of reactive bromine (Bry) to the lower

stratosphere. Although significant uncertainties remain,

analyses with the AER 2-D chemical transport model in

Salawitch et al. (2005) suggest that this constant amount

of Bry added to the lower stratosphere has a noticeable

impact on past trends in column ozone, as shown in Figure

2-12. Similar results were found using the University of

Bremen 2-D model (Sinnhuber et al., 2006). Indeed, the

inclusion of additional Bry

VSLS may lead to better agreement between modeled and observed stratospheric ozone

trends, particularly during times of perturbed aerosol

loading, especially in the Northern Hemisphere (Salawitch

et al., 2005; Sinnhuber et al., 2006). (Similar findings are

discussed in Chapter 3, Section 3.4.3.2, and shown in

Figure 3-25.) Both of these simulations assumed that

Bry

VSLS was constant over time, due to primarily natural

sources. Increased ozone depletion due to Bry

VSLS was

caused by the ability of bromine to enhance ozone loss by

ClO derived from CFCs, as outlined above (Chapter 3 also

discusses impacts of Bry

VSLS on ozone trends). The effect

of anthropogenic emissions of Bry

VSLS or possible longterm changes in Bry

VSLS due to climate change or natural

variability, have yet to be studied (see Section 2.7).

Several recent studies (e.g., von Glasow et al., 2004;

Yang et al., 2005) suggest that inorganic bromine from

halogenated VSLS could lead to reductions in tropospheric

ozone. The effect on tropospheric ozone should largely

be regional in nature, but may have implications on globalscale ozone concentrations. Reductions in ozone occur

due to a direct effect (increased loss by the BrO + HO2

cycle) and an indirect effect (reduced production due to

decreased levels of NO) that is ultimately driven by

increased efficiency of the BrONO2 hydrolysis sink of NOx

(Lary, 2005). More accurate evaluations will be dependent

on improved knowledge about the bromine budget in the

troposphere. If anthropogenic production of brominated

VSLS becomes important in the future, the effect of inorganic product gas on tropospheric ozone will have to be

evaluated to quantify the impact of these compounds on

total column ozone.

VERY SHORT-LIVED SUBSTANCES

2.44

Figure 2-12. Calculated change in column ozone

relative to 1980 levels using the AER 2-D model for

stratospheric Bry

VSLS of 0 (red), 4 (green), and 8 (blue)

ppt for 35°N-60°N (top) and for 35°S-60°S (bottom),

compared with observed trends in total ozone (black

line). Based on Salawitch et al. (2005).
2.6.2 ODPs for Halogenated VSLS

There have been only a few new reported studies of

the ODPs for VSL SGs since the last Assessment. The

additional model estimates of SGI and PGI fluxes provide

an opportunity to update the estimates for ODP using

Equation 2.15 from Ko and Poulet et al. (2003). Using the

results discussed in Section 2.5.1 as a guideline, it is reasonable to assume that for a source gas with a local tropospheric lifetime of about 25 days and uniform landmass

emissions, 1% of the emitted flux enters the stratosphere

via the SGI and PGI pathways. We make the additional

assumption that the residence time of the Bry introduced

by either pathway is similar to that of the Bry produced by

long-lived SGs. The order of magnitude estimate for the

ODP of a species with molecular weight comparable to

that of CFC-11, and a local lifetime of about 25 days, can

be determined to be c. 0.003 if it contains 1 chlorine atom,

and c. 0.18 if it contains 1 bromine atom (assuming an

alpha factor of 60 for bromine). A simple scaling for iodinated compounds is not possible because of their shorter

lifetimes and the uncertainty of the alpha factor for iodine.

The resulting ODP for any given compound with the above

characteristics would be linearly dependent on the number

of chlorine or bromine atoms.

Using the UIUC zonally averaged 2-D model, Li et

al. (2006) have re-evaluated the ODPs for surface emissions of trifluoroiodomethane (CF3I), a compound that is

proposed as a replacement for bromotrifluoromethane

(halon-1301, CBrF3) in firefighting and inhibition applications. The model used an updated representation of

iodine chemistry relative to an earlier study of the ODP of

CF3I (Solomon et al., 1994). They found an ODP of 0.013

for emissions evenly distributed over the Northern

Hemisphere (compared with <0.008 from Solomon et al.

(1994)), 0.011 for emissions distributed over 30°N to

60°N, and 0.018 for emissions distributed over 5°S to

30°N. Li et al. (2006) also consider the effects of potentially using CF3I in aircraft. The resulting ODPs depend

heavily on the altitude and latitude of the emissions. For

example, emissions peaking at 6-9 km were calculated to

give an ODP of 0.084 for emissions occurring at 30°N40°N and 0.25 for emissions occurring at 0-10°N. Another

proposed replacement for halon-1301 in aircraft engine

fire applications, phosphorus tribromide (PBr3), has not

been evaluated, but should also have non-zero ODPs.

The ODP of n-propyl bromide was considered at

length in WMO (2003), but there have been no new evaluations since that date.

2.7 POTENTIAL FUTURE CONSIDERATIONS

The impact of halogenated VSLS on stratospheric

halogen loading and ozone depletion is, as discussed in

Sections 2.5.1 and 2.6.1, currently thought to be relatively small (but significant) compared with that from

the long-lived halocarbons. The present-day impact

arises principally from the brominated VSLS, whereas

the contributions from chlorinated and iodinated gases

(for instance, to EESC) are small.

This situation could potentially change in the future.

Small anthropogenic emissions of halogenated VSLS

already exist (e.g., n-C3H7Br, CHCl3, CH2Cl2, C2Cl4, etc.;

see Section 2.2.4). These could increase in the future

(although several of these have evidently decreased in the

last decade, as previously discussed), or new products

could enter commercial use (e.g., CF3I; Sections 2.2.4.3

and 2.6.2). However, as noted previously, the majority of

halogenated VSLS have partly or wholly natural origins,

and therefore contribute to the natural background halogen

loading of the stratosphere. This is set against a future

trend of declining concentrations of long-lived anthropogenic halogenated gases, and hence a potentially larger

relative impact from halogenated VSLS.

The exact sources and intensities of natural VSLS

emissions are largely linked to biogenic activity in both

the marine and terrestrial biospheres (there is also evidence of some abiotic marine emissions; see Section

2.2.3.2). It is conceivable that biogenic emissions could

change in response to, for example, changing global

temperatures, atmospheric CO2, rainfall, land use, vegetative cover, and oceanic productivity. Evidence for this is

presently sparse (Section 2.2). In one study (Butler et al.,

2006), positive correlations between open ocean flux rates

of CHBr3 and CH3I with sea surface temperature (SST)

were reported. The authors noted, however, that SST

might simply be a proxy for some other driver (e.g., incident radiation, latitude, etc.). They further noted that while

SST (and also wind speed) largely dictate sea-air flux rates,

other significant parameters, such as nutrient supply and

stratification of surface waters, are also affected by SST

and wind speed, and may have important modifying effects

(both positive and negative) on overall emission rates.

Another study, reported in Section 2.2.2, showed that

measurements of organo-halogen emissions from rice paddies indicate an apparent positive correlation between temperature and CH3I emissions (Redeker and Cicerone,

2004). On the other hand, firn records of the atmospheric

abundance of several halogenated VSLS suggest minimal

VERY SHORT-LIVED SUBSTANCES

2.45
long-term changes throughout the second half of the 20th

century, other than small changes in the Northern Hemisphere atmosphere for a few gases with evident anthropogenic sources (Worton et al., 2006; Sturges et al., 2001).

It is noted that inorganic VSLS in the boundary

layer, notably BrO from halogen activation reactions

(Sections 2.2.5 and 2.5.1.2), might be sensitive to climatic

changes. Hollwedel et al. (2004), for instance, suggest a

link between observed increases in springtime BrO column

density with increased annual sea ice coverage, but we do

not consider it likely that boundary layer VSLS from high

latitudes will be transported to the stratosphere. There

might, however, be a potential link between inorganic

halogen released from sea salt in the free troposphere and

wind speed-dependent generation of marine aerosol (see

also Section 2.2.5).

It is significant that several of the limited studies to

date point toward maximum flux rates of halogenated

VSLS in the tropics (Section 2.2). This coincides with the

region where strong convection can most effectively and

quickly transport VSLS to the tropopause region (Section

2.4). Changes in circulation and meteorology in the tropics

might also exert a powerful modifying influence on the

contribution of VSLS to stratospheric halogen loading.

These effects could be either positive or negative. For

example, a stronger future Brewer-Dobson circulation

(as suggested by Butchart et al., 2006) could increase the

rate of vertical transport of VSLS from the TTL into the

stratosphere, whereas enhanced precipitation could decrease the efficiency of the PGI route (Section 2.3.4.1).

Changes in the latitudes of the subtropical jets, or equivalently the extent of the tropospheric Hadley circulation

(Fu et al., 2006), might be associated with changes in

quasi-horizontal transport of VSLS from the TTL into

the lowermost stratosphere.

Future changes in the chemical processing of VSLS

and their product gases could also influence their atmospheric concentrations (Section 2.3), for example, through

changes in OH abundance.

Our understanding of these many processes is

presently too poor to allow any quantitative prediction of

possible future halogenated VSLS trends.

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2.57
CHAPTER 3

Global Ozone: Past and Present

Contributors:

G. Bodeker

P. Demoulin

W. Feng 

L. Flynn

S. Frith

S. Guillas

M. Ilyas

B. Knudsen

B. Liley

R. McPeters

A.J. Miller

R. Müller

S. Oltmans

Y. Orsolini

I. Petropavlovskikh

W. Steinbrecht

H. Struthers

D. Tarasick

Y. Terao

Lead Authors:

M.P. Chipperfield 

V.E. Fioletov 

Coauthors:

B. Bregman

J. Burrows

B.J. Connor

J.D. Haigh

N.R.P. Harris

A. Hauchecorne

L.L. Hood

S.R. Kawa

J.W. Krzyscin  ́

J.A. Logan

N.J. Muthama

L. Polvani

W.J. Randel

T. Sasaki

J. Stähelin

R.S. Stolarski

L.W. Thomason

J.M. Zawodny

Final Release: February 2007

From Scientific Assessment of Ozone Depletion: 2006

CHAPTER 3

GLOBAL OZONE: PAST AND PRESENT

Contents

SCIENTIFIC SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.1

3.1 INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.5

3.2 UPDATE OF OZONE CHANGES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.5

3.2.1 Statistical Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.5

3.2.2 Changes in Total Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.7

3.2.3 Changes in the Vertical Distribution of Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.8

3.2.3.1 Upper Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.10

3.2.3.2 Lower Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.13

3.2.3.3 Consistency of Column and Profile Trends . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.17

3.2.4 Tropospheric Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.18

3.3 SUMMARY OF OTHER OBSERVATIONS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.19

3.3.1 Stratospheric Aerosol and Its Precursors . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.19

3.3.2 Nitrogen Dioxide (NO2) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.20

3.4 UNDERSTANDING AND INTERPRETATION OF OZONE CHANGES . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.21

3.4.1 Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.21

3.4.2 Dynamical Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.22

3.4.3 Chemical Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.25

3.4.3.1 Update of Relevant Kinetics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.25

3.4.3.2 Lower Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.26

3.4.3.3 Upper Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.30

3.4.4 Solar Cycle Variations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.30

3.4.5 Assessment Model Simulations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.32

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.41

APPENDIX 3A: DESCRIPTION OF THE OZONE DATASETS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.55

3A.1 Column Ozone Data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.55

3A.2 Ozone Profile Measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.57

GLOBAL OZONE

3.1

SCIENTIFIC SUMMARY

Total Column Ozone

• Global mean total column ozone values for 2002-2005 were approximately 3.5% below 1964-1980 average

values. The 2002-2005 values are similar to the 1998-2001 values and this indicates that ozone is no longer

decreasing. These changes are evident in all available global datasets, although differences of up to 1% between

annual averages exist between some individual sets.

• Averaged for the period 2002-2005, total column ozone for the Northern Hemisphere (NH) and Southern

Hemisphere (SH) midlatitudes (35°-60°) are about 3% and 5.5%, respectively, below their 1964-1980 average

values and are similar to their 1998-2001 values. However, as noted in previous Assessments, the behavior of

column ozone in the two hemispheres during the 1990s was different. The NH shows a minimum around 1993, followed by an increase. The SH shows an ongoing decrease through the late 1990s, followed by the recent leveling

off.

• There are seasonal differences in ozone changes over midlatitudes between the NH and the SH. Changes

since the pre-1980 period over northern midlatitudes (35°N-60°N) are larger in spring, while those over southern

midlatitudes (35°S-60°S) are nearly the same throughout the year.

• Total column ozone over the tropics (25°S-25°N) remains essentially unchanged. Total ozone trends in this

region for the period 1980-2004 are not statistically significant. These findings are consistent with the findings of

the previous Assessments.

Vertical Ozone Distribution

• Upper stratospheric ozone declined during 1979-1995, but has been relatively constant during the last decade.

Measurements from the Stratospheric Aerosol and Gas Experiment (SAGE I+II) and Solar Backscatter Ultraviolet

(SBUV(/2)) satellite instruments show significant declines through 1995 when averaged over 60°N-60°S and altitudes of 35 to 50 kilometers (km). The net ozone decrease was ~10-15% over midlatitudes; smaller but significant

changes occurred over the tropics. Available independent Umkehr, lidar, and microwave ozone measurements confirm these findings.

• Lower stratospheric ozone declined over the period 1979-1995, but has been relatively constant with significant variability over the last decade. At midlatitudes of both the Northern and Southern Hemispheres, measurements by SAGE I+II and SBUV(/2) showed declines of up to 10% by 1995 between 20 and 25 km altitude. These

decreases did not continue in the last decade. 

• In the lowermost stratosphere, between 12 and 15 km in the Northern Hemisphere, a strong decrease in

ozone was observed from ozonesonde data between 1979 and 1995, followed by an overall increase from 1996

to 2004, leading to no net long-term decrease at this level. These changes in the lowermost stratosphere have a

substantial influence on the column. The Southern Hemisphere midlatitude data do not show a similar increase

since 1995 at these altitudes.

• Significant ozone decreases of ~3% between the tropopause and 25 km are found in the SAGE satellite measurements between 1979 and 2004 at 25°S-25°N. Since no change is found in total ozone over the tropics, this

could be explained by significant increases in tropospheric ozone in the tropics. While regional increases in tropical

tropospheric ozone have been seen, not all tropical regions or datasets show an increase.
GLOBAL OZONE

3.2

Understanding Past Changes in Ozone

• There has been no change in our basic understanding that, analyzed over all latitudes and seasons, halogen

increases have been the principal driver of ozone depletion over the past few decades. There is good overall

agreement between observed long-term changes in extrapolar ozone and model simulations that include the effects

of increasing halogens. The models generally reproduce the observed ozone changes as a function of altitude, latitude, and season, confirming our understanding that halogen changes are the main driver of global ozone changes.

This link is supported by the statistical fit of globally averaged ozone observations with Equivalent Effective

Stratospheric Chlorine (EESC).

• Empirical and model studies have shown that changes in tropospheric and stratospheric dynamics have been

partially responsible for the observed NH midlatitude winter ozone decline from 1979 to the mid-1990s and

ozone increase thereafter. Whether this is due to dynamical variability or results from a long-term trend in stratospheric circulation is not yet clear. Estimates of these dynamical effects on long-term trends range from ~20% up to

50% for the winter period.

• The impact of aerosols (dynamical and chemical) on midlatitude ozone was greatest in the early 1990s after

the eruption of Mt. Pinatubo in 1991. The observed decrease in NH column ozone in 1993 agrees with chemicaldynamical models that include these effects. An outstanding issue is the absence of an effect of the Mt. Pinatubo

eruption on the observed SH ozone column, which contrasts with model predictions.

• Several independent modeling studies covering periods in the 1990s confirm that dilution of ozone-depleted

polar air makes a substantial contribution to midlatitude ozone depletion, especially in the Southern

Hemisphere (due to the much larger polar ozone loss there). Long-term annually averaged model-based estimates indicate that dilution by polar air contributes about one-third of the midlatitude ozone depletion in the Northern

Hemisphere, with large interannual variation; in the Southern Hemisphere midlatitudes, the contribution is estimated to be about one-half.

• The inclusion of additional inorganic bromine (Bry) from very short-lived substances (VSLS) in models leads

to larger ozone destruction at midlatitudes, compared with studies including only long-lived bromine source

gases. The enhanced ozone loss occurs in the lower stratosphere via interactions of this bromine with anthropogenic chlorine. Midlatitude ozone loss is most enhanced during periods of high aerosol loading. Ozone loss

through cycles involving bromine and odd-hydrogen (HOx) is also enhanced at midlatitudes under all conditions.

The impact on long-term midlatitude ozone trends (1980-2004), assuming constant VSLS Bry, is calculated to be

small because aerosol loading was low at the start and end of this time period.

• The profile shape of upper stratospheric ozone trends from 1980-2004 is generally consistent with our understanding of gas-phase chlorine chemistry as the cause, modulated by changes in temperature and other gases

such as methane (CH4). However, global dynamical-chemical models have not demonstrated that they can simultaneously reproduce realistic trends in all relevant parameters, although observations over the full time period are

limited. Non-interactive models obtain ozone change that peak at about 14% for 1980-2004 (in altitude coordinates), consistent with SAGE observations. Interactive models, which calculate their own temperatures, obtain

ozone trends that are nearly half in magnitude.

• Both two-dimensional (2-D) and three-dimensional (3-D) models perform better in reproducing observed

past changes in the NH than in the SH. Consistent with results presented in the previous Assessment, 2-D models

show large model-model differences in the SH due to different treatments of the Antarctic ozone loss. Off-line 3-D

chemical transport models (CTMs) are now also routinely available to study past changes. These models are inherently better in simulating the polar regions and this leads to relatively small model-model differences, especially in

the SH. CTMs also clearly perform better at reproducing long-term changes in the NH than in the SH. The ongoing

disagreement between model-observation comparisons in the NH versus the SH indicates that we do not yet have a

full understanding of the processes controlling ozone changes at midlatitudes.
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3.3

Solar Cycle

• Identification of the solar cycle signal in observed ozone has been improved due to the absence of major volcanic eruptions over the past 15 years. The deduced solar cycle variation in column ozone has a mean amplitude

of 2-3% (from minimum to maximum) in low to midlatitudes from the extended data series.

• Model estimates and measurements suggest that the amplitude of the solar cycle in ozone concentration is

less than 4% throughout the stratosphere, although there are apparent differences between the models and

observations at some altitudes. The vertical structure of the ozone solar cycle variation is subject to observational

uncertainties (e.g., short record lengths, instrument intercalibration problems) that make it difficult to test models

critically.

3.1 INTRODUCTION

The abundance of stratospheric ozone is determined

by a combination of chemical and dynamical processes.

The importance of these processes changes with location

(particularly altitude and latitude) and with time. These

processes also act on different time scales. For example,

the influence of the solar cycle on total column ozone

varies over its 11-year cycle, the quasi-biennial oscillation

(QBO) over an approximately 27-month cycle, and strong

volcanic eruptions have an effect lasting a few years.

The long-term changes in global ozone were last

reviewed in Chapter 4 of WMO (2003). At that time,

global total ozone values were about 3% below the pre1980 levels. This reduction had occurred primarily in the

mid- to high latitudes of both hemispheres, with no significant trends reported in the tropics. Clear differences

between the two hemispheres were noted. Annually averaged total ozone over the northern midlatitudes was 3%

lower than pre-1980 levels, with twice as much decline in

winter than in summer. Over the southern midlatitudes,

by contrast, a long-term decrease of 6% had occurred,

roughly constant through the year. 

WMO (2003) concluded that models including

observed changes in halocarbons, other source gases, and

aerosols captured the long-term behavior in midlatitudes.

At that time, the differences between the Northern and

Southern Hemispheres were not explained. However, the

report did conclude that there was increased evidence that

the observed changes in atmospheric dynamics had a significant influence on ozone over northern midlatitudes on

decadal time scales. It also noted that the magnitude of

these influences, largest in the lower stratosphere, had not

been quantified unambiguously. At higher altitudes (35-

50 kilometers (km)), ozone trends were up to 8%/decade

over midlatitudes in both hemispheres, consistent with

expectation due to the observed changes in chlorine concentrations.

In this chapter, the same issues are examined with

the benefit of an additional four years of measurements

and a number of new analyses and interpretations.

Equivalent effective stratospheric chlorine (EESC) peaked

in the late 1990s to early 2000s, and so there are now several years of observations since that peak. There is an

encouraging consistency between the observations of

ozone and the behavior expected from EESC changes.

Major emphases of this chapter are whether the gradual

reduction in EESC, described in Chapter 1, can be found

in the observational record when broken down by latitude

and altitude, and whether the observational record can be

reproduced by models. The decadal influence of the

changes in dynamical processes is again an important issue

for Northern Hemisphere midlatitudes, and one that is critical for a complete understanding of total ozone changes

over the last decades.

Both dynamical and chemical processes may

contribute to decadal ozone (O3) changes in the lower

stratosphere. There is no single cause and the relative

contributions of different processes will change with

time period considered and between the hemispheres.

Chemical changes in midlatitudes may result from

changes in O3 loss processes occurring locally or those

at polar latitudes (see Chapter 4), whose effects are subsequently transported to midlatitudes. Both gas-phase

and heterogeneous chemical processes are important in

these processes. A major chemical driver for long-term

lower stratospheric ozone changes is increases in chlorine and bromine.

Previous Assessments described long-term ozone

changes in terms of linear trends, estimated using regression analysis. The decision to fit a linear trend was based

on the expected response to the approximately linear

increase in ozone-depleting substances (ODSs). However,

it has become clear that the change of ODSs after the mid1990s is no longer linear with time. In addition, ozone

data in the late 1990s to early 2000s do not follow the trend

line drawn by fitting the past data to a linear function.

Recent studies of long-term changes in ozone are focused

on detection of the ODS-related signal in the available

data. Typically this is done by including a term in statistical models that is proportional to the stratospheric burden

of ODSs, such as EESC or a similar function.

In this chapter, Section 3.2 presents the basic observations of ozone, with particular emphasis on behavior over

the past 4 years and a statistical interpretation. Section 3.3

presents observations of long-term changes in aerosols and

nitrogen dioxide (NO2) that are relevant to stratospheric

ozone. Finally, Section 3.4 discusses the dynamical and

chemical processes that lead to ozone changes and summarizes the use of “physical” models for quantifying the

contributions of the key processes to these changes.

3.2 UPDATE OF OZONE CHANGES

3.2.1 Statistical Methods

The general purpose of the use of statistical models

is to characterize and quantify the relationship between

“predictor” or “explanatory” variables and ozone. Some

of the explanatory variables describe natural variations in

ozone, while others are included in the model to reflect

possible anthropogenic changes in ozone. Multiple linear

regression is typically used to estimate statistical model

GLOBAL OZONE

3.5
parameters (e.g., SPARC, 1998; Weatherhead et al., 1998,

2000; Reinsel et al., 2002, 2005; Svenby and Dahlback,

2004; Krzy ́scin, 2006).

Seasonal cycle, solar flux (see Section 3.4.4), and

quasi biennial oscillation (QBO) indices are commonly

used as proxies for natural ozone variability. Characteristics of volcanic aerosols, such as aerosol optical

depth, are used in some models to account for effects of

volcanic eruptions, particularly for the El Chichón and Mt.

Pinatubo eruptions (see Section 3.3.1).

The way proxies are included in a regression model

can be an issue. Steinbrecht et al. (2004a) used harmonic

oscillations to describe ozone variations related to the 11-

year solar cycle, while Newchurch et al. (2003) assume

strict proportionality to the 10.7-cm solar radio flux with

no seasonal variation. Debate over the accuracy, robustness, and physical meaning of such choices of proxies is

found in Steinbrecht et al. (2004b) and Cunnold et al.

(2004). Stolarski et al. (2006) used the 10.7-cm solar radio

flux proxy and allowed for possible seasonal variations.

Harris et al. (2003) and Malanca et al. (2005) used a polynomial fit to separate long-term ozone changes from natural variability.

A linear function was used in earlier ozone

Assessments (WMO, 1995; 1999) to describe long-term

changes in ozone due to ODSs. As noted above, the

increase of ODSs is not approximated well by a linear

function over the time period from the late 1970s through

the 1990s and beyond. The piecewise linear trend concept, in which different linear fits are used before and after

a turning point, has been incorporated in several recent

trend analyses (e.g., Newchurch et al., 2003; Reinsel et

al., 2005; Miller et al., 2006). The turnaround date has

been chosen to be January 1996 (Reinsel et al., 2002; 2005)

or 1994 (Krzyscin et al., 2005).  ́

Instead of fitting a linear (or piecewise) trend to the

ozone time series, the changes in ozone can be analyzed

using the EESC function (Chapter 1). The EESC is an

index reflecting the amount of ozone-depleting chlorine

and bromine in the stratosphere (Chapter 8). Daniel et al.

(1995) suggested that a change in total column ozone can

be assumed to be generally proportional to the EESC. The

EESC trend model has been used in several recent statistical analyses of ozone (e.g., Newman et al., 2004; Guillas

et al., 2004; Yang et al., 2005; Fioletov and Shepherd,

2005; Dhomse et al., 2006; Stolarski et al., 2006). This

chapter is focused on detection of the EESC-related signals in available ozone records and so we use this method.

Piecewise trends and other methods used for the detection

of ozone recovery are discussed further in Chapter 6.

The contribution of different explanatory variables

to ozone fluctuations is illustrated in Figure 3-1 (upper

panel). It shows area-weighted seasonal mean total ozone

values for the region 60°S-60°N estimated from groundbased data, and the contribution (in Dobson units, DU) of

major components of the ozone variability to the total integral: the seasonal cycle, solar cycle-related signal, QBO,

and an estimated volcanic-related component. The EESCrelated term included in the regression model allows for

possible seasonal variations. The residuals are also shown.

Possible anthropogenic changes in ozone are isolated by

removing known natural components of ozone variability.

Total ozone for the region 60°S-60°N, with seasonal components and volcanic-, solar cycle-, and QBO-related components removed, are shown in Figure 3-1 (lower panel).

The seasonal component of the EESC-related signal is also

removed from the data to reduce seasonal variations of the

residuals. The plot illustrates the good agreement between

the long-term changes in ozone and the EESC.

For trends derived as a regression onto the EESC

time series as shown in Figure 3-1, the results can be

expressed in several ways:

(a) The results can be simply expressed as ozone change

per unit of EESC (e.g., % per 0.1 parts per billion by

volume (ppbv) EESC).

(b) The EESC time series is nearly linear for 1979-1989

(see Figure 3-1), with a net change of approximately

1.0 ppbv of chlorine. Thus the EESC fit to ozone can

be expressed in terms of linear changes for this time

period, with results reported in ozone changes (% or

DU) per decade (as in Stolarski et al., 2006).

(c) Ozone changes associated with EESC can also be

expressed as simple differences between two time

periods.

Here we follow (b) and show results expressed as percent

per decade for 1979-1989. The net change in ozone

between 1979 and 2005 is approximately 1.4 times these

linear trend values (i.e., a −5% per decade trend corresponds to approximately a −7% net change for 1979-

2005). Approach (b) also makes it easier to compare

EESC-related trends discussed here with linear ozone

trends reported in the earlier Assessments.

There are certain issues related to the use of EESC

as a regressor. The EESC calculation depends on the age

of air spectrum, which varies with altitude and latitude

(see Newman et al. (2006) and Box 8-1 of Chapter 8 for

details). The age of air is about 3 years in the midlatitude

lower stratosphere. It is longer, about 6 years, over the

poles and in the upper stratosphere. Ozone dilution from

the Antarctic, and to a lesser degree from the Arctic, affects

lower stratospheric ozone over midlatitudes (Section

3.4.3) and, strictly speaking, the “polar” EESC should also

GLOBAL OZONE

3.6
be used as a contributing factor when midlatitude trends

are considered. However, due to the shape of the EESC

curve, a 3-year difference in the age of air has limited effect

on the estimated EESC-related regression coefficient. For

example, the annual trend for 60°S-60°N is −1.6% ±

0.17% and −1.3% ± 0.15% per decade (2σ uncertainties

here and elsewhere in this chapter) for EESC calculated

with an age of air of 3 and 6 years, respectively. However,

the age of air plays a major role when ozone recovery is

considered (see Chapter 8).

Care must be used when comparing trends derived

from data in different vertical coordinate systems. In the

presence of a temperature trend, the various representations of the changes in ozone profiles will give different

trends because of the changing altitudes of pressure surfaces and the changing air density on pressure surfaces

with time. These trend differences will depend on the magnitude of the temperature trend and are important only in

the upper stratosphere. Rosenfield et al. (2005) demonstrated that between 1979 and 1997, ozone trend differences at 3 hectoPascals (hPa) are as large as 1-2%/decade

depending on how ozone changes are reported, with the

largest differences in the southern high latitudes. The most

negative trend is computed when ozone profiles are given

as number densities on geometric altitude, while the least

negative trend is computed for number densities on pressure levels. The Stratospheric Aerosol and Gas Experiment (SAGE) instrument measures ozone profiles as

number densities on geometric altitude, while Solar

Backscatter Ultraviolet (SBUV) data are provided as

mixing ratios on a pressure grid. Thus, upper stratospheric

ozone trends computed from SAGE data will be larger

than those computed from SBUV data when the trends are

computed in the natural vertical coordinate for each instrument. For comparison with SBUV, SAGE data are typically converted from number density on altitude levels to

mixing ratio on pressure levels (e.g., Wang et al., 1996;

McPeters et al., 1994). However, errors in the temperature data used to carry out this conversion can then affect

the computed trends. Li et al. (2002) estimated that adjustment for the temperature trends would reduce differences

in SBUV trends relative to SAGE by about 1%/decade, in

agreement with Rosenfield et al. (2005).

3.2.2 Changes in Total Ozone 

Since the previous Assessment (WMO, 2003),

satellite data from Total Ozone Mapping Spectrometer

(TOMS) and SBUV(/2) instruments have been

reprocessed with the new version 8 algorithm (Bhartia et

al., 2004). This has resulted in better agreement between

satellite and ground-based datasets, particularly in the

1980s and early 1990s. The version 8 algorithm was also

GLOBAL OZONE

Figure 3-1. Top: Ozone variations for 60°S-60°N estimated from ground-based data and individual components

that comprise ozone variations. Bottom: Deseasonalized,

area-weighted total ozone deviations estimated from

ground-based adjusted for solar, volcanic, and QBO effects,

for 60°S-60°N. The thick yellow line represents the EESC

curve scaled to fit the data from 1964-2005.

3.7
applied to 1970-1977 measurements from the Nimbus 4

Backscatter Ultraviolet (BUV) instrument, and these data

are shown in figures below. A new version of Global

Ozone Monitoring Experiment (GOME) data was also

used here. Details of the instruments and datasets can be

found in Appendix 3A.

The approach used in WMO (2003) is again used

here (Fioletov et al., 2002). Five datasets of 5°-wide zonal

averages of total ozone values are analyzed in this

Assessment and described in the Appendix. Areaweighted seasonal (3-month) averages are calculated for

different latitude belts and for the globe. Each dataset has

been deseasonalized with respect to the period 1979-1987,

followed by subtraction of the average of the monthly

mean anomalies for 1964-1980, estimated from groundbased data. Deviations are expressed as percentages of

the ground-based time average for the period 1964-1980.

As a part of the deseasonalizing process, the longterm trends for individual seasons have been removed and

replaced by the average of seasonal trends, as was previously done in WMO (1999). Unlike WMO (1999), where

the trend was approximated by a straight line, the EESC

curve is used here as a proxy for the trend function. For

regions with missing data, ozone deviations are assumed

to be identical to the surrounding latitude belts where data

are available.

The total ozone deviations for the 90°S-90°N, 60°S60°N, 35°-60°N, and 35°-60°S latitude belts are shown in

Figures 3-2 and 3-3. These plots differ slightly from those

of WMO (2003) because the satellite data processing algorithms have changed. The ozone deviations in the most

recent years are similar to those in the late 1990s-early

2000s, with no clear change in magnitude. A decadal periodic component related to the 11-year solar cycle is discussed in Section 3.4.4. Another distinct feature seen in

the unsmoothed data plot is a periodic structure with a

period of 2-3 years that is likely associated with the QBO.

These components are estimated and removed using statistical models as described above. Furthermore, to

remove short-term fluctuations, annual averages were calculated (Figure 3-2 and 3-3, bottom). Comparing hemispheres, midlatitude ozone values in 2002-2005 are about

3% lower than the pre-1980 level in the Northern

Hemisphere (NH) and about 5.5% lower in the Southern

Hemisphere (SH) (Figure 3-3), i.e., they are similar to

those reported in WMO (2003).

Figure 3-4 shows the total ozone deviations in the

tropics (25°S-25°N), where about 40% of global ozone is

located. No long-term changes in total column ozone have

been observed over the equator (e.g., Reinsel et al., 2005).

Outside the equatorial region, total ozone at 10°-25°S and

10°-25°N shows some signs of decline, but the trends are

not statistically significant. As discussed in Section 3.4.4,

there is a strong decadal periodic component in tropical

total ozone variations, probably related to the solar cycle.

The solar flux at 10.7 cm is shown at the top of Figure 3-

4. Pre-1980 ozone deviation estimates from ground-based

data for the tropics are less reliable because the number of

stations in that region was very limited and there were no

regular Dobson instrument intercomparisons prior to 1974

(Bojkov and Fioletov, 1996). Ground-based data do seem

to agree with Nimbus 4 BUV measurements; however,

BUV measurements were very sparse after 1972 (Stolarski

et al., 1997).

Total ozone changes in different seasons (DJF,

MAM, JJA, and SON) from 35°-60°S and 35°-60°N are

shown in Figure 3-5 for the period 1979-2005. The variability is high during the seasonal ozone buildup period in

winter and spring and is less through summer and autumn.

Long-term ozone declines in winter and spring seasons

are similar between hemispheres; ozone values are 5-6%

lower than the pre-1980 values. Summertime ozone deviations from the pre-1980 level are smaller in the NH (about

2%) than in the SH (about 5%).

Figure 3-5 shows that there is a clear difference in

long-term ozone variations over NH and SH midlatitudes.

Ozone variations over northern midlatitudes have a minimum in 1992-1995 (see also Figure 3-3 for annual averages). The large anomalies in winter-spring of 1992 and

1993 have been associated with the Mt. Pinatubo volcanic

eruption in June 1991 (Gleason et al., 1993; Bojkov et al.,

1993; Kerr et al., 1993; Hofmann et al., 1994). Low values

in the mid-1990s were followed by increases in the late

1990s-early 2000s (Reinsel et al., 2005; Yang et al., 2006).

This increase in column ozone is associated with the rise

in ozone content in the lower stratosphere below 18 km,

as described in Section 3.2.3 and further discussed in

Section 3.4.5. Similar features are not seen over southern

midlatitudes, where strong ozone minima occurred in 1985

and 1997 (Brinksma et al., 1998; Connor et al., 1999;

Cordero and Nathan, 2002). Differences in the seasonal

cycle of long-term ozone loss between NH and SH midlatitudes are partly caused by the export of ozone-depleted

air from the polar vortex (see Section 3.4.3.2). This effect

exists in both hemispheres, but its effect on spring- and

summertime ozone is stronger in the SH due to the larger

and more regular springtime ozone depletion over the

Antarctic.

3.2.3 Changes in the Vertical Distribution

of Ozone

There are a number of sources of vertically resolved

ozone data. Ozonesondes make in-situ measurements up

GLOBAL OZONE

3.8
GLOBAL OZONE

3.9

35°N - 60°N 35°S - 60°S

Ground-based

Figure 3-3. Top: Deseasonalized, area-weighted total ozone deviations from datasets for the latitude bands 35°N60°N (left) and 35°S-60°S (right). Anomalies were calculated with respect to the time average for the period 1964-

1980. Updated from Fioletov et al. (2002) and WMO (2003). Bottom: The same plot, but for annual averages.

Figure 3-2. Top: Deseasonalized, area-weighted seasonal (3-month average) total ozone deviations, estimated from different global datasets: ground-based, merged satellite dataset, National Institute of Water and

Atmospheric Research (NIWA) assimilated dataset, SBUV(/2), and GOME total ozone data. See Appendix 3A

for details. Each dataset was deseasonalized with respect to the period 1979-1987, the average of the monthly

mean anomalies for 1964-1980 estimated from ground-based data was then subtracted from each anomaly

time series, 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 (left) and the entire globe (90°S-90°N) (right). Updated

from Fioletov et al. (2002) and WMO (2003). Bottom: The same plot, but for annual averages.
to 30-35 km. Lidar and microwave instruments measure

ozone vertical profiles from the lower stratosphere to about

50 km (higher for microwave) and their records are now

long enough to assess long-term ozone changes (e.g.,

Schneider et al., 2003; Hartogh et al., 2004; Steinbrecht et

al., 2006). The Umkehr method provides profile data to

about 50 km from ground-based observations by Dobson

and Brewer instruments. A new version of the Umkehr

algorithm, UMK04, has been developed recently

(Petropavlovskikh et al., 2005a, b). Long records of ozone

profile retrieval from satellite observations are available

from the SAGE and SBUV and SBUV/2 series of instruments, although there are also satellite instruments with

shorter records, such as the Halogen Occultation

Experiment (HALOE) (e.g., Nazaryan et al., 2005). The

SBUV and SBUV/2 instruments have been in operation

since 1978, with six SBUV(/2) instruments having flown

since then. A new version 8 of SBUV(/2) data has recently

been released (Bhartia et al., 2004). First results have

demonstrated a significant improvement of the SBUV(/2)

data quality (McPeters et al., 2004). While there are some

differences between the SBUV and SAGE time series, they

both demonstrate similar features of long-term changes in

the vertical distribution of ozone. The biases between the

SAGE II and the SBUV record vary with time, according

to which of the individual SBUV(/2) sensors are used in

the long-term record (Nazaryan and McCormick, 2005;

Fioletov et al., 2006; Terao and Logan, 2006). These

biases, of several percent, are largest in the upper stratosphere and will contribute to differences in trends derived

from SAGE II and SBUV data. Data sources are discussed

in detail in the Appendix.

3.2.3.1 UPPER STRATOSPHERE

Figure 3-6 shows the evolution of upper stratospheric ozone relative anomalies, averaged between 35

and 45 km altitude, from lidar and microwave measurements, for five stations (Steinbrecht et al., 2006). The corresponding zonal averages from satellite data (SAGE,

SBUV, and HALOE) are also plotted. Anomalies are

defined as the deviation of individual monthly means from

the average climatological annual cycle. All records

starting before 1990 clearly show the long-term decline of

ozone in the upper stratosphere. Steinbrecht et al. (2006)

estimated that ozone in the upper stratosphere shows a

long-term decline of 10 to 15% since 1980. Upper stratospheric ozone trends before 1997 were about −6%/decade

at the northern midlatitude stations, almost −8%/decade at

Lauder in southern midlatitudes, and only −4.5%/decade

at subtropical Hawaii (Steinbrecht et al., 2006). This confirms similar findings of interhemispheric and latitudinal

GLOBAL OZONE

3.10

Ground-based

Figure 3-4. Deseasonalized,

area-weighted total ozone

deviations from five datasets

for the latitude bands 25°S25°N. The solar flux at 10.7 cm

is shown as a proxy for solar

variability.
GLOBAL OZONE

3.11

Ground-based

5°N - 60°N

5°S - 60°S

Figure 3-5. Seasonal 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). Updated from WMO (2003).
differences reported from SAGE data by Li et al. (2002)

and from SBUV(/2) data by Rosenfield et al. (2005).

Figure 3-6 also shows that ozone levels seem to be

inversely following the chlorine curve, and at many stations the previously steep ozone decline has not continued

in the last five years. Newchurch et al. (2003), using

SAGE data, also emphasized the lack of decline of ozone

in the upper stratosphere since 1997.

The vertical structure of ozone trends has been

derived from SAGE data (Wang et al., 2002; Li et al., 2002;

Randel and Wu, 2006) and SBUV data (Rosenfield et al.,

2005). Updated ozone trend estimates attributed to the

changes in the EESC as a function of latitude and altitude

estimated using SAGE I+II data and SBUV(/2) data for

1979-2004 are shown in Figure 3-7. Values are given in

% per decade for the period of linear change of the EESC,

and were calculated in the native units of each instrument.

GLOBAL OZONE

3.12

Figure 3-6. Upper stratospheric

ozone anomalies as measured by

ground-based lidars and microwave

radiometers at selected stations of

the Network for the Detection of

Atmospheric Composition Change

(NDACC), and by the corresponding

zonal means from the satelliteborne Stratospheric Aerosol and

Gas Experiments (SAGE, v6.20),

Halogen Occultation Experiment

(HALOE, v19), and Solar Backscatter

Ultraviolet instruments (SBUV). Anomalies are averaged over the altitude range from 35 to 45 km, or the

pressure range from 6.4 to 1.6 hPa

for SBUV (V8-merged dataset). For

clarity, 5-month running means are

plotted. The EESC in the figure is

for the age of air spectrum with

mean age of 6 years and width of 3

years, and zero bromine (see Box

8-1 of Chapter 8). Thin black curve

at the bottom shows the solar flux at

10.7 cm. Adapted from Steinbrecht

et al. (2006).
Also, as noted in Section 3.2.1, the SAGE data give larger

trends than SBUV data in the upper stratosphere due to

the temperature trend and the use of pressure (SBUV)

rather than geometric altitude (SAGE) coordinates. Some

of the differences are also likely caused by different biases

between ozone measured by SAGE and by the different

SBUV(/2) instruments. The strongest decline (7-9% per

decade, or 10-15% since 1979) occurred at 40-45 km over

midlatitudes. The overall pattern of upper stratosphere

ozone trends shown in Figure 3-7 is similar to that reported

in WMO (2003) from SAGE data for a shorter period and

using a linear trend function.

The vertical profiles of trends in ozone over

northern midlatitudes (35°-60°N) derived from SAGE,

SBUV, and Umkehr instruments are very similar for 30-

38 km. SAGE trends are larger than the SBUV and

Umkehr trends above 40 km (Figure 3-8). Similarly, over

35°-60°S, SAGE and SBUV trends are essentially the

same for 30-38 km and the SAGE trends are larger above

40 km. Despite the difference in the magnitude of the

trend, both SAGE and SBUV show that the decline at 40

km over SH midlatitudes is slightly larger than that over

NH midlatitudes, but the difference is within the error bars.

A slightly stronger decline over the SH was previously

reported from SAGE data in WMO (2003).

Upper stratospheric negative trends are generally

largest in winter and smallest in summer (Hood et al.,

1993; Hollandsworth et al., 1995). Ozone trends for four

seasons estimated using the EESC from version 8 SBUV

data for 1979-1997 are shown in Figure 3-9 (Rosenfield et

al., 2005). In both hemispheres the winter decline is larger

(maximum of −5.7 ± 2.8%/decade in the NH and −6.9 ±

2.8%/decade in the SH) than the summer decline (−3.1 ±

1.5%/decade and −4.9 ± 1.6%/decade, respectively).

Spring and autumn negative trends are also larger in the

SH than in the NH. The spring and autumn upper stratospheric trends lie in between the winter and summer trends.

3.2.3.2 LOWER STRATOSPHERE

Ozone values in the lower stratosphere from the late

1970s to the present from measurements made by sondes,

SAGE, SBUV(/2), and HALOE are shown over Europe

(Figure 3-10a) and over Lauder, New Zealand (Figure 3-

10b) (updated from WMO, 2003 for sondes; Terao and

Logan, 2006). There is good agreement among the various datasets. Ozone reached minimum values in 1993 in

the NH from 13 to 25 km (158-25 hPa) in the NH. Since

then, from 13 to 16 km (158-100 hPa), ozone values have

increased and are about the same as in the early 1980s,

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3.13

S AGE I+II SBUV(/2)

Figure 3-7. SAGE I+II and SBUV(/2) annual trends in percent per decade as a function of latitude and altitude

for the period 1979-2004. The trends were estimated using regression to an EESC curve and converted to

%/decade using the variation of EESC with time in the 1980s. The SAGE trends were calculated in geometric

altitude coordinates, and the SBUV(/2) trends were calculated in pressure coordinates. Shadings indicate that

the changes are statistically significant at the 2Û level.
while from 22 to 25 km (40-25 hPa), they are about 5-8%

lower than in the early 1980s.

Estimating long-term ozone changes below 20 km

from satellite data is still a challenge. SBUV(/2) version

8 Layer 1 data represent the partial ozone column from

approximately 7 to 19 km. However, it is a fairly new data

product and more validation is required. The use of SAGE

data is limited in this altitude range because the present

version of SAGE I data is not reliable below 20 km. The

ozone amount below a certain level in the lower stratosphere can be estimated by subtracting SAGE I and II partial column ozone above that level from total column

measured by TOMS. Figure 3-11 shows estimates of partial column ozone below 19 km from satellite data for

northern and southern midlatitudes from SBUV(/2) and

SAGE/TOMS data. (The SAGE partial column above 19

km is derived from a regression analysis of the SAGE profile measurements, as described in Randel and Wu, 2006.)

Figure 3-11 demonstrates that ozone levels below 19 km

in the NH in the 2000s are similar to these from the early

1980s but are ~10% lower in 1992-1995. In the SH, the

data also show a decline between 1980 and the mid-1990s;

however levels in the 2000s remain ~10% lower than the

levels of the late 1970s to early1980s.

GLOBAL OZONE

35°N - 60°N 35°S - 60°S Figure 3-8. Vertical profile of

ozone trends over northern and

southern midlatitudes estimated

from ozonesondes, Umkehr,

SAGE I+II, and SBUV(/2) for the

period 1979-2004. The trends

were estimated using regression

to an EESC curve and converted

to %/decade using the variation

of EESC with time in the 1980s.

The trends were calculated in

geometric altitude coordinates for

SAGE and in pressure coordinates for SBUV(/2), sondes, and

Umkehr data, and then converted

to altitude coordinates using the

standard atmosphere. The 2Û

error bars are shown.

3.14

Figure 3-9. Season-pressure sections of SBUV midlatitude linear ozone trends for (top) Northern

Hemisphere and (bottom) Southern Hemisphere for

1979-1997 in % per decade (Rosenfield et al., 2005).

Shaded areas are significant at the 2Û level.
Earlier Assessments reported a large negative trend

(more than 5%/decade) in the lower stratosphere below

18 km (WMO, 1992, 1999; SPARC, 1998; Logan et al.,

1999) over northern midlatitudes for the period up to the

mid-1990s. Amaximum decline of about −10% per decade

was reported at 15 km (see Figure 3-12). However, the

ozonesonde trends reported by WMO (2003) for 1980-

2000 at 12-18 km (200-63 hPa) are half as large as those

for 1980-1996. A similar reduction of the trend magnitude is shown by Umkehr data analysis (Bojkov et al.,

2002). Recent trend estimates over northern midlatitudes

for the period up to 2004 show no decline at 15 km (Figures

3-8, 3-12) and ozone values at 13-16 km are similar to

those in the early 1980s (Figure 3-10).

The dramatic change in ozone trend estimates at 12-

18 km results from a decline in ozone during the 1980s

and early 1990s followed by a rapid increase thereafter

(see Figure 3-10). This behavior can be partially attributed to low ozone values in the lower stratosphere in 1992

and 1993 (Figure 3-10) when the stratosphere was influenced by the Mt. Pinatubo eruption. Miller et al. (2006)

fit the sonde data from 13 NH stations for 1970-2003 with

a piecewise linear trend, with the change in trend specified at January 1996. They removed two years of data

after June 1991 to reduce the influence of the Mt. Pinatubo

eruption. Nevertheless, they found that the trend is about

−3 to −4%/decade at 15-18 km for 1970-1995, with a statistically significant change in trend of +9 to +13%/decade

GLOBAL OZONE

3.15

M hly ly O [%]

ear ear

Figure 3-10. (a) Monthly ozone anomalies for Europe as measured by ozonesondes (black line), SAGE II (red

circles), HALOE (blue circles), and SBUV(/2) (green line) at at four pressure layers. The monthly anomalies

were calculated as the difference between a given monthly mean and the average of monthly means for 1985-

1990 for each dataset, except for HALOE where 1994-1999 was used; the average of the monthly mean ozonesonde anomalies for 1979-1981 was then subtracted from each anomaly time series. The ozonesonde data are

the average of measurements at three European stations: Hohenpeissenberg, Payerne, and Uccle. The satellite data were selected within a grid box of 45°N-55°N and 10°W-30°E. A 3-month running mean was applied to

the data. For SAGE II and HALOE, results are shown when the number of observations is larger than 4 per

month. (b) Same as (a), but for Lauder, New Zealand. The monthly anomalies were calculated using the 1987-

1991 monthly means for each dataset, except for HALOE where 1994-1999 was used. The satellite data were

selected within a grid box of 40°S-50°S and 150°E-170°W. Adapted from Terao and Logan (2006).
after January 1996, i.e., there was a positive trend of 6 to

11% per decade in 1996-2003. Similar results are reported

by Yang et al. (2006) and their study is discussed in detail

in Chapter 6. The piecewise method gives an increase in

ozone after 1996 because of the low ozone values in the

first half of the 1990s and the higher values in the early

2000s (Figure 3-10a). The pre-1996 trends for 12-18 km

are largest in spring, as in earlier studies (Logan et al.,

1999), and the trend change is also largest in spring (Miller

et al., 2006).

Ozonesonde data from Lauder, New Zealand

(45°S), show no significant trends at 12-19 km (Figure 3-

10b). The lack of trend in the Lauder record, which begins

in late 1986, is not inconsistent with the satellite data in

Figure 3-11, because much of the ozone decline over

southern midlatitudes occurred before 1986. An ozonesonde record from earlier years is available from

Melbourne, Australia. It demonstrates a significant decline

of about 10% per decade at 100 hPa (~16 km) for 1970-

1990 (Logan, 1994; WMO, 1995). However, the record is

affected by infrequent measurements for many years, as

well as changes in sonde type and site location, and further analysis is required for trend studies.

At 18-25 km, ozone values have been relatively flat

in the NH since the late 1990s, as shown in Figure 3-10.

According to Miller et al. (2006), at 20-22 km, the trend is

−4.5%/decade for 1970-1995, with a change in trend of 2-

4%/decade after January 1996. Yang et al. (2006) suggested that the cessation of ozone depletion at 18-25 km is

consistent with a leveling of the EESC curve. The ozone

trend estimated using the EESC curve is about −3 to

−4%/decade at 18-24 km (Figures 3-7 and 3-8) in the both

hemispheres. In the past five years, ozone values at 22-25

km (40-25 hPa) were about 5% below those in the early

1980s. As shown by the ozone time series discussed above

and the EESC curve, this decrease occurred prior to about

1996.

As was previously reported by SPARC (1998) and

WMO (1999, 2003), the trend magnitude over midlatiGLOBAL OZONE

3.16

35°N - 60°N

35°S - 60°S

Figure 3-11. Deseasonalized, seasonal (3-month

average) ozone deviations from SBUV and SAGE II

data for the latitude bands 35°N-60°N (top) and 35°S60°S (bottom) below 19 km. Anomalies were calculated with respect to the time average during 1979-

1982. The blue line represents SBUV(/2) Layer 1

data (1000-63 hPa). The red line represents the difference between TOMS column ozone and SAGE

I+II data vertically integrated above 19 km. The

SAGE partial column above 19 km is derived from a

regression analysis of the SAGE profile measurements, as described in Randel and Wu (2006).

Figure 3-12. Ozone trends over northern midlatitudes

(36°N-59°N) as a function of altitude in %/decade and

DU/km/decade estimated from ozonesonde data for

1980-1996 (dotted line) and 1980-2004 (solid line).

The trends were estimated using regression to an

EESC curve and converted to %/decade and DU/-

decade using the variation of EESC with time in the

1980s. The error bars represent the 95% confidence

limits. Updated from Logan et al. (1999).
tudes has its minimum near 30 km. The SAGE I+II data

and the ozonesonde data give almost no trend related to

EESC at 30 km in northern midlatitudes, while the SBUV

data give a small negative trend (−2 to −3%/decade)

(Figures 3-7 and 3-8).

In the tropics, the SAGE data in Figure 3-7 show

negative trends in the lower stratosphere, while the SBUV

data show no significant trends in tropical ozone from 22

to 30 km. WMO (2003) showed significant negative

trends in the tropical lower stratosphere below 22 km

based on SAGE I+II data, and these changes have been

accentuated by several recent years of low ozone (Figure

3-7; Randel et al., 2006). However, these results should

be viewed with caution as there are no other long-term

records for ozone in the tropical lower stratosphere. This

is a region of small ozone concentrations and large vertical gradients, which present challenges for satellite

observations. On the other hand, interannual changes in

SAGE II data show good agreement with tropical ozonesonde data from the SHADOZ (Southern Hemisphere

Additional Ozonesondes) network (Thompson et al., 2003)

for 1998-2004 (Randel et al., 2006). 

3.2.3.3 CONSISTENCY OF COLUMN AND PROFILE

TRENDS

The vertical profile trends over midlatitudes discussed in Section 3.2.3 are in general consistent with

column ozone trends there. Less than 10% of the midlatitude total column ozone is located above 35 km, and a 10-

15% decline there accounts for a total column decline of

about 1%. The lower stratosphere between 20 and 25 km

contains 20-25% of the total column ozone, and an 8-10%

decline in this layer contributes a total column decline of

about 2%. Ozone located below 18 km contributes about

30% to the total column. As shown in Figure 3-11, there

is no ozone decline below 19 km over northern midlatitudes and about a 10% decline at southern midlatitudes,

yielding a total column decline of about 3%. Thus, the

overall total column decline from the vertically integrated

profiles is about 3% over northern midlatitudes and about

6% over southern midlatitudes; these are similar to the

declines in column ozone described in Section 3.2.2.

As mentioned in Appendix 3A, SBUV(/2) total

ozone used in Section 3.2.2 is the vertically integrated

SBUV(/2) ozone profiles. Therefore, SBUV(/2) column

ozone and integrated profile trends are necessarily consistent. Section 3.2.2 shows that SBUV(/2) column ozone

data agree with total ozone data from the other sources. 

Figure 3-13 shows a latitudinal profile of the vertically integrated SAGE I+II ozone changes for 1979-2005

compared with corresponding changes in column ozone,

derived from the merged satellite ozone data. The SAGE

results represent a vertical integral of the profile trends

from the tropopause to 50 km. Results from both datasets

show reasonable agreement for net ozone decreases over

the extratropics. There are significant differences, however, in the magnitude of changes in the tropics, with the

integrated SAGE results giving larger changes than the

column measurements (which are near zero in the equatorial region). This difference could have several implications. There could be compensating positive trends in

tropical tropospheric column ozone, with increases of ~6

DU over 1979-2005 (~15% of background levels).

Tropical tropospheric ozone trends for this period are not

well known. Lelieveld et al. (2004) have reported substantial increases in near-surface ozone in the tropical

Atlantic ocean for 1972-2002, and Bortz et al. (2006) suggest a 20% increase in tropical upper tropospheric ozone

during 1994-2003, based on aircraft measurements. On

the other hand, Ziemke et al. (2005) suggested that there

have not been significant trends in tropical tropospheric

ozone for this period based on satellite observations. An

alternative explanation of the tropical differences in

GLOBAL OZONE

3.17

Figure 3-13. Latitudinal profile of ozone trends

derived from the vertically integrated SAGE I+II data,

compared with results from the merged TOMS/SBUV

column ozone dataset. The trends were estimated

using regression to an EESC curve and converted to

%/decade using the variation of EESC with time in

the 1980s. The SAGE I+II results are a vertical integral of the changes in Figure 3-7, integrated between

the tropopause and 50 km. For comparison, the

dashed line shows the SAGE results integrated over

a limited range of 25-50 km (for 30°N-30°S). Adapted

from Randel and Wu (2006).
Figure 3-13 is that the ozone changes derived from SAGE

data are simply too large, particularly the lower stratospheric maximum (as discussed above). For comparison,

Figure 3-13 also includes the SAGE I+II changes integrated over the limited altitude range of 25-50 km (for

30°N-30°S), and this produces smaller trends that border

on the statistical uncertainty levels for the tropical column

trends.

3.2.4 Tropospheric Ozone

A recent analysis of long-term changes in tropospheric ozone emphasizes that trends vary regionally and

that within a given region, trends have changed over the

past 25-35 years (Oltmans et al., 2006). This section

updates the discussion in WMO (2003) and focuses on

ozonesonde data and surface data at remote, or relatively

remote, sites.

Tarasick et al. (2005) showed that tropospheric

ozone over Canada (53°-75°N) decreased from 1980 to

around 1994 and then increased up to the present, with

ozone values in the early 2000s similar to those in the early

1980s. They also found that annual average ozone anomalies in the troposphere at the Canadian sites are correlated

with those in the lower stratosphere (r = ~0.65). Ozone

over the eastern U.S.A. (Wallops Island) shows relatively

small changes from 1970 to the present (<5%), with highest

ozone in the late 1980s (Oltmans et al., 2006). An analysis

of short-term trends in MOZAIC (Measurements of Ozone

and Water Vapor by In-service Airbus Aircraft) aircraft data

finds an increase of 12%/decade in the middle troposphere

over New York from mid-1994 to 2001; the increase is

much larger in winter (20%/decade) than summer (6%/

decade). Ozone was particularly high in 1998 and 1999

(Zbinden et al., 2006).

Over Europe, the sonde data from Hohenpeissenberg

(the longest self-consistent record) show a steady increase

from 1966 to the mid- to late 1980s, followed by a slow

decrease and considerable interannual variability. Values in

the past few years are similar to those in the early 1980s

(Oltmans et al., 2006). Sonde data from Payerne and Uccle

show little change in ozone in the mid-troposphere in the

last decade. Oltmans et al. (2006) compare mean values for

1995-2004 to those for 1985-1994 for Hohenpeissenberg

and the nearby mountain station of Zugspitze (~3000 m).

The sonde data show little change except a small decrease

in May-July, while the Zugspitze data show a small increase

in summer, and a larger increase in winter. Simmonds et al.

(2004) report an increase in ozone of 13%/decade in background air at Mace Head (Ireland) from 1987 to 2003, with

highest values in 1998 and 1999. The MOZAIC data over

Frankfurt and Paris also give an increase in ozone, 7 and

15%/decade, for 1994 to 2001, with the largest trend in

winter and spring and almost no trend in summer (Zbinden

et al., 2006). 

Jonson et al. (2006) recently summarized reports

on trends over Europe and noted that the differences

between the trends given by sondes, and those from

MOZAIC, mountain sites, and Mace Head cannot easily

be reconciled. They also note that surface sites within

Europe generally show substantial downward trends in

high ozone (95th or 98th percentiles) over the past 10-15

years, as well as increases in ozone at polluted surface

sites. Both these features are caused by substantial

decreases in emissions of ozone precursors in Europe,

leading to less ozone formation in summer, and a reduction in titration of ozone by nitric oxide (NO) in winter

(e.g., Jonson et al., 2006).

Long-term ozonesonde measurements over Japan

(32°-43°N) show an overall increase since 1970, but most

of this occurs early in the record, and the last few years

show a downturn in ozone (Oltmans et al., 2006). Naja

and Akimoto (2004) find that ozone below 3 km in regionally polluted air over Japan increased by 11-22% from the

1970s to the 1990s. The MOZAIC data over Japan give

an increase of 8%/decade for 1994-2001, with no trend in

summer and an increase of ~10%/decade in other seasons

(Zbinden et al., 2006). 

In summary, the sonde data for northern midlatitudes for the last decade show increases in ozone over

Canada and slight decreases over Europe and Japan. The

average trend for the sonde data for 1980 to 2004 is zero

(Figure 3-12).

There has been a long-term increase in ozone over

Hawaii of 3.5 ± 1.5%/decade during autumn and winter,

in both surface data at 3.4 km and sonde data at 3-6 km,

that appears to be related to shifts in transport patterns

(Oltmans et al., 2006). There is little change in ozone since

1985 in spring, when transport events from Asia occur.

Indian ozonesonde data demonstrate a statistically

significant increase of tropospheric ozone (43% ± 25%

per decade at 800-1000 hPa and more than 50% per decade

at 200-500 hPa) over Delhi for the period 1972-2001

(Saraf and Beig, 2004). However, Indian sondes have

demonstrated large, more than 30%, uncertainties in the

troposphere during international ozonesonde intercomparisons (Attmanspacher and Dütsch, 1970; SPARC, 1998)

and these trend results should be interpreted with caution.

Analysis of MOZAIC data indicates that that tropical ozone at cruise altitudes (7.7-11.3 km) has increased

by ~1 ppb/year, or by ~20%/decade, from 1994 to 2003

(Bortz et al., 2006). This is twice as large as the increase

reported for surface ozone over the tropical Atlantic from

shipboard data by Lelieveld et al. (2004). Surface and

GLOBAL OZONE

3.18
sonde data up to the mid-troposphere from Samoa (14°S)

show a small decrease from the late 1980s to the most

recent decade in austral winter, and no change the rest of

the year (Oltmans et al., 2006). Unfortunately, the longest

tropical sonde record from Natal (Brazil) suffers from

changes in sonde types and procedures, as well as data

gaps, and the record is not suitable at present for deriving

reliable trends.

The Samoa sonde data since 1995 do not show the

increase seen in the MOZAIC data for the southern tropics.

The MOZAIC data are from South America, the Atlantic,

and Africa, while the sonde data are from the Pacific.

Ziemke et al. (2005) found no trend in the tropospheric

column of ozone derived from TOMS data in the tropics

from 1979 to 2003. There appear to be inconsistencies

among the various data records for tropospheric ozone.

In southern midlatitudes, ozone at Lauder (45°S)

increased from 1986 to present by ~5%/decade below 3

km, with a marginally significant increase of 2%/decade

at 3-6 km, and no trend at 6-9 km (Oltmans et al., 2006).

Sonde data from the South Pole show a small increase in

ozone of 0.5%/decade since 1986 in the mid-troposphere,

primarily in spring and summer. No such increase is seen

at Syowa (69°S) (Oltmans et al., 2006).

It appears that the short-term trends from MOZAIC

data for both the midlatitudes and the tropics may be inconsistent with sonde data, but more analysis is needed.

3.3 SUMMARY OF OTHER OBSERVATIONS

The observed long-term changes in temperature and

water vapor are discussed in Chapter 5. In addition, observations of past changes in long-lived source gases and

halogen-containing species are discussed in Chapter 1. In

this subsection we therefore focus on the remaining parameters that are important for long-term changes in ozone

and for which there are relevant data records.

3.3.1 Stratospheric Aerosol and Its

Precursors

The source of stratospheric aerosol observed since

the late 1970s has been dominated by the injection of

sulfur dioxide (SO2) from a few major volcanic events

including El Chichón in 1982 and Mt. Pinatubo in 1991.

The inference of a nonvolcanic stratospheric background

is hampered by very limited periods without volcanic

influence since systematic measurements began; however, the best indications are that there is no long-term

trend in the background aerosol level (Deshler et al.,

2006). For background periods, observed since the late

1990s, the dominant stratospheric aerosol precursor gases

are carbonyl sulfide (OCS) and SO2. Through SO2,

human-related activities may influence the observed

background stratospheric aerosol. There is general agreement between measured OCS and modeling of its transformation to sulfate aerosol, and observed aerosols.

However, there is a significant dearth of SO2 measurements, and the role of tropospheric SO2 in the stratospheric aerosol budget, while significant, remains a matter

of some guesswork. In addition, it is not well understood

whether global human-derived SO2 emissions, which are

decreasing, or emissions in low latitude developing countries, such as China, which are increasing, dominate the

human component of SO2 transport across the tropical

tropopause (e.g., Notholt et al., 2005).

As shown in Figure 3-14, aerosol loading during

the past 20 years has varied by as much as 2 orders of

magnitude. As a result, the quality of the measurements

and the agreement among diverse measure systems varies

strongly as a function of the level of aerosol loading. For

instance, results from the SPARC Assessment of Stratospheric Aerosol Properties (ASAP) (SPARC, 2006)

demonstrate that space-based and in situ measurements

of aerosol parameters, in particular surface area density

(SAD), tend to be consistent following significant volcanic events like El Chichón and Mt. Pinatubo. However,

during periods of very low aerosol loading, this consistency breaks down and significant differences exist

between systems for inferred parameters like SAD and

directly measured parameters like aerosol extinction coefficient (SPARC, 2006). Although integrated aerosol

GLOBAL OZONE

3.19

Figure 3-14. The Assessment of Stratospheric Aerosol

Particles (ASAP) reconstruction of the tropical (solid

line) and northern midlatitude (dotted) 1020-nm

aerosol optical depth from 1979 through 2002 based

on a combination of SAGE, SAGE II, and groundbased and airborne lidar data. From SPARC, 2006.
quantities such as surface area density and effective radius

can be calculated without approximation from a known

size distribution, the satellite and in situ observational

bases for size distributions are controlled by a priori

assumptions regarding the distribution itself or by having

coarse size resolution, respectively. ASAP showed that

during volcanically quiescent periods, models and observations disagree significantly mainly because of the modeled fraction of the surface area density that resides in

particles too small to be measured, especially near nucleation regions. While there are some model shortcomings

relative to observations, particularly in the lower stratosphere, it seems likely that space-based datasets underestimate, perhaps significantly, aerosol surface area density

in the lower stratosphere during low loading periods.

In the past, the effect on ozone variability of stratospheric aerosol driven by sporadic volcanic events (primarily the El Chichón and Mt. Pinatubo eruptions) has

given concern that it could mimic the effects of the 11-

year solar cycle. However, since stratospheric aerosol has

remained near nonvolcanic background levels since the

late 1990s through much of the recent solar maximum,

such concerns have been greatly mitigated (see Section

3.4.4).

3.3.2 Nitrogen Dioxide (NO2)

Long-term ground-based observations of column

NO2 have been made at Lauder, New Zealand (45°S), since

late 1980, by UV/visible observations of sunlight scattered

from the zenith. Trends in the observations were discussed

in Liley et al. (2000) and WMO (2003). Here the observations and analysis are updated through early 2006. A

statistical analysis of geophysical cycles (solar cycle,

QBO, El Niño-Southern Oscillation), trends, and volcanic

eruptions was done to assess correlations with the observed

NO2. Linear increases of 6.2 ± 1.8% per decade (am) and

5.7 ± 1.1% per decade (pm) are inferred. These are not

significantly different from those reported in Liley et al.

(2000). Figure 3-15 shows the time series for monthly

averages of morning and afternoon column NO2 at Lauder,

after subtraction of the seasonal cycle.

NO2 columns have also been measured at the

Jungfraujoch (47°N) since 1985 (Mahieu et al., 2000). The

Jungfraujoch measurements are made by Fourier transform spectroscopy (FTS) in solar absorption in the

infrared. The monthly averaged vertical column densities

from 1985-2005 are also shown in Figure 3-15. Analysis

of these data using the same algorithm applied to the

Lauder data produces a linear trend of 1.5 ± 1.0% per

decade. The trend reported by Mahieu et al. for 1986-1998

(7 ± 3% per decade) may have been significantly affected

by not explicitly accounting for the effect of the Mt.

Pinatubo eruption from 1992-1994. SAGE II satellite

observations of profile NO2 (Cunnold et al., 1991) are consistent with the Lauder data and show that the bulk of the

increase occurs in the lower stratosphere. The SAGE II

data also show a large hemispheric asymmetry, with no

significant trend at Northern Hemisphere midlatitudes.

Thus, the available evidence suggests that the trend

of NO2 in the Southern Hemisphere is significantly larger

than that in the Northern Hemisphere. The trend at the

Jungfraujoch is consistent with the 2.4%/decade increase

in tropospheric N2O reported in WMO (2003). Other factors may be required to explain the larger increase at

Lauder. WMO (2003) concluded that observed nitrous

oxide (N2O) increases and deceases in ozone explained a

trend in NO2 of 5 ± 1%/decade. The origin of the apparent

hemispheric difference in trend remains unexplained,

although uncertainties in the analysis due to the relatively

short data records should be noted.

GLOBAL OZONE

3.20

Figure 3-15. Deseasonalized anomalies of the NO2

slant column measured at two sites, the Jungfraujoch, Switzerland (47°N), and Lauder, New

Zealand (45°S). The Jungfraujoch data are measured in direct, solar absorption in the infrared, at a

wide range of zenith angles, while the Lauder data

are UV/visible absorption of sunlight scattered from

the zenith at solar zenith angles of 90°. Sunrise

measurements are shown as dashed lines, and

sunset measurements as solid lines.
3.4 UNDERSTANDING AND INTERPRETATION

OF OZONE CHANGES

3.4.1 Models

In order to quantify the effect of certain processes

on long-term ozone changes, we need to describe these

processes by numerical models. These models can either

be “physical” models, which solve the governing equations of the physics/chemistry of the atmosphere, or “statistical” models, in which changes in the long-term ozone

time series are linked with forcing terms, typically by multiple regression (see Section 3.2.1).

In this chapter, we use the term “physical model”

to describe a numerical model that integrates the relevant

equations of the chemistry and dynamics of the atmosphere to predict the evolution of the atmosphere with time.

These models are a mathematical representation of our

current understanding of the chemistry and physics of the

atmosphere. These models can be 0-D (box), 1-D

(column), 2-D (latitude-height), or 3-D models. The 3-D

models can be classified as “off-line” chemical transport

models (CTMs), which are forced by specified winds, or

general circulation models (GCMs), which calculate their

own winds and temperatures. The components of such

models were described in Chapter 4 of WMO (2003).

GCMs that contain interactive, detailed chemistry schemes

are called coupled Chemistry-Climate Models (CCMs).

The various types of models are also described in more

detail in Chapter 5 (Box 5-1).

Aside from increased computer power allowing

more 3-D simulations, and therefore a better understanding

of the performance of these models, a major development

in these tools over the past four years for understanding

past ozone changes has been the new availability of longterm, whole-stratosphere meteorological reanalyses that

can be used to force CTMs.

The European Centre for Medium-range Weather

Forecasts (ECMWF) and the National Centers for

Environmental Prediction (NCEP) have extended the

upper boundary of their operational meteorological

analyses to the lower mesosphere (0.1 hPa, more recently

0.01 hPa) and upper stratosphere (1 hPa), respectively.

The United Kingdom Met Office (UKMO) has provided

meteorological analyses extending to 0.3 hPa since 1991

(Swinbank et al., 2002). The Global Modeling and

Assimilation Office (GMAO) of the National Aeronautics

and Space Administration (NASA) also produces meteorological analyses for the troposphere and stratosphere.

The Goddard Earth Observing System, version 4 (GEOS-4)

data assimilation system (Bloom et al., 2005) has an upper

boundary near 80 km (0.01 hPa), with data constraints up

to about 55 km. Furthermore, both ECMWF and NCEP

have performed 40- and 50-year reanalyses labeled ERA40 (Simmons et al., 2005; Uppala et al., 2005) and REAN

or NCEP-50 (Kanamitsu et al., 2002), respectively. ERA40 extends to 0.1 hPa, although REAN only extends up to

10 hPa. These reanalyses have been performed with a 3-

D variational analysis (3D-VAR), while the operational

analyses are performed with the computationally more

expensive 4-D variational analysis (4D-VAR), which is

regarded as more accurate, due to the re-integration in

time.

Some general issues relating to the quality of wind

datasets used to force CTMs have been identified.

Schoeberl et al. (2003), Strahan and Douglass (2004), and

Tan et al. (2004) showed that GMAO and UKMO winds

are too dispersive as a result of the assimilation procedure.

The procedure causes excessive mixing in the tropical

lower stratosphere, forcing large-scale meridional tracer

circulation in the stratosphere. The consequence is that

the mean residence time of air in the stratosphere becomes

too short, which has a negative impact on long-term tracer

integrations. Schoeberl et al. (2003) showed that winds

from GCMs are much less dispersive and contain a weaker

residual circulation, leading to a mean age of air closer to

observations.

The ERA-40 reanalyses extend from 1957 to 2001

and, because they cover the whole stratosphere, were

expected to provide a much better representation than previous analyses, which extended only up to 10 hPa. Such

analyses (e.g., ECMWF ERA-15 reanalyses for the period

1978 to 1993) were used in studies discussed in WMO

(2003). These ERA-40 reanalyses have now been quite

extensively tested in long-term transport studies and the

winds have been applied for ozone integrations. These

studies have shown that the reanalyses are a very important resource for diagnosing past ozone changes but, due

to some clear problems, some caution is needed in their

use. The ERA-40 data has a much stronger circulation

than the operational data, due to the less balanced assimilation procedure in 3D-VAR, which has been shown by

ozone integrations (Laat et al., 2006). In addition, ERA40 suffers from inhomogeneity, introduced by the use of

spaceborne observations in the data assimilation as well

as by changes in the satellite instrumentation. This was

illustrated from different diagnoses, e.g., downward ozone

fluxes (Van Noije et al., 2004) and water vapor distributions (Bengtsson et al., 2004). They showed significant

discontinuities in the strength of the ozone flux at 100 hPa.

Different datasets of assimilated winds and temperature (operational and reanalyses) have been evaluated for

the Arctic winter (Manney et al., 2003) and the 2002

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3.21
Antarctic polar vortex (Manney et al., 2005) by comparisons with observations. The large-scale dynamical structures were well represented in all datasets, but noticeable

differences appeared in several diagnoses. The REAN and

NCEP show warm biases in the temperatures, while the

ERA-40 dataset shows unrealistic spurious temperature

oscillations, which were also reported by Simmons et al.

(2005). Manney et al. (2005) also show that between the

different meteorological datasets, there is little consensus

in mixing of vortex air with midlatitudes. Overall, they

concluded that caution must be taken when applying a certain assimilated wind dataset, and that the use of some

reanalysis data for detailed studies of polar dynamics and

chemistry is not recommended. Additional evidence of

inhomogeneities in the reanalysis datasets was given by

temperature and tropopause height trends analyses (Santer

et al., 2004). However, they also concluded that the ERA40 reanalysis has much improved compared with the earlier ERA-15 and NCEP-50 reanalyses. Birner et al. (2006)

also noted the effect of data inhomogeneities on

tropopause characteristics. Despite the caveats, the interannual variability of tracers using ERA-40 is simulated

quite well (Hadjinicolaou et al., 2005; Feng et al., 2006).

Even when applying the operational 4D-VAR

ECMWF data, the large-scale circulation in pressurecoordinate CTMs remains too strong (Meijer et al., 2004;

Van Noije et al., 2006). However, there are different

approaches to the use of analyzed winds to force an offline CTM. For example, the use of an isentropic vertical

coordinate results in less noisy vertical motion, shown by

a reduction of air parcel dispersion in the tropical lower

stratosphere and increased mean age of air (Mahowald et

al., 2002; Schoeberl et al., 2003; Chipperfield, 2006).

Another approach to reducing the noise in the wind data

uses forecasts instead of analyses (Meijer et al., 2004;

Scheele et al., 2005; Laat et al., 2006). In addition, the

use of 3-hourly instead of the usual 6-hourly analyses has

been shown to improve stratospheric tracer distribution at

northern midlatitudes (Berthet et al., 2005).

Nevertheless, the apparent dispersive character of

assimilated winds remains, and it seems inherently connected to the assimilation procedure. In addition, reducing

present spurious variability and inhomogeneity as a result

of temporal changes in spaceborne instrumentation and

inter-instrumental biases is a major challenge for future

improvements of reanalyses datasets. Data assimilation

for the stratosphere is the subject of ongoing development.

Improvements in the 3D-VAR and 40-VAR assimilation

procedures are currently being evaluated (Polavarapu et

al., 2005; Monge Sanz et al., 2006), which may lead to

improved reanalyses datasets. In the meantime, despite

its shortcomings, the ERA-40 dataset is the best description we have of the meteorological state of the stratosphere

over the past few decades (Randel et al., 2004). These

reanalyses have been used in chemical and dynamical attribution studies discussed below (Sections 3.4.2 and 3.4.5).

3.4.2 Dynamical Processes

As reviewed in Section 4.6 of WMO (2003),

changes in two specific dynamical transport processes can

significantly influence midlatitude ozone trends. These

are: 

(1) interannual and long-term changes in the strength of

the stratospheric mean meridional (Brewer-Dobson)

circulation, which is responsible for the winter-spring

buildup of extratropical ozone (e.g., Fusco and Salby,

1999; Randel et al., 2002; Weber et al., 2003; Salby

and Callaghan, 2004; Hood and Soukharev, 2005);

and

(2) changes in tropospheric circulation, particularly

changes in the frequency of local nonlinear synoptic

wave forcing events, which lead to the formation of

extreme ozone minima (“mini-holes”) and associated

large increases in tropopause height (Steinbrecht et

al., 1998; Hood et al., 1997, 1999, 2001; Reid et al.,

2000; Orsolini and Limpasuvan, 2001; Brönnimann

and Hood, 2003; Hood and Soukharev, 2005; Koch et

al., 2005).

It is therefore important to consider interannual

changes in both the Brewer-Dobson circulation and the

nonlinear synoptic wave forcing when estimating the component of interannual ozone variability and trends that can

be attributed to dynamical transport processes. The bulk

of the studies on this subject have looked at the northern

midlatitudes, and this is reflected in the following discussion. The lack of published studies for the Southern

Hemisphere is partly due to the apparently weaker signal

of dynamical changes on ozone trends compared with the

Northern Hemisphere.

When the stratospheric polar vortex is strong (positive North Atlantic Oscillation or Arctic Oscillation index),

tropospheric wave forcing is weaker, the Brewer-Dobson

circulation is weaker, and less ozone is transported to the

extratropics in winter and spring. Also, when the polar

vortex is strong, the zonal wind field in the midlatitude

lower stratosphere is less cyclonic, implying a greater frequency of anticyclonic, poleward wave breaking events

that lead to ozone mini-holes and localized tropopause

height increases (Peters and Waugh, 1996; Hood et al.,

1999; Orsolini and Limpasuvan, 2001). Therefore, at

northern midlatitudes in winter-spring, these two dynamical transport mechanisms tend to reinforce one another.

GLOBAL OZONE

3.22
During a period of increasing AO and NAO indices, such

as the 1980s and early 1990s (Appenzeller et al., 2000;

Hurrell, 1995; Graf et al., 1998; Zhou et al., 2001), a negative dynamically induced contribution to column ozone

trends at northern midlatitudes is to be expected. As

assessed below, several studies aimed at estimating empirically the contribution to ozone trends at northern midlatitudes due to one or both of the above transport mechanisms have been published during the past four years.

Reinsel et al. (2005) reported a multiple regression

analysis of the version 7 TOMS and SBUV/SBUV(/2) total

ozone dataset over the 1979 to 2002 period. The statistical model included terms proportional to the EliassenPalm (EP) planetary wave flux averaged between 30° and

90° latitude in each hemisphere as well as the

Arctic/Antarctic Oscillation (AO/AAO) indices. It was

found that both dynamical variables had a substantial influence on total ozone at latitudes higher than ~40° in both

hemispheres. Evidence was also obtained for a “large positive and significant” change in trend after 1996. The latter

result was obtained both with and without the inclusion of

dynamical index terms in the regression model. For the

Southern Hemisphere, Malanca et al. (2005) found a sizable, latitude-dependent slowdown in the ozone loss from

the early 1990s, with significant longitudinal variations in

the size of the change in trends. Although the zonal mean

behavior is strongly linked to variations in chemical loss

in the Antarctic vortex, the longitudinal asymmetry is due

to dynamical influences.

Inclusion of AO/AAO or NAO indices alone in a

regression model does not necessarily account completely

for ozone variability associated with synoptic wave

forcing. Also, the EP flux is more a measure of the ozone

tendency associated with changes in the Brewer-Dobson

circulation (e.g., Fusco and Salby, 1999; Randel et al.,

2002) than of ozone itself. So this statistical analysis could

not definitively determine whether or not the change in

trend after 1996 was caused by dynamical transport contributions. Nevertheless, the statistical evidence for a

change in trend after 1996 is not affected by these issues.

A similar statistical analysis of total ozone from groundbased observations in Europe by Krzy ́scin et al. (2005) has

also concluded that a positive change in ozone trend in

this region has occurred since 1994.

Continuing the earlier work of Fusco and Salby

(1999), Salby and Callaghan (2002, 2004) showed that

total ozone interannual variability and trends in the

Northern Hemisphere are both characterized by an out-ofphase relationship between low latitudes and high latitudes. This relationship is consistent with that expected

from changes in the Brewer-Dobson circulation. It was

therefore argued that a “systematic weakening” of the

latter transport influence was responsible for a major portion of the ozone decline between the 1980s and 1990s.

Randel et al. (2002) also used correlative and regression

relationships to estimate more specifically that net

decreases in strength of the Brewer-Dobson circulation

may have caused 20% to 30% of the column ozone trend

at northern midlatitudes over the 1979 to 2000 period.

The importance of nonlinear synoptic wave forcing

for ozone interannual variability and trends at northern

midlatitudes has been indicated by several recent studies.

Wohltmann et al. (2005) analyzed total ozone variability

measured at European ground-based Dobson stations

using a multiple regression statistical model that included

an explanatory variable based on the equivalent latitude

(i.e., potential vorticity or PV) profile at a given station.

Their technique takes advantage of the near conservation

of both ozone and PV on isentropic surfaces on a time scale

less than a few weeks in order to transform the PV profile

into a synthetic ozone profile using an ozone climatology.

The resulting synthetic ozone column is then used as the

explanatory variable. It accounts for both vertical lifting

(sinking) of isentropes by tropospheric pressure systems

and for horizontal isentropic advection of ozone. It was

concluded that 30% to 50% of the observed long-term

trend over Europe during the 1970 to 2002 period was

attributable to long-term changes in tropospheric pressure

systems. The latter consist primarily of increases in the

number and amplitudes of anticyclonic systems associated with poleward planetary and synoptic wave events

(e.g., Hood et al., 1999). Consistent with their results,

Brönnimann and Hood (2003) analyzed historical data to

show that low-ozone events over northwestern Europe in

winter were much more frequent in 1990-2000 than in

1952-1963 and that changing atmospheric circulation

strongly contributed to the observed increase in frequency.

In an analysis of balloonborne ozonesonde data at several

Canadian stations, Tarasick et al. (2005) report evidence

for a positive change in ozone trends at all levels below

63 hPa after about 1993. Interannual variability was found

to correlate well with the wintertime frequency of laminae

in the ozone profile (defined here as a thin layer bounded

by sudden changes in the ozone profile of at least 20 hPa),

which represent southward excursions of ozone-rich

Arctic, or polar vortex, air. It was suggested that much of

the positive change in trends after the early 1990s could

result from changes in circulation, since both the laminae

time series and the ozone time series show a similar change

in trend (see also Krizan and Lastovicka, 2005).

As also reviewed in the previous Assessment

(Section 4.6.3.3 of WMO, 2003), it may be reasonably

questioned whether the separate influences of the BrewerDobson circulation and tropospheric synoptic wave

GLOBAL OZONE

3.23
forcing on zonal mean ozone trends at a given latitude can

simply be added together, because these two dynamical

processes may not be entirely independent of one another.

For example, tropospheric planetary-scale waves that are

dominantly responsible for driving the Brewer-Dobson

circulation are also associated with synoptic wave events

and local tropopause height changes.

Hood and Soukharev (2005) used correlative and

regression methods to estimate separately and in combination the portion of ozone interannual variability and

trends over the 1979 to 2002 period that can be attributed

to long-term changes in both the Brewer-Dobson circulation and nonlinear synoptic wave forcing. In approximate

agreement with Randel et al. (2002), it was estimated that

18% to 25% of the observed maximum negative trend in

February and March is due to long-term changes in the

Brewer-Dobson circulation. In addition, 27% to 31% of

the observed maximum midlatitude trend was estimated

to be caused by synoptic wave forcing. No significant

correlations were found between monthly mean EP flux

variations (representing changes in Brewer-Dobson circulation strength) and monthly mean PV variations (representing changes in synoptic wave forcing) at northern

midlatitudes. A significantly higher correlation, up to 0.7,

was obtained using both the time-integrated EP flux and

zonal mean PV as predictor variables than using either

predictor variable alone. This indicates that contributions

to total ozone variability from these processes are, at least

to first order, independent and summable at northern midlatitudes. Together, these transport components could

explain almost 50% of the observed interannual variance

and maximum negative trend at northern midlatitudes in

February and March (Figure 3-16). The empirical regression model was able to simulate approximately the leveling off and slight increase in column ozone anomalies

that have been observed for some months and latitudes

since the middle 1990s (Figure 3-16). Furthermore,

Brönnimann et al. (2004a, b) provided evidence for

dynamic coupling of strong and long-lasting El NiñoSouthern Oscillation events with stratospheric ozone in

the NH by analyzing the record high total ozone values

observed in northern midlatitudes in the early 1940s.

Changes in dynamical processes also affect the

polar vortex conditions and, as a result, polar ozone loss.

The midlatitude ozone is influenced by polar loss via airmass mixing after the polar vortex breakup in early spring,

as discussed in Section 3.4.3.2. Using regression analysis,

Dhomse et al. (2006) concluded that this mechanism is

one of the main factors responsible for the recent increase

in NH total ozone.

Finally, estimates of the dynamically induced contributions to ozone interannual variability and trends can

be derived by using CTMs driven by observed temperature and wind fields (Hadjinicolaou et al., 1997, 2002,

2005). Most recently, Hadjinicolaou et al. (2005) have

reported further integrations of the SLIMCAT 3-D chemical transport model with a parameterized ozone tracer and

using the ECMWF ERA-40 meteorological analyses for

1979-2002. They find that the model simulates a large

part of the observed ozone changes at northern midlatitudes (represented by the merged satellite dataset (see the

Appendix 3A.1) averaged over the 35°N to 60°N latitude

range), including the positive tendency after the mid1990s. A linear regression analysis indicated that about

one-third of the observed ozone trend from 1979 to 1993

at these latitudes and, within uncertainties, all of the positive trend thereafter could be explained by the model,

which includes only transport-related changes. They argue

that these changes are consistent with the changes in the

Brewer-Dobson circulation in the observed positive tendency after the mid-1990s, but emphasized that this transGLOBAL OZONE

3.24

Ozone DU Ozone DU (a)

(b)

79 89 8 83 85 87 9 93 95 97 99 01

420

420

400

400

340

360

380

380

360

340

320

79 81 83 85 87 89 9 93 95 97 99 0

Ozone February, 45°N

r = 0.67

Ozone March, 45°N

Observed

Pred_HFLX_PV

Observed

Pred_HFLX_PV

r = 0.65

Figure 3-16. (a) Comparison of observed and empirically estimated column ozone at 45°N for February

when both 350 K PV and EP flux, averaged over the

40°N to 70°N latitude band and integrated over January

and February, are considered as predictor variables in

the empirical model. (b) As in (a) but for March. In this

case, EP flux is integrated over the January to March

period. From Hood and Soukharev (2005).
port process alone is unlikely to explain the relatively large

change in slope of the ozone trend (modeled and observed)

and that synoptic wave forcing may contribute significantly. Results from this study are included with full

chemistry CTM studies, which also quantify the important role for halogen forcing, in the discussion of assessment models (Section 3.4.5 below).

The direct effects of dynamical factors on ozone

described above are mainly limited to winter and spring.

Through the summer season the total ozone changes from

month to month are mainly controlled by nitrogen oxides

(NOx) abundance and photochemistry (Brühl et al., 1998).

While the photochemical lifetime does not vary from year

to year (e.g., Randel et al., 2002), the initial conditions for

summertime ozone vary and depend on the winter-spring

ozone abundance. As a result, ozone anomalies observed

in late summer and early autumn are highly correlated with

winter-spring anomalies (Fioletov and Shepherd, 2003;

Dhomse et al., 2006). This suggests that summertime

ozone is also indirectly influenced by dynamical processes

in the preceding winter-spring.

In summary, both empirical and model studies using

observed circulation changes continue to indicate that a

major fraction (almost 50% in February-March) of the

observed column ozone trends at northern midlatitudes

from 1979 to the mid-1990s can be attributed to long-term

changes in lower stratospheric circulation. Such circulation changes may also have been responsible, at least in

part, for the increase that has been observed at these latitudes since the mid-1990s. The primary dynamical transport processes are (1) interannual and long-term changes

in the strength of the mean meridional (Brewer-Dobson)

circulation and (2) interannual and long-term changes in

nonlinear synoptic wave forcing. Contributions from these

two forcings appear to be comparable in amplitude and

are, at least to first order, independent of one another.

Zonal mean ozone decreases at northern midlatitudes in

winter-spring caused by both forcings are larger when the

tropospheric AO/NAO circulation indices are positive.

The causes of the observed long-term trends in lower stratospheric circulation (including trends in the AO/NAO

indices) are unclear. This question can be answered ultimately only with the aid of GCMs. Such models have so

far indicated that it is unlikely that radiative changes from

chemically induced ozone losses are responsible (e.g.,

Langematz et al., 2003). Remaining possibilities include

long-term natural climate variability and greenhouse-gasinduced climate change.

3.4.3 Chemical Processes

3.4.3.1 UPDATE OF RELEVANT KINETICS

Based on the Jet Propulsion Laboratory (JPL) 2002

recommendation (Sander et al., 2003), updates to stratospheric reaction kinetics since the previous Assessment

have generally been minor. No major new reaction pathways or rate changes have been uncovered since WMO

(2003), which relied primarily on the JPL 2000 evaluation

(Sander et al., 2000). Several open issues from the previous Assessment have been resolved and some relatively

minor updates are discussed here. Note that here we do

not discuss any results from the very recent JPL 2006 recommendation (Sander et al., 2006), which appeared after

this Assessment was finalized.

Gas Phase

The JPL 2002 recommendation (Sander et al.,

2003) highlights several noteworthy changes from the previous stratospheric evaluation. Rates of some hydrocarbon

reactions in the upper troposphere are considered, although

their effect on the lower stratosphere is probably small.

Reactions with hydroxyl radical (OH) and photolysis cross

sections and quantum yields have been updated and added

for numerous halocarbons. These will affect halocarbon

lifetimes to some degree but do not have a major impact

on stratospheric ozone.

Progress has been made in resolving several

kinetics issues that were raised in WMO (2003).

Regarding the product pernitrous acid (HOONO) in the

reaction OH + NO2 + M, JPL 2002 recommends neglecting

this channel in stratospheric models. The OH + ClO →

HCl product channel has increased from 5 to 7% at 298 K

in JPL 2002, and the HO2 + ClO updates discussed in

WMO (2003) are included in JPL 2002, however, the

impacts of these changes in models have not been published specifically. Peroxynitric acid (HO2NO2) overtone

photolysis has been tested in comparison with aircraft and

balloon data by Salawitch et al. (2002) and in a global

model by Evans et al. (2003). This process significantly

changes the model abundance of HO2NO2 globally and

odd hydrogen (HOx) particularly near sunrise/sunset, and

improves comparison with observations. The maximum

change of O3 in the lower stratosphere is 1-2%. This

process is not included JPL 2002. Data on HOx reactions

with O3 discussed in WMO (2003) are now included in

JPL 2002, which is consistent with atmospheric measurements of OH/HOx.

GLOBAL OZONE

3.25
More recently, the reaction NO + HO2 → OH + NO2

(a) has been found to have a second channel at low temperatures: NO + HO2 → HNO3 (b), branching ratio (kb/ka)

= 0.18 (298 K) and 0.87 (223 K) (Butkovskaya et al.,

2005). Inclusion of this product channel will produce a

HOx sink in the upper troposphere and decrease

NOx/HNO3 partitioning in the lower stratosphere, perhaps

by as much as 10-15%. Global impacts have yet to be

tested in models.

Low-temperature oxidation rates for several hydrocarbon species have been updated, including: CH3CHO +

OH → CH3CO + H2O (Sivakumaran and Crowley, 2003),

H2CO + OH (Sivakumaran et al., 2003), and CH3OH +

OH → products (Dillon et al., 2005). These changes will

affect upper troposphere calculations but are not expected

to have a significant impact on stratospheric chemistry.

The assessment models discussed in Section 3.4.5

use an updated version of the JPL 2002 kinetics. No significant differences from simulations using the standard

JPL 2002 are found.

Heterogeneous Chemistry

JPL 2002 provides updates to several heterogeneous processes, with particular focus on liquid binary

H2SO4

.

H2O uptake, i.e., HOCl + HBr, HO2NO2 + HCl,

and α (HOI), as well as HCl + HNO3 on liquid ternary

solution. None of these updates is expected to have a significant impact on global ozone, although a specific test

has not been published.

Recent laboratory experiments yielded new information regarding uptake of gases on ice surfaces under

atmospheric conditions corresponding to the upper troposphere/lower stratosphere (UT/LS). So far the knowledge

of gas uptake on ice particles was primarily based on

clouds in cold polar stratospheric conditions and no information was available for competitive and reactive uptake.

Hynes et al. (2002) and Cox et al. (2005) found that the

uptake of HCl was suppressed when HNO3 was present

on the ice surface at temperatures higher than 208 K

(Hynes et al.) or 218 K (Cox et al.). Hence for temperatures greater than 208 K, there is a tendency to reduce

uptake of hydrochloric acid (HCl) when competitive

uptake is considered, which will reduce the potential for

chlorine activation on cirrus clouds. Fernandez et al.

(2005) found significant uptake of chlorine nitrate

(ClONO2) on pure and doped (HNO3, HCl) ice surfaces.

No model study has included this information to estimate

the effect of heterogeneous chemistry on aerosol and ice

particles on ozone.

3.4.3.2 LOWER STRATOSPHERE

The major chemical processes expected to contribute to global ozone depletion were reviewed in previous Assessments (e.g., WMO 2003). There have been

no significant changes to the understanding of these

processes, although studies have provided new quantification. There have been relatively minor refinements in

chemical rate constants and uncertainties (Section 3.4.3.1)

but no major changes in our understanding of chemical

processes. However, some new observations, related to

likely minor processes of cirrus activation and pyroconvection, are discussed here. Then we discuss recent findings related to the more established and important process

of aerosol chemistry, the role of bromine in the lower stratosphere, and export from the polar vortex.

Cirrus (Including Subvisible Cirrus)

In addition to the observed cirrus climatologies in

the tropopause region reported earlier (see WMO, 2003),

there is new observational evidence of thin cirrus clouds

in the cold tropical (Peter et al., 2003) and midlatitude

tropopause region (Ström et al., 2003; Keckhut et al.,

2005a). Supporting these findings, there is further observational evidence of supersaturation with respect to ice in

the same altitude range (Ovarlez et al., 2002; Spichtinger

et al., 2004), though it is likely that the thin “ice” clouds

consist of nitric acid trihydrate (NAT) (Popp et al., 2006).

Gao et al. (2004) found observational evidence for a new

HNO3-containing ice particle (delta-ice). However, the

laboratory results from Delval and Rossi (2005) do not

support the Gao et al. findings and suggest that small NAT

particles, even H2O-rich, can survive much longer in a

subsaturated atmosphere than previously thought.

Following the first observations of possible chlorine activation in the lowermost stratosphere (Keim et al.,

1996; Borrmann et al., 1997), Solomon et al. (1997)

showed that these findings could lead to significant chemical ozone loss in the midlatitude lowermost stratosphere,

but Bregman et al. (2002) calculated considerably less loss

(see WMO, 2003). More recent observations by Thornton

et al. (2003) showed much more in-situ observational evidence of chlorine monoxide (ClO) of several tens of parts

per trillion by volume (pptv) in the lowermost stratosphere

in the northern extratropics. The observed mixing ratios

are significantly higher than can be explained by background gas-phase and heterogeneous chemistry on

aerosols (Thornton et al., 2005). The observations were

primarily performed during wintertime in the Northern

Hemisphere and the enhancements were found between

55°N-70°N. These observed levels of ClO were in very

GLOBAL OZONE

3.26
close agreement with those calculated by Bregman et al.

(2002), although no evaluation could be made regarding

the temporal and spatial distribution of the ClO enhancements. However, the modeled bromine monoxide (BrO)

levels in Bregman et al. were 2-4 pptv, which is a factor of

two too small, according to recent estimates from

Salawitch et al. (2005), implying that the calculated ozone

loss is likely too small; consequently, this process may

need further study.

The chemical composition of the ice cloud particles is important for their potential to activate halogens.

In the tropical tropopause region there is ongoing discussion about the existence and occurrence frequency of the

HNO3-containing particles (Gao et al., 2004; Popp et al.,

2006; Delval and Rossi., 2005).

Pyroconvection

It is well known that the presence of enhanced

aerosols in the stratosphere can cause significant chemical ozone loss through heterogeneous chemical reactions,

as demonstrated by many studies of the impact of the

eruption of Mt. Pinatubo (see WMO, 2003). Apart from

volcanic eruptions, a recently identified process, pyroconvection, may also cause enhanced stratospheric

aerosol concentrations. Pyroconvection is induced by

boreal fires or biomass burning, likely in combination

with strong convective activity (Fromm and Servranckx,

2003; Jost et al., 2004; Fromm et al., 2005). The aerosol

enhancements have been found in the lower stratosphere

over all longitudes in the northern mid- and high latitudes.

The particles likely consist of soot and smoke (Fromm

and Servranckx, 2003).

Blumenstock et al. (2006) observed chlorine activation in the lower stratosphere in late Arctic winter under

conditions well above the NAT temperatures. This was

attributed to heterogeneous chemistry on an enhanced

aerosol plume, originating from forest fires and injected

by strong convection, as described in Gerding et al. (2003).

Ray et al. (2004) and Jost et al. (2004) attribute

enhanced carbon monoxide (CO) and aerosol observed by

aircraft measurements in the subtropical stratosphere to

pyroconvection. Livesey et al. (2004) found enhanced

acetonitrile (CH3CN) in stratospheric data from the spaceborne Microwave Limb Sounder (MLS), associated with

forest fires and thunderstorm lofting. Injection of aerosol

and other fire-produced chemical species will affect ozone

locally.

The occurrence frequency of pyroconvection

remains an open issue, as does a quantification of its

impact on ozone. However, given the impact of other

established processes included in models that can generally reproduce past ozone changes, we can surmise that

any effect will be relatively small.

Aerosol Effects 

Sulfate aerosol in the lower stratosphere provides

surfaces for the activation of chlorine. The distribution of

sulfate surface area depends on background sulfur emissions and volcanic eruptions. The atmosphere is currently

(since about 1999) near a background minimum (Section

3.3.1) and this sets the minimum for the nonvolcanic

aerosol loading. Thus, the Mt. Pinatubo-to-present period

gives a good span of potential heterogeneous effects, barring a huge future volcanic eruption. The reactions, as

reviewed in previous Assessments, do not suggest significant missing processes or inaccurate rates. Sulfate surface area in models can be, and usually is, specified from

observations for past-to-present runs.

A large chemical effect from volcanoes has been

quantified in 2-D models (e.g., Tie et al., 1994; WMO

2003, Section 4.5.3.4) and 3-D models (e.g., Chipperfield,

1999; 2003; Stolarski et al., 2006). The volcanic effect on

column ozone results from heterogeneous suppression of

NOx, via N2O5 + H2O(l), which then interferes less with

the halogen and HOx ozone loss cycles in the lower stratosphere. Thus, chemical ozone losses from volcanic sulfate injection are largest at times of peak chlorine and

bromine, and volcanic impact on ozone at preindustrial

halogen levels is small or even positive (Tie and Brasseur,

1995).

Dynamical changes resulting from the Mt. Pinatubo

eruption also contribute to ozone change and are present

in meteorological analyses (Hadjinicolaou et al., 2005; see

Sections 3.4.2, 3.4.5). An outstanding puzzle is the lack

of a Pinatubo effect on observed ozone in the SH while

the effect on NO2 there is clear. Models all show a SH

effect as large, or larger than, the NH effect. Stolarski et

al. (2006) show that much of this apparent hemispheric

discrepancy may be the result of interannual variability

masking the volcanic effect in the ozone time series

analysis.

Inorganic Bromine

A variety of observations have shown that that inorganic bromine in the stratosphere is underestimated when

based on release from methyl bromide (CH3Br), halons,

and other long-lived source gases (see Chapter 2), as is

usually assumed in global models. Although full resolution of this issue awaits further study (Sinnhuber et al.,

2005), the potential impact of increased background stratospheric Br on O3 loss is significant (Salawitch et al.,

GLOBAL OZONE

3.27
2005). Assuming the purported additional Br is of biogenic origin, and thus is constant with time, it contributes

slightly to the O3 column change from 1979 to present

through reaction with Cl at low background aerosol

amounts. The effect of the additional Br is large, however,

during periods of volcanically enhanced stratospheric

aerosol following El Chichón and Mt. Pinatubo (see

Figures 3-14 and 3-25). Ozone loss via the BrO + ClO

catalytic cycle is enhanced by a factor of two or more by

the additional Br during periods of relatively high Cl

resulting from heterogeneous reactions on volcanic sulfate. This enhancement generally improves the global

model simulation of decadal O3 change in comparison to

observations (Chapter 2 Figure 2-12; see also 3-D results

in Figure 3-25 below). This potential additional source of

stratospheric Br is important to resolve and characterize

because it will impact O3 future projections as Cl decreases

and if the Br source and/or transport processes change with

climate. 

Export from Vortex 

The export of ozone-depleted or activated polar

vortex air may have a significant contribution to observed

ozone loss at midlatitudes in spring-summer. This effect

exists in both hemispheres but is expected to be larger in

the Southern Hemisphere due to the larger and more regular ozone depletion in the Antarctic vortex. Although

transport is clearly involved in this process, the ultimate

cause is chemical O3 loss in the polar regions by Cl and

Br species (see Chapter 4). Figure 3-17 shows that the

mass of “missing” ozone in the ozone hole has the same

order of magnitude as the mass deficit over southern

middle and high latitudes in summer, and illustrates the

strong correlation between loss in the ozone hole and in

summer.

Konopka et al. (2003) investigated the fate of Arctic

vortex remnants, and the chemistry occurring in those remnants, during the spring of 1997 and 2000 using a

Lagrangian CTM. They found a different behavior in the

lower and in the mid-stratosphere. Above 20 km, vortex

remnants (Orsolini, 2001) remain long-lived in the

summer westward circulation. Using balloonborne in-situ

measurements of water vapor and methane, Durry and

Hauchecorne (2005) indeed found evidence for such

vortex remnants in the midlatitude summer stratosphere.

Below 20 km, the subtropical jet bounds the meridional

propagation of those remnants (Piani et al., 2002), and the

remnants’ lifetime is considerably reduced due to

enhanced stirring by synoptic eddies. This picture is corroborated by satellite measurements of the breakup of the

Antarctic polar vortex in 2004. Manney et al. (2005)

reported that vortex remnants persisted in tracer measurements for over a month in the mid-stratosphere, but no

more than a week in the lower stratosphere.

Seasonal model studies have been conducted to

quantify the impact of the export of polar air on midlatitudes in particular years. Knudsen and Grooß (2000)

found that about 40% of the observed midlatitude depletion in May 1995 and 1997 is due to dilution with polarprocessed air. Including all years with large Arctic ozone

depletion (1993, 1995, 1996, 1997, and 2000) Andersen

and Knudsen (2002) found that the dilution explains 33%

of the 1979-2002 midlatitude trend in the April-May

period. The dilution has a tendency to occur mainly over

the Eurasian continent, which helps to explain why ozone

trends are larger over this region (Knudsen and Andersen,

2001). Using a high resolution CTM, Marchand et al.

(2003) estimated that the dilution of polar vortex air during

the cold winter 1999-2000 contributed up to 50% of ozone

depletion at northern midlatitudes (45°-55°N) over the

period from December 10 to April 30. Millard et al. (2003)

studied the connection between polar and midlatitude

ozone loss with a series of seasonal 3-D CTM simulations.

During years 1994 to 2000, they found a large interannual

variability (from −25 DU to almost 0) in the contribution

of polar vortex air (90°-70°N) to midlatitude (60°-30°N)

ozone loss from December 1 to May 31.

For the Southern Hemisphere, Ajti ́c et al. (2004)

quantified the seasonal dilution effect of the Antarctic

ozone hole by calculating ensemble diabatic trajectories

GLOBAL OZONE

3.28

O3 Mass Difference from 79 81 Mean (x109 kg)

ON, 60°S–90° S

D F, 30°S–60° S

D JF, 60°S–90° S

Figure 3-17. The change in mean ozone mass from

the average of 1979-1981 over southern polar latitudes in spring (September-November; blue curve)

and summer (December-February; red curve) compared with that over southern midlatitudes in summer

(green curve). Derived from the NIWA Assimilated

Dataset (see Appendix 3A).
of ozone-depleted air parcels from October 15 to January

15 for a series of years. The mean calculated ozone reduction in 1998, 1999, and 2000 was 16-19 DU (5-6% of total

column) between 30°S-60°S. This is 83-95% of the

change observed by TOMS between 1979-1980 and 1998-

2000. Part of this overestimate is due to the fact that photochemistry was not included. Based on Prather et al.

(1990), who suggested there would be 20-30% production

during this period, the dilution may therefore account for

58-76% of the summertime ozone depletion at southern

midlatitudes

Multiannual model studies have been performed to

investigate the accumulated effect of polar processing on

midlatitudes. Hadjinicolaou and Pyle (2004) investigated

the dilution of polar ozone loss in the 1990s using a decadal

3-D CTM run with parameterized ozone. They found a

large interannual variability in the north related to winterspring planetary wave activity and calculated a year-round

midlatitude depletion of about 1% between 40°N-60°N,

compared with 2-4% in the south. Chipperfield (2003)

used a full chemistry 3-D CTM for 1979-1995 and

reported that for the overall modeled decrease of midlatitude ozone due to halogens, about 30% (NH) to 50% (SH)

of the change is caused by ozone loss within the polar

vortex.

Once the vortex breaks down, ozone is rapidly

mixed throughout the extratropics, which makes summertime year-to-year ozone variations outside the tropics

homogeneous in latitude. This holds for both short-term

variations and long-term trends (Fioletov and Shepherd,

2005). Figure 3-18 (top) shows that the magnitudes of

the summertime trends over middle and polar latitudes

are nearly identical, while the corresponding springtime

trends are quite different. Polar ozone depletion and

export of depleted air from the vortex thus affects the seasonal structure of midlatitude ozone trends, and thereby

accounts for the different seasonality of the midlatitude

trends in the two hemispheres. Over 35°-60°N, the longGLOBAL OZONE

3.29

5°N–80°N 35°S–80°S

5°S–80°S

5°S–60°S

0°S–80°S

5°N–80°N

5°N–60°N

0°N–80°N

Figure 3-18. Top: The observed total ozone trends for 35°-60° (red), 60°-80° (blue), and 35°-80° (black) latitudinal belts for the Northern Hemisphere (left) and Southern Hemisphere (right). Bottom: The observed ozone

trends for 35°-80°N and S (black line), and the trends estimated from the March, April, and May trends for the

NH (November, December, and January trends for the SH) and regression coefficients estimated from

detrended data. The trends were estimated for the period 1979-2003 using regression to an EESC curve and

converted to DU/year using the variation of EESC with time in the 1980s. The error bars represent the 95%

confidence intervals. From Fioletov and Shepherd (2005).
term trends are in line with interannual variability, with a

maximum in spring and a minimum in autumn. In particular, the trend magnitudes from late spring through to

early autumn are related to the trend in April in the same

way that the corresponding monthly ozone anomalies are

related to the April anomaly in the detrended data

(Fioletov and Shepherd, 2003). In contrast, the summer

ozone long-term decline over 35°-60°S is stronger than

one would expect purely from photochemical decay of

the spring trend there, and is in fact of comparable magnitude to the spring trend. However for the entire 35°-

80° belt, the long-term trends in SH ozone have a clear

seasonal structure and are in line with interannual variability, as shown in Figure 3-18 (bottom). The implication is that the near constancy of the SH midlatitude trends

throughout the year, in striking contrast to the behavior in

the NH, reflects the much greater influence of polar

processes in the SH compared with the NH. Indeed,

Fioletov and Shepherd (2005) estimated from total ozone

observations that about 39% of the observed southern

midlatitude long-term ozone decline in December can be

attributed to the polar ozone depletion achieved up to

November. In the Northern Hemisphere, the corresponding contribution is about 15%, but the statistical uncertainty is too large to make an accurate estimate.

3.4.3.3 UPPER STRATOSPHERE

There have been no changes in our understanding

that ozone trends in the upper stratosphere (around 40 km)

are caused by changes in chlorine acting through the ClO

+ O cycle. The chemistry of the upper stratosphere will

be affected by the trend in temperature (Chapter 5) and the

partitioning of ClO and HCl, which will affect loss via the

ClO + O cycle, will also be affected by changes in methane

(CH4) (Chapter 1). Trends in these quantities also need to

be observed and compared with models for a rigorous test

of our understanding. Overall, our understanding, as

expressed in current models, is qualitatively consistent

with observations, though as discussed in Section 3.4.5,

there appears to be quantitative discrepancies in the simultaneous modeling of all relevant trends. This, however,

should be considered as a problem for these models to simulate realistically the distribution of the precursors and

species that control ozone, rather than a measure of uncertainty in the upper stratosphere chemical processes for

which we have a good understanding (e.g., WMO, 1999).

Chapter 6 (Section 6.5.1) summarizes observations of

recent changes in ozone in the upper stratosphere, in the

context of EESC levels, and discusses the implications for

recovery from halogen-induced loss.

3.4.4 Solar Cycle Variations 

A decadal variation of total column ozone is

observed to be in phase with the solar cycle, with an amplitude of 2-3% from solar minimum to maximum in the

tropics (Figure 3-4) and over the 60°S-60°N band (Figure

3-1). The detailed vertical structure of the ozone variation

is still a topic of current research (for a review, see Hood,

2004) but, despite some uncertainty due to the limited

record length (see discussion below), this decadal variation appears to be a dominant form of long-term ozone

variability and should be carefully considered when evaluating anthropogenic trends at all latitudes and altitudes

(Steinbrecht et al., 2004a, b; Cunnold et al., 2004).

The stratospheric ozone response to 11-year solar

forcing has been estimated as a function of altitude, latitude, and season by a number of analysts based mainly on

long-term, near-global satellite remote sensing datasets

(Chandra and McPeters, 1994; McCormack and Hood,

1996; McCormack et al., 1997; Hood, 1997; Wang et al.,

1996; Lee and Smith, 2003; Soukharev and Hood, 2006).

Figure 3-19 compares the annually averaged solar cycle

response derived from two long-term satellite ozone profile datasets (SBUV and SAGE) with the predictions of

four models that account for observed 11-year changes in

solar ultraviolet spectral irradiance. The observed ozone

change from solar minimum to maximum is estimated

using a multiple regression statistical model containing

QBO, volcanic aerosol, solar cycle, and linear trend

explanatory variables as well as a first-order autoregressive term. Figure 3-19a presents the signal in ozone averaged over the 55°S-55°N band (essentially the whole

region covered by the satellites), while Figure 3-19b shows

only the tropical response (25°S-25°N). The satellite

datasets used are SBUV(/2) version 8 from 1979-2003

(Frith et al., 2004; updated from Hood, 2004) and SAGE

I+II version 6.20 from 1979 to 2005, excluding several

years after the Mt. Pinatubo eruption (updated from

Stolarski and Randel, 1998). Analyses of SAGE II and

UARS HALOE data (Soukharev and Hood, 2006), as well

as comparisons between SBUV and TOMS data (e.g.,

Hood, 1997), suggest that most or all of the apparent total

ozone solar cycle variation originates primarily in the

lower stratosphere. Thus the observed variation in the

ozone column of 2-3% (as discussed above) can be taken

to represent the integrated values below about 25 km and

30 km for the 55°S-55°N and 25°S-25°N bands.

The theoretical curves in Figure 3-19 include results

from two 2-D radiative-chemical transport models

(Brasseur, 1993; Haigh, 1994) and two fully interactive,

3-D Chemistry-Climate Models (Tourpali et al., 2003;

Egorova et al., 2004). Although the models have very difGLOBAL OZONE

3.30
ferent dynamical formulations, their predictions for the

ozone solar response in the low-to-middle stratosphere

(15-35 km) are very similar, and also representative of the

results from other published modeling studies. In the

upper stratosphere the model predictions diverge, probably due to their different representations of photochemical processes in the mesosphere.

Over the 55°S-55°N band (Figure 3-19a), the model

predictions lie within the error bars of the SAGE data

analysis, and of the SBUV analysis in the upper and lower

stratosphere, but tend to overestimate the signal found in

the middle stratosphere in the SBUV data. The models

are also generally consistent with the 1.5% lower stratospheric (ozone column) signal. However, in comparing

the predicted model ozone responses to the observationally derived responses in the tropics (Figure 3-19b), some

possible differences on the 11-year time scale are evident.

Near the stratopause (~50 km altitude), the observed

response appears to be larger than expected from the

models while in the middle stratosphere (~30 – 35 km),

the reverse is true, especially relative to SBUV data. In

the lower stratosphere, the mean observed response

appears to increase again while the model responses

decrease. 

A number of possible explanations for the apparent

differences between the observationally derived and

model-predicted tropical ozone responses may be considered. First, it is possible that there are significant uncertainties in the observations because of problems in the

cross-calibration of instruments (see Section 3.2.3),

overall short record lengths, and the relative size of the

solar effect (a few percent). Only after another several

decades of data are acquired will these uncertainties be

reduced.

One physical explanation suggested for the unexpected altitude dependence of the tropical ozone response

is increased odd nitrogen in the upper stratosphere caused

by enhanced precipitation of energetic electrons when the

GLOBAL OZONE

3.31

Percent

Figure 3-19. Annual mean solar

cycle response calculated from two

long-term satellite ozone profile

datasets with the predictions of four

stratospheric models that account for

observed 11-year changes in solar

ultraviolet spectral irradiance, for (a)

55°S-55°N and (b) 25°S-25°N. Datasets are the monthly zonal mean version 8 SBUV/SBUV(/2) ozone profile

dataset of Frith et al. (2004) over the

1979 through 2003 period (updated

from Hood, 2004) and version 6.20

SAGE I and II ozone profile data over

the 1979 to 2005 period, except for

several years following the Mt.

Pinatubo eruption (updated from

Stolarski and Randel, 1998). Model

results are from the 3-D models of

Tourpali et al. (2003) and Egorova et

al. (2004) and the 2-D models of

Brasseur (1993) and Haigh (1994).
Sun is less active (Callis et al., 2000, 2001). Some model

simulations (Langematz et al., 2005; Rozanov et al., 2005)

have produced a response of opposite sign (a “self-healing

effect”) in the mid-stratosphere but only in response to

massive electron precipitation events producing very large

NOx signals, which have not been detected in satellite data

(Hood and Soukharev, 2006). At this time, therefore, there

is no evidence for particle-induced NOx being responsible

for the apparent difference between models and observations in the tropics.

Several possible explanations for the reduced solar

cycle ozone response in the tropical middle stratosphere

may be considered. First, it is possible that the multiple

regression statistical analysis is being biased by interference from the QBO and major volcanic eruptions. Lee and

Smith (2003) have applied a two-dimensional chemicaldynamical-radiative model with a prescribed QBO to

show that such interference can occur, especially from

the QBO, and can potentially lead to small or even negative solar cycle regression coefficients in the equatorial

lower and middle stratosphere. It remains to be seen

whether some of this response is an artifact due to the

highly idealized form used by Lee and Smith for the QBO,

but interference from the QBO should be carefully considered in future evaluations of the observed tropical

response minimum. It is also possible that the QBO itself

may be modulated slightly by the solar cycle (Salby and

Callaghan, 2000; McCormack, 2003; but see also Hamilton, 2002). If so, then additional dynamical effects on

the net ozone change over a solar cycle could be implied.

Finally, one relatively simple possible explanation for the

tropical ozone response minimum is that the unexpectedly large ozone response in the upper stratosphere effectively reduces the ozone production rate in the middle

stratosphere through increased absorption of UV radiation (a form of “self-healing”), but this remains to be

tested. On short time scales (i.e., that of the ~27-day solar

rotation period), the observed ozone response to solar UV

variations in the tropical middle and upper stratosphere

agrees very well with photochemical model predictions

(see, e.g., Chen et al., 1997).

The origin of the apparent solar cycle ozone variation in the lower stratosphere also remains uncertain. It

has been questioned whether the lower stratospheric

decadal ozone variation is of solar origin at all because of

the occurrence of two major volcanic perturbations about

9 years apart in 1982 and 1991 (e.g., Solomon et al., 1996).

However, the increase in tropical total ozone approaching

the most recent solar maximum, when no major eruptions

have occurred, supports the view that volcanic eruptions

alone cannot explain this decadal variation. A remaining

hypothesis is that the direct upper stratospheric effects of

solar ultraviolet radiation (and possibly particle precipitation) are able to modify the development of stratospheric

circulation in such a way as to modify the effective

upwelling rate in the tropical and subtropical lower stratosphere. Possible mechanisms for this, and for the subsequent transmission of a solar signal into the troposphere,

are discussed in Section 5.2.1.1 of Chapter 5.

3.4.5 Assessment Model Simulations

A range of model runs have been performed for this

Assessment to simulate both the past and future atmosphere. Different types of two-dimensional (2-D) latitudeheight models and three-dimensional (3-D) models are

described in Box 5-1 (Chapter 5). Simulations using 2-D

models and 3-D coupled Chemistry-Climate Models

(CCMs) are described in Chapter 6 (Appendix 6A), where

they are used to investigate future changes in stratospheric

ozone. These models were also used, in certain experiments, to simulate the past atmosphere. In this section we

make use of these simulations of the recent past to understand the causes of past changes in ozone. Descriptions of

the 2-D models and 3-D CCMs used in this chapter may

be found in Tables 6-3 and 6-4 of Chapter 6. In addition

to the 2-D models and CCMs, we also use results from offline 3-D chemical transport models (CTMs). Because the

CTM results are only used in this chapter, we summarize

details of the runs available in Table 3-1. Generally, for

the CTMs we have used existing, published simulations

and reanalyzed the output to compare with recent observations. The NASA Goddard Space Flight Center (GSFC)

CTM provided one simulation (Stolarski et al., 2006),

which was forced by GCM winds. The Global Modeling

Initiative (GMI) CTM, which is technically very similar

to the GSFC CTM and is based on much of the same code,

provided two simulations. These used repeating climatological GCM winds corresponding to a “cold” and a

“warm” Arctic winter. The SLIMCAT CTM was forced

by analyzed ECMWF winds. In one experiment the model

used a parameterized ozone tracer (see Hadjinicolaou et

al., 2005). The model was also used to perform three

experiments with full chemistry: A run with timedependent halogen loading, a similar run but without an

assumed 5 pptv of inorganic bromine (Bry) from shortlived species, and a run with fixed halogen loadings.

Two-dimensional models have been used extensively in previous Assessments and we use them again

here for comparison with the 3-D CTM runs. We have

used output from the P5 runs (see Chapter 6), as these

have the most realistic bromine loadings. In addition to

the basic model setups described in Appendix 6A, the

non-interactive GSFC 2-D model also performed a run in

GLOBAL OZONE

3.32
which the model was forced with an interannually varying

circulation (run GSFC-IDV in discussion below). These

2-D models and the 3-D CTMs have been compared with

observations for ozone column and profiles in both the

Northern and Southern Hemispheres, which extends on

published work. The status of CCMs is much less mature

and for these simulations we show just a basic comparison of the range of models with column ozone.

Figure 3-20 shows the deseasonalized zonal mean

column ozone for global (60°N-60°S), NH (60°N-35°N)

and SH (35°S-60°S) regions for a selection of the 2-D

models and 3-D CTMs. There is a wide range in the predicted absolute column O3 from the different models,

which is due mainly to differences in the model transport.

Clearly, the 2-D models, with their climatological circulation, do not capture the interannual variability in the observations. The GSFC CTM, forced by GCM winds, does

well in tracking the variations in global ozone (as shown

in Stolarski et al., 2006) but overestimates the SH values

by around 20 DU. The interannual variability in this

model, driven by the GCM winds, would not be expected

to correlate with observed variations. For the 2-D models

(except GSFC-IDV, not shown in Figure 3-20), the weak

interannual variations will be driven by only aerosol

changes and source gas trends. The SLIMCAT CTM,

which is forced by analyzed winds, does capture more variability. This was demonstrated for the NH by

Hadjinicolaou et al. (2005), who ran the model for the

stratosphere with parameterized O3 and showed the very

good agreement for 35°N-60°N. The full chemistry

SLIMCAT run extends to the surface and tends to have

larger columns due to overestimating O3 in the lowermost

stratosphere. However, these model runs are characterized by a large, spurious positive deviation in O3 in the

late 1980s that is caused by inaccurate interannual variability in the forcing (ERA-40) analyzed winds.

Figure 3-21 shows the global, equatorial, NH midlatitude, and SH midlatitude anomalies of the annual

average O3 from the 2-D models and 3-D CTMs. As seen

in WMO (2003), the 2-D models perform better in reproducing the observed variations in the NH, although there

is still a spread in model results. In particular, the decrease

in these models around 1992 is due to enhanced aerosol

from the Mt. Pinatubo eruption. In the SH and globally,

the 2-D models produce a wider spread of results, likely

reflecting different treatments of the Antarctic ozone hole.

The 3-D CTMs also appear more successful in simulating

the NH midlatitudes than the SH midlatitudes. Aside from

the large positive (late 1980s) and negative (early 1990s)

deviations in the SLIMCAT runs, which use ERA-40

winds, the CTMs perform well in capturing the overall

variations, which will be dominated by changes in aerosol

loading and halogens. The SLIMCAT parameterized O3

run is that discussed in Hadjinicolaou et al. (2005), where

they show results for 35°N-60°N (see Section 3.4.2). This

run has no time-dependence in the chemistry parameterization and clearly does not capture the overall trend from

1980-2004. However, the run does produce an increase in

GLOBAL OZONE

3.33

Table 3-1. Participating 3-D chemical transport models.

Model Name Institution(s) Investigators Forcing Winds Resolution Chemistry / Reference

and and Domain PSC Scheme

Temperatures

GSFC CTM NASA A. Douglass GEOS-4 GCM 2° × 2.5° Full chemistry / Stolarski

Goddard, S.R. Kawa Surface to Considine et al. (2006)

U.S. R. Stolarski 0.4 hPa et al. (2000)

GMI CTM Several A. Douglass GEOS-4 GCM 2° × 2.5° Full chemistry / Douglass

institutions, S. Strahan Surface to Considine et al. (2004)

U.S. R. Stolarski 0.015 hPa et al. (2000)

(33 levels)

SLIMCAT University of M.P. ECMWF 7.5° × 7.5° Full chemistry / Chipperfield

Leeds, U.K. Chipperfield (ERA-40, Surface to Equilibrium (1999)

W. Feng operational) 55 km Feng et al.

(2006)

University of P. Hadjinicolaou ECMWF 5.6° × 5.6° Parameterized Hadjinicolaou

Cambridge, J. A. Pyle (ERA-40, 8 to 55 km O3 tracer / (No et al. (2005)

U.K. operational) time dependence)
GLOBAL OZONE

3.34

Figure 3-20. Time series of deseasonalized mean column O3 for (top

panel) 60°S-60°N, (middle panel)

35°N-60°N, and (bottom panel) 60°S35°S from 1980-2004 for three

selected 2-D models and three 3-D

CTMs. Also shown are the observations from the merged satellite dataset (black line). See Tables 6-3 and

6-4 of Chapter 6 for information about

the models.

Figure 3-21. (On the right.) Time series

of annual mean column ozone variations from observations (black line) and

from eight 2-D model runs (left) and

five 3-D CTM runs (right) for latitude

bands of (top) 60°S-60°N, 35°N-60°N,

25°N-25°S, and (bottom) 60°S-35°S.

Changes are calculated in percent with

respect to 1980 values for each time

series. Observations are from the

merged satellite dataset; see Appendix

3A. The error bars on the observations

represent the spread of different

datasets, which provides a minimum

estimate of their uncertainty. The

SLIMCAT run with parameterized O3

does not include any chemical changes

and is therefore not expected to follow

the other model curves. See Tables 6-

3 and 6-4 of Chapter 6 for information

about the models.
GLOBAL OZONE

3.35
NH column O3 after the mid-1990s and this was presented

as evidence that circulation changes are leading to O3

increases. Interestingly, the GMI CTM runs with repeating

meteorology also model an increase in O3 during this

period, which, like any changes in the 2-D models except

GSFC-IDV, is related to aerosol trends. The GMI run with

“cold” Arctic meteorology gives around 1% lower O3 than

the run with “warm” meteorology, which is a measure of

the impact of different polar loss in this run. The different

CTM runs give similar simulations for the SH, again with

more variability in the SLIMCAT run, but the full chemistry CTMs clearly underestimate the observed O3 from

the early 1990s on. These models produce a realistically

large O3 hole and this appears to be having a large impact

on the model SH midlatitudes. The models shown here

are a mixture of those that include feedback between chemistry and dynamics (e.g., some of the 2-D models) and

models without any coupling (e.g., the off-line 3-D

CTMs). However, the impact of the modeled midlatitude

O3 loss on dynamics is not likely to be important. Using a

GCM with parameterized O3 chemistry, Braesicke and

Pyle (2003) showed the impact of an imposed midlatitude

O3 loss did not produce an additional feedback on the O3

distribution. The effect of dynamical changes on ozone is

discussed in Section 3.4.2.

Results from the 2-D models and 3-D CTMs have

been analyzed for ozone trends. This was done with a similar statistical model to that used in Section 3.2.1 and the

trend term was regressed onto an EESC curve. Figure 3-

22 shows the modeled O3 trends from 1980-2004 (fitted

using EESC and converted to %/decade using the EESCtime variation in the 1980s) as a function of latitude. As

shown in WMO (2003) the 2-D models tend to reproduce

the observed trends in the NH midlatitudes but show a

wide variation in the SH. The 2-D models also produce a

negative trend of 1-2%/decade in the tropics where the

observations indicate a zero trend. The 3-D CTMs also

fall in the range of observations in the NH and produce a

negative trend in the tropics with the exception of the

SLIMCAT run forced by analyzed winds. This run produces a zero trend in the tropics (see also Figure 3-21),

which may indicate that the ECMWF analyzed winds succeed in isolating this region more efficiently (see Monge

Sanz et al., 2006). In the SH, the CTMs are quite consistent with each other but again tend to overestimate the

observed midlatitude trend.

Figure 3-23 shows the modeled and observed

column ozone trends versus month for the NH and SH.

The NH observations show a seasonality with the largest

trend of around −4%/decade in March and a minimum

GLOBAL OZONE

3.36

Figure 3-22. Latitudinal profile

of annual mean column ozone

trends from observations (black

line) compared with results from

2-D models and 3-D CTMs for

1980-2004. The trends were

estimated using regression to

an EESC curve and converted

to %/decade using the variation

of EESC with time in the 1980s.

Observations are from the

merged satellite dataset; see

Appendix 3A. The error bars for

the observations indicate the

2σ uncertainties.
trend in late summer. The models generally capture the

magnitude of this trend and its seasonal variation. There

is a wide variation in the 2-D model results, which is similar to WMO (2003); this is expected, as these models will

not have evolved substantially over this period. The CTMs

capture the seasonal variations but tend to overestimate

the trend in late spring. In contrast, the comparison in the

SH is very poor. The observations show a trend of −4/%

decade with only a small seasonal variation. This lack of

seasonal variation is produced by some 2-D models but

there is a wide range in magnitude of trend, reflecting the

extent of the models’ SH polar loss. The 3-D CTMs, with

their expected more realistic treatment of the SH polar loss

and vortex dynamics, produce a trend which is larger than

observed and shows a clear seasonal cycle. In these runs

the formation of the Antarctic ozone hole in AugustSeptember seems to impact midlatitudes and increase the

trend there. Given the better physical basis for the 3-D

models, the apparently better agreement for the seasonality of the trend from some 2-D models should not be

taken to mean that this modeling approach is better.

The trends in the vertical profile of O3 from the

models and observations are shown in Figure 3-24. The

observations show a characteristic double peak with loss

around 40 km due to the ClO + O cycle and loss in the

lower stratosphere due to halogen/aerosol chemistry and

the impact of loss at higher latitudes. In the upper stratosphere the 3-D CTMs and model 2-D models produce

trends similar to the observations, with a peak around 40-

45 km. The Leeds-Bremen, NOCAR, and GSFC-INT

2-D models produce a trend that is about a factor of two

less than observed. These three models, in contrast to the

other 2-D models and CTMs, calculate interactive temperatures in the upper stratosphere. On this basis they

might be expected to be more realistic than models that

use fixed temperatures, although yearly varying analyses

should also be realistic. However, this depends on the

accuracy of the modeled temperature trends and any trends

that causes in circulation. The interactive 2-D models produce a peak upper stratospheric cooling over the same time

period of around 1-1.4 K/decade (not shown), which is in

reasonable agreement with observations (Chapter 5) and

clearly better than models that assume no trend. These

interactive 2-D models also produce the largest CH4 trend

in the upper stratosphere, indicating a feedback in these

models that increases the meridional circulation (not

shown). Both of these factors act to reduce the magnitude

of the ozone trend. Overall these comparisons shown that

while we have a reasonable understanding of upper stratospheric ozone depletion, we still do not have full quantitative agreement between models and observations. Around

20 km in the NH lower stratosphere, the models again

overall produce a similar trend to that observed. There are

no trend estimates from observations below 20 km in the

SH but, consistent with the SH comparisons shown above,

there is a large variation between the models. The CTMs

give similar results while the 2-D models can give very

small trends, or in the case of the AER model, a large trend

between 10-15 km.

GLOBAL OZONE

3.37

Figure 3-23. Seasonal variation of observed (black

line) and modeled column ozone trends at (top) 35°N60°N and (bottom) 35°S-60°S. The model results

are from 2-D models and 3-D CTMs. The trends were

estimated using regression to an EESC curve and

converted to %/decade using the variation of EESC

with time in the 1980s. Observed trends were estimated from the merged satellite dataset, with error

bars denoting 2σ uncertainties.
On the basis of these comparisons, it is evident that

many broad features of the observed ozone trends are consistent with models that are forced with observed variations in halogens and aerosols. The models generally perform better in the Northern Hemisphere than in the south.

Overall, the 3-D CTMs perform as least as well as the 2-D

models, although there are fewer of them to compare. In

particular, the behavior of the models in the SH is more

similar, though they do show a significant discrepancy

with the observed magnitude and seasonality of the trend.

We now use results from the SLIMCAT CTM runs

to diagnose the roles of chemistry and dynamics in past

ozone changes in more detail. As discussed in Section

3.4.2, CTMs forced by meteorological analyses have been

used to estimate the role of dynamics on (NH) midlatitude

ozone. Full chemistry CTMs can also be used to quantify

the effect of different chemical forcings. For example,

Chipperfield (2003) reported a similar SLIMCAT CTM

integration to Hadjinicolaou et al. (2002) but including

detailed chemistry for cases with and without the inclusion of halogen loading and time-dependent aerosol. For

most months and latitudes, the run with halogen loading

and aerosol data yielded a better agreement with observed

ozone time series. Moreover, the magnitude of the modeled halogen effect was similar to the observed long-term

decrease. It was therefore concluded that the modeled

ozone decrease at all latitudes up to the early 1990s,

including northern midlatitudes, was dominantly caused

by halogen-related trends in combination with heterogeneous chemistry on lower stratospheric polar stratospheric

clouds (PSCs)/aerosols. The runs discussed above (Table

3-1) update these studies using ERA-40 winds

(Hadjinicolaou et al., 2005; Feng et al., 2006). Feng et al.

(2006) ran the CTM from 1977 until 2005, using ERA-40

winds until the end of 2001 followed by operational

analyses. In addition to the basic model run with timedependent source gases, they performed a run with fixed

(1980 troposphere) halogen loadings and a run without a

6 pptv contribution of Bry from very short-lived substances

(VSLS). As shown in Figure 3-25, the comparison of the

basic model runs with the TOMS/SBUV anomaly is similar to that shown in Hadjinicolaou et al. (2005) (Figure 3-

21), where the model captures some observed variability

but, through the use of ERA-40 winds, produces some features (variability) which are not observed. As discussed

in Section 3.4.1, the use of meteorological analyses for

GLOBAL OZONE

3.38

Figure 3-24. Vertical profiles of ozone trends over the period 1980-2004 from observations compared with 2-D

models and 3-D CTMs for (left) 60°N-35°N and (right) 35°S-60°S. The trends were estimated using regression

to an EESC curve and converted to %/decade using the variation of EESC with time in the 1980s. The observations are from SAGE I+II in the upper stratosphere and from ozonesondes in the lower stratosphere (NH only).
trend studies needs to be treated with caution. Assuming

that the analyses are able to give a realistic circulation in

the model, then the CTM can be used for chemical experiments, although these are not coupled (i.e., any indirect

circulation effect of a chemical forcing will not be captured). Figure 3-25 shows that when the model is run with

fixed halogen loadings, there is a difference of about 4-

7% in the NH and SH anomalies, or 20-22 DU of the

column, in the model ozone by 2000 relative to the run

with time-varying halogens, confirming the important

overall role of the halogen trends. The 3-D CTM results

in Figure 3-25 also confirm the 2-D results of Salawitch et

al. (2005) (see Section 3.4.3). An additional 6 pptv of Bry

in the lower stratosphere decreases the model O3 column

by about 10 DU (Feng et al., 2006) but, when expressed

as an anomaly relative to 1980 (a low aerosol period), an

additional relative depletion is only noticeable around the

time of the Mt. Pinatubo eruption. Finally, we can note

that by about 2000, and before the model results are

affected by the switch of analyses, the modeled contribution from halogens to midlatitude ozone loss is starting to

decrease slightly, i.e., there is only a small contribution of

halogen changes to recent changes in ozone. Using a photochemical box model, Yang et al. (2006) argued that the

increase in ozone for altitudes below 18 km is most likely

driven by changes in transport (see Section 3.4.2), rather

GLOBAL OZONE

3.39

Figure 3-25. Time series of annual

mean anomalies (with respect to

1980) for three simulations of the

SLIMCAT 3-D CTM compared with

observations (black line). The

SLIMCAT simulations considered

full time-dependent halogen chemistry (run P5, green dotted line), a

simulation without 6 pptv Bry from

short-lived species (run P1, red

dashed line), and a simulation with

fixed (1980 tropospheric) halogen

loadings (blue dashed line). The

vertical line at the start of 2002 indicates where the CTM forcing changed from ECMWF ERA-40 analyses

to ECMWF operational.
GLOBAL OZONE

3.40
GLOBAL OZONE

3.41

than by declining chlorine and bromine, but the cessation

of ozone depletion between 18-25 km altitude is consistent with a leveling off of stratospheric abundances of chlorine and bromine.

As discussed in Chapter 6, there are now many coupled Chemistry-Climate Models (CCMs) that have been

developed to study past and future changes in the stratosphere. These models have chemistry schemes similar to

those in 2-D models and CTMs, but they calculate their

own interactive circulation, which will be much more variable than interactive 2-D models. The models are also

computationally expensive and often can only run one scenario either as a single simulation or as a few-member

ensemble. Hence, these models have not yet been applied

to diagnosing past changes in midlatitude ozone; rather

the models themselves are still at the validation stage.

Andersen et al. (2006) presented results of three CCMs

compared with the 2-D models used in WMO (2003); the

CCM results varied significantly. More recently, Eyring

et al. (2006) compared results from 13 CCMs that ran with

the same external forcings. These were the REF1 simulations described in Chapter 6 of this Assessment. The

CCMs produce a good simulation of the global temperature fields. Results of the stratospheric age of air and tropical tape recorder signal showed a large spread of results,

although about half of the models were realistic; this shows

an improvement compared with CCMs used in previous

Assessments. Eyring et al. (2006) also discussed the performance of these models in reproducing polar ozone

depletion. Overall the differences in modeled transport,

and likely the treatment of lower stratospheric chemical

processes at all latitudes, leads to a range in modeled global

ozone trends over the past 20 years.

Figure 3-26 shows the results from 13 CCMs for

the seasonal cycle in column ozone and its anomaly for

various latitude bands (see Chapter 6 for model details),

extending on the analysis of Eyring et al. (2006). At NH

and SH midlatitudes the models all capture the expected

seasonal cycle but the modeled column varies by up to 160

DU. The models also all tend to overestimate the tropical

column ozone. In the tropics some CCMs capture the

observed small long-term change but many, as in the case

of 2-D models and 3-D CTMs, produce a large long-term

decrease. Note that only some of the CCM simulations

include a treatment of the solar cycle (see Chapter 6). For

the midlatitude anomalies, the models generally show a

decreasing trend through to the 1990s, in response to the

imposed halogen and aerosol variations, but again there is

a wide range in the model predictions. However, as noted

in Chapter 6, some of the outlying simulations should be

ignored due to unrealistic levels of inorganic chlorine in

the lower stratosphere. Still, given the expected dynamical variability in the models, it is not possible to draw

robust conclusions from this range of model runs.

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Appendix 3A

DESCRIPTION OF THE OZONE DATASETS

3A.1 Column Ozone Data

Ground-Based Data

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, 2003), and

therefore only a brief description of new developments 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. The quality of total ozone data available from 17 Dobson sites prior to 1957 was assessed by

Brönnimann et al. (2003a, b). About half of these pre-1957 datasets are reliable for analysis of total ozone variability;

however, uncertainties in the absolute calibrations make all sets unsuitable for trend analysis. Several old Dobson records

have been recently re-evaluated: from Tromsø, Norway, 1935-1972 (Hansen and Svenøe, 2005); Svalbard, Spitzbergen,

Norway, 1950-1962 (Vogler et al., 2006); and Hradec Kralove, Czech Republic, 1962-2003. Griffin et al. (2006) reanalyzed datasets of 1953-1972 astronomical observations from Haute Provence, France, and Jungfraujoch, Switzerland, to

calculate nightly column ozone. The retrieved values agree well with Arosa Dobson observations, but with a bias of

6-7%. The Brewer spectrophotometer was developed in the early 1980s (Kerr et al., 1981). There are now about 200

instruments installed around the world. They are regularly calibrated against a traveling standard. The traveling standard

itself is calibrated against the set of three Brewer instruments located in Toronto and known as the Brewer Reference

Triad (Fioletov et al., 2005). Small systematic differences between Dobson and Brewer total ozone measurements have

been reported (Staehelin et al., 2003). Filter ozonometers are widely used in the former Soviet Union countries, and longterm reliable records are available from 1972 (Shalomyanskij et al., 2004). This instrument is less accurate than the

Dobson and Brewer instruments, and the calibration is traceable to the Dobson reference. Dobson, Brewer, and filter

ozonometer data are available from the World Ozone and UV Data Centre (WOUDC) at http://www.woudc.org. The

ground-based dataset used in the Assessment is an updated version of the ground-based zonal mean dataset (Fioletov et

al., 2002).

Total Ozone Mapping Spectrometer (TOMS) 

Data from a series of TOMS instruments (on Nimbus 7, Meteor 3, and Earth Probe) have been reprocessed using a

new version 8 algorithm (Bhartia, et al., 2004). The calibration was reviewed and updated during this process to be consistent from instrument to instrument. The TOMS version 8 algorithm uses a more detailed, improved a priori in the

retrieval. The Nimbus 7 TOMS calibration has a slightly different time dependence than in version 7 because of a reassessment of how to handle the uncertainties in noise when TOMS started exhibiting occasional chopper non-synchronization

effects in the mid-1980s. The Earth Probe TOMS data began to have a scan-mirror problem starting in about year 2000

that could not be corrected. The data should probably not be used for trend determination beyond year 2000. By 2002,

there were latitude-dependent errors of 4% to 10%. TOMS data prior to July 1999 were included in the Merged satellite

dataset described below. Data are available from the TOMS web site at http://toms.gsfc.nasa.gov/.

Backscatter Ultraviolet (BUV)

The BUV instrument launched on Nimbus 4 in April of 1970 provided the first global measurements of ozone from

space (Stolarski et al., 1997). Data from this instrument are valuable for extending the satellite data record into the preozone depletion period. Data from 1970 through 1972 provide good global coverage, but thereafter coverage is more

sparse after the partial failure of the solar power array. Data from BUV were recently processed through the latest (version 8) SBUV algorithm (Bhartia et al., 2004). The calibration of BUV has always been questionable because of very

rapid degradation of its solar diffuser plate. For the recent reprocessing, Umkehr data were used to establish the basic

long-term calibration. The soft calibration techniques that were developed to calibrate SBUV instruments were used to

establish the absolute accuracy of BUV. Comparison with Dobson data from that period shows that the resulting data

were reasonably well calibrated.
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Solar Backscatter Ultraviolet (SBUV, SBUV/2)

The SBUV data have also been reprocessed using a new profile retrieval algorithm (Bhartia et al., 2004). This one

is also designated version 8, and while there are some similarities, it is not the same algorithm as used for TOMS. The

SBUV algorithm, in contrast to TOMS, now has a time-independent a priori to remove the possibility that a trend in the a

priori could contribute to the deduced trend in ozone. The calibrations for the SBUV instrument on Nimbus 7 and for the

SBUV/2 instruments on the sequence of NOAA satellites (9, 11, 14, 16, 17, and 18) have now been put on a consistent

scale. The data for upper stratosphere profile can now be used for trend analyses directly without adjustment (though with

some associated uncertainties). Data are available from http://www.cpc.ncep.noaa.gov/products/stratosphere/sbuv2to.

The observational data used for this Assessment are zonally averaged, integrated vertical profile SBUV/2 data (version 8)

from the period from 1979 to 2005 (updated from Miller et al., 2002).

Global Ozone Monitoring Experiment (GOME)

The GOME on board the European Space Agency’s ERS-2 satellite is the first European experiment dedicated to

global ozone measurements (Burrows et al., 1999). ERS-2 was launched in April 1995. GOME has been operational

since June 1995, but spatial coverage has been limited since July 2003 due to problems with tape storage on ERS-2.

GOME measures the backscattered radiances from 240-790 nm in the nadir-viewing geometry. In the relevant region for

total ozone retrieval (320-340 nm), the spectral resolution is about 0.17 nm. The maximum scan width in the nadir is 960

km across track on the ground and global coverage is achieved within three days.

Three algorithms were developed for GOME applications in 2003 and used to reprocess the GOME total ozone data.

These algorithms are all based on the DOAS (Differential Optical Absorption Spectroscopy) technique. They are (1) the

WFDoas algorithm (Coldewey-Egbers et al., 2005; Weber et al., 2005; http://www.iup.uni-bremen.de/gome/wfdoas/); (2)

the TOGOMI/TOSOMI algorithm (Eskes et al., 2005; http://www.temis.nl/protocols/O3total.html); and (3) the GDOAS

algorithm (Van Roozendael et al., 2006; http://wdc.dlr.de/sensors/gome/gdp4.html). For all three algorithms, retrieved

GOME total ozone show excellent agreement with each other and with ground-based Brewer and Dobson measurements at

midlatitudes, so only one GOME dataset is shown in Figures 3-2, 3-3, 3-4, and 3-5 (WFDoas version 1 algorithm).

Merged TOMS+SBUV/2

The merged dataset consists of monthly-mean zonal and gridded average datasets constructed by merging individual TOMS and SBUV/2 satellite datasets (Frith et al., 2004). An external calibration adjustment has been applied to

each satellite dataset in an effort to calibrate all the instruments to a common standard. All data in the present version of

the merged dataset have been derived using the TOMS version 8 and SBUV version 8 algorithms. A recent manuscript

(Stolarski and Frith, 2006) evaluates the instrument record uncertainty for one of these datasets and shows that the variations among them are within the uncertainties. For this Assessment, BUV total ozone data are also included in the merged

dataset without any adjustments. There are inherent uncertainties is the possible drift of each instrument and in the offset

of overlapping instruments. For the merged dataset, Stolarski and Frith (2006) estimated an overall drift uncertainty of a

little more than 1%/decade due to instrument effects. This corresponds to about 8 DU over the 25-year period of the

measurements. The data, and information about how they were constructed, can be found at http://hyperion.gsfc.nasa.gov/-

Data_services/merged/mod_data.public.html.

National Institute of Water and Atmospheric Research (NIWA) Assimilated Dataset

The assimilated database combines satellite-based ozone measurements from 4 Total Ozone Mapping Spectrometer

(TOMS) instruments, 3 different retrievals from the Global Ozone Monitoring Experiment (GOME), and data from 4 Solar

Backscatter Ultraviolet (SBUV) instruments (Bodeker et al., 2005). The dataset used in this analysis is an update and extension of the homogenized total column ozone dataset developed by Bodeker et al. (2001). Specific changes include: version 8

Nimbus 7 and Earth Probe TOMS data are used rather than version 7 (http://toms.gsfc.nasa.gov/). GOME data from the

European Space Agency (ESA) are updated from version 2.4 (used in Bodeker et al., 2001) to version 3.1. Assimilated total

column ozone fields from the Royal Netherlands Meteorological Institute (KNMI) (http://www.knmi.nl/goa) based on GOME

measurements are included. Total column ozone fields from the KNMI total ozone algorithm from GOME using the Ozone

Monitoring Instrument algorithm (http://www.temis.nl/protocols/O3total.html) are included. Version 8 SBUV (Solar

Backscatter Ultraviolet) data from the NASA Nimbus 7, NOAA 9, NOAA 11, and NOAA 16 satellites (http://-

www.cpc.ncep.noaa.gov/products/stratosphere/sbuv2to/) are included. The analysis period is extended to the end of 2004.
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3A.2 Ozone Profile Measurements

Stratospheric Aerosol and Gas Experiment (SAGE)

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 (Wang et al., 1996). SAGE II ozone profile data commenced in November 1984

and the science mission ended in August 2005. The analysis here used SAGE II version 6.2 (v6.2), which was released in

its entirety in September 2005. In mid-2000, SAGE II experienced a failure of the azimuth gimbal in the pointing system.

After an approximate two-month interruption in science data gathering, normal science operation resumed at a reduced

duty-cycle. The failure precluded the ability to take both sunrise and sunset observations on each orbit. Events of only one

type occur in blocks of approximately 35 days. The event type switched from one type to the other after each spacecraft

yaw-maneuver (performed in order to maintain proper illumination of the spacecraft’s fixed solar panels). The spatial and

temporal sampling in the reduced mode of operation does not introduce any new significant gaps or biases in the latitudinal

and seasonal coverage. The species profiles have a vertical resolution of 1 km or better and are placed on a 0.5 km altitude

grid. The species separation and inversion in v6.2 is similar to that used in v6.1 (the basis for the SAGE II ozone data in

WMO (2003)), but v6.2 incorporates an improved spectral characterization of the 940 nm channel optical properties that

results in significantly improved water vapor retrievals. In general, the ozone results are of good quality down to the

tropopause. As in previous Assessments, the data analyzed here exclude time periods of enhanced aerosol extinction following the eruption of Mt. Pinatubo in June 1991. A small percentage of SAGE II data is affected by errors caused by

interference from aerosol (including stratospheric aerosol from major volcanic eruptions), clouds, and other, not always

known, factors and should be removed (Rind et al., 2005). The data are available from http://eosweb.larc.nasa.gov/.

Solar Backscatter Ultraviolet (SBUV, SBUV/2)

The SBUV/2 instrument is a scanning double monochromator measuring backscattered solar radiation in 12 discrete

wavelength bands ranging from 252.0 to 339.8 nm. In previous SBUV algorithms, total column ozone was retrieved using

the longest 4 wavelengths, and then a profile was retrieved using the shortest 8 wavelengths. In the version 8 algorithm

released in 2004, an ozone profile is retrieved using all 12 wavelengths, and total column ozone is the integral of the profile

(Bhartia et al., 2004). The version 8 algorithm is optimized to provide a self-consistent long-term ozone record. The

SBUV(/2) satellite data used here are reprocessed with the version 8 algorithm and are available at http://daac.gsfc.nasa.gov/-

data/dataset/TOMS/DVD-ROMs and from http://www.cpc.ncep.noaa.gov/products/stratosphere/sbuv2to. The data are

available as column ozone in Dobson units (DU) for 13 layers, each 3.2 km thick except for Layer 1, which is about 18 km

in depth, and as ozone mixing ratios at 15 levels. To avoid errors related to volcanic aerosols, it is recommended to exclude

from the analysis data between 40°S and 40°N for the 1-year periods following the major volcanic eruptions in 1982

and 1991 (SPARC, 1998). Results of SBUV(/2) ozone profile comparisons with other data sources are discussed by

Petropavlovskikh, et al. (2005a), Nazaryan and McCormick (2005), Fioletov et al. (2006), and Terao and Logan (2006).

Combined datasets are now becoming available for profile data. Stolarski and Frith have put together the SBUV series

from the version 8 results (http://hyperion.gsfc.nasa.gov/Data_services/merged/mod_data.public.html). Brunner et al.

(2006) reconstructed the vertical ozone distribution in equivalent latitude coordinates.

Ozonesonde Data 

A network of stations provides ozone profile information from balloonborne sondes, which measure ozone from

the ground to about 33 km, with a vertical resolution of ~150 m. Details of ozonesonde measurements and data processing are discussed in SPARC (1998). Data are available from the WMO World Ozone and UV Data Centre (WOUDC)

(http://www.woudc.org). Data from many sonde stations are normalized to the overhead ozone column measured on the

same day by a Dobson or Brewer instrument. The normalizing factor (called the correction factor) is commonly used to

screen the data, as described in Logan et al. (1999). Data are available from about 1970 for stations in North America,

Europe, and Japan. The European stations make measurements 2-3 times a week, while other stations make weekly
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measurements. The long-term records from Payerne and Uccle have been reprocessed since the SPARC (1998) assessment, and the Japanese sonde data since 1994 are now available with much higher vertical resolution than previously.

Umkehr Data

The Umkehr technique is used to derive the vertical distribution of ozone from ground-based measurements of

zenith sky radiation. The method is based on principles of differential absorption of solar UV light by atmospheric ozone

and molecular scattering. The ratio of zenith blue sky radiation at two UV wavelengths is measured over a range of solar

zenith angles between 60° and 90°. A representative total ozone measurement is required as part of the data input. The

resulting profile is reported as mean partial pressure values for ten pressure layers, where the pressure at the top of the layer

is half of the pressure at the bottom of the layer, and layer 10 contains ozone integrated above the ~1 hPa pressure level.

The analysis is based on a climatological first guess and an iterative solution is reached. The dependence of Umkehr

retrievals on a priori information becomes a serious issue for trend analysis. (Dütsch and Staehelin, 1992; Mateer et al.,

1996). In the presently used operational UMK92 algorithm, the a priori profiles are constructed using total ozone column

measured by the same instrument. This makes it difficult to ascertain whether the retrieved long-term changes are forced

by a priori or whether they reflect information contained in the measurements. The new UMK04 algorithm includes revised

a priori profiles based on the new ozone climatology (McPeters et al., 2003) that vary with season and latitude, but have no

day-to-day or long-term variability. It also uses new look-up tables and an improved forward model. In addition, the

impacts of both algorithm smoothing and measurement errors on the retrieved ozone profiles have been further optimized

(Petropavlovskikh et al., 2005a, b). The overall accuracy of the Umkehr method (1σ) is estimated to be ±25% for the troposphere (layers 0 and 1), ±15% for low stratosphere (layers 3 and 2), less than ±10% for the middle stratosphere (layers 4 to

6), less than ±10% for the upper stratosphere (layers 7 and 8), and errors slightly increase in ozone integrated above layer

7. The algorithm description and the data are available at http://www.srrb.noaa.gov/research/umkehr/. Zanis et al. (2006)

described re-evaluation of the 1956-2003 Umkehr series from Arosa, Switzerland. The vertical profiles of ozone trends

from Umkehr data shown in Figure 3-8 are the average of trends derived from measurements at Arosa (1979-2004), HauteProvence (1984-2004), Boulder (1979-2004), and Belsk (1979-2004), processed with the new UMK04 algorithm.

Lidar Data

The Differential Absorption Lidar (DIAL) ozone measurement technique retrieves the vertical profile of ozone

absorption by comparing atmospheric return at two UV wavelengths, one absorbed by ozone (usually 308 nm) and one

much less absorbed (353 or 355 nm) (Donovan et al., 1997; Guirlet et al., 2000; Leblanc and McDermid, 2000, 2001;

Godin et al., 2001, 2002). The DIAL method is differential in wavelength and altitude, which makes it self calibrating and

ideally suitable for long-term routine measurements. Altitude resolution of the measurements ranges from 1 km (below

30 km) to 5 km (at 45-50 km), and the altitude range is typically from 15 to 45-50 km, with accuracies of approximately

3% below 35 km and 10% above 40 km. Observations are made during clear nights. High quality, continuous time series

are available since the late 1980s (Keckhut et al., 2005b). The data are available from the Network for the Detection of

Atmospheric Composition Change (NDACC) site at http://www.ndacc.org/.

Microwave Ozone Measurements

The method is based on measurements of microwave thermal radiation in the ~100-200 GHz range, emitted by

ozone and measured at high spectral resolution. As the observed emission line is broadened by pressure, information

about the vertical distribution of ozone in the atmosphere can be retrieved from the measured spectra by means of an

inversion algorithm. Individual rotational spectral lines with frequencies most often either 100 or 142 GHz, are used.

A review of the methodology and instrumentation used for ozone observations in the microwave frequency range is

given, for example, in Parrish (1994) and Clancy and Muhleman (1993). Ground-based microwave radiometers allow

unattended, continuous observations of the atmosphere that are nearly independent of weather conditions. A vertical

resolution is typically ~10 km and the temporal resolution is 1-2 hours. The altitude range is typically 20-70 km, with

accuracies of ~5%. High quality, continuous time series are available since 1992 (Tsou et al., 2000; Klein et al., 2002;

Calisesi et al., 2003; Nagahama et al., 2003; Ricaud et al., 2004). The data are available from the NDACC site at

http://www.ndacc.org/. 
CHAPTER 4

Polar Ozone: Past and Present

Contributors:

G.E. Bodeker

K. Frieler

F. Goutail

M. López-Puertas

G.L. Manney

E.R. Nash

C.E. Randall

H.J. Singer

C.S. Singleton

R.M. Stimpfle

S. Tilmes

M. Weber

Lead Authors:

P.A. Newman

M. Rex

Coauthors:

P.O. Canziani

K.S. Carslaw

K. Drdla

S. Godin-Beekmann

D.M. Golden

C.H. Jackman

K. Kreher

U. Langematz

R. Müller

H. Nakane

Y.J. Orsolini

R.J. Salawitch

M.L. Santee

M. von Hobe

S. Yoden

Final Release: February 2007

From Scientific Assessment of Ozone Depletion: 2006

CHAPTER 4

POLAR OZONE: PAST AND PRESENT

Contents

SCIENTIFIC SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1

4.0 INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.3

4.1 POLAR STRATOSPHERIC OBSERVATIONS UPDATE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.4

4.1.1 Polar Stratospheric Dynamics and Transport . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.4

4.1.1.1 Temperatures and PSC Formation Potential . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.4

4.1.1.2 Vortex Parameters . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.7

4.1.1.3 Transport . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.9

4.1.2 Polar Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.11

4.1.2.1 Ozone and Other Constituents . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.11

4.1.2.2 Polar Ozone Loss Studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.14

4.2 PROGRESS IN OUR UNDERSTANDING OF THE PHYSICAL AND CHEMICAL PROCESSES . . . . . . 4.20

4.2.1 Polar Ozone Chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.20

4.2.1.1 Chlorine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.21

4.2.1.2 Bromine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.24

4.2.1.3 Ozone Loss Rates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.25

4.2.2 PSC Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.28

4.3 RECENT POLAR OZONE CHANGES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.30

4.3.1 2002 Antarctic Ozone Hole . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.30

4.3.1.1 Observations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.30

4.3.1.2 Modeling of the Warming . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.31

4.3.1.3 Theoretical Understanding . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.33

4.3.2 Is Polar Ozone Getting Worse? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.34

4.3.2.1 Arctic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.34

4.3.2.2 Antarctic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.34

4.4 THE INFLUENCE OF PRECIPITATING CHARGED PARTICLES ON POLAR OZONE . . . . . . . . . . . . . 4.35

4.4.1 Odd Nitrogen (NOy) Enhancements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.36

4.4.2 Ozone Decreases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.37

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.37

POLAR OZONE

4.1

SCIENTIFIC SUMMARY

Arctic

• Arctic spring total ozone values over the last decade were lower than values observed in the 1980s. In addition, spring Arctic ozone is highly variable depending on dynamical conditions. For current halogen levels,

anthropogenic chemical loss and variability in ozone transport are about equally important for year-to-year Arctic

ozone variability. Colder-than-average vortex conditions result in larger halogen-driven chemical ozone losses.

Variability of temperatures and ozone transport are correlated because they are both driven by dynamic variability.

• For the coldest Arctic winters, the volume of air with temperatures low enough to support polar stratospheric

clouds (VPSC) increased significantly since the late 1960s. This change of climate conditions is much larger than

expected from the direct radiative effect of increasing greenhouse gas concentrations. The reason for the change is

not clear and it could be due to long-term natural variability or an unknown dynamical mechanism.

• Column ozone loss in the 2004/2005 Arctic winter was among the largest ever observed. The 2004/2005 Arctic

stratosphere was exceptionally cold, particularly below 18 kilometers (km), leading to a value of VPSC 25% larger

than the previous record value. Various independent studies and methods suggest that the chemical column ozone

loss in 2004/2005 was among the largest ever observed. However, dynamical processes resulted in March-average

Arctic total ozone amounts being comparable to those in other recent winters, while the large losses contributed to

low total column ozone over parts of Europe during March 2005.

Antarctic

• Over the last decade (1995-2005), Antarctic ozone depletion has stabilized. Most ozone hole diagnostics show

a leveling off after the mid-1990s. Saturation of ozone loss inside the ozone hole due to complete ozone destruction

over a broad vertical layer plays a role in this leveling off. Ozone hole area, ozone mass deficit, and higher minimum column amounts were observed to be below average in some recent winter years. These improvements in the

ozone hole resulted from higher levels of dynamical forcing, and not decreases in equivalent effective stratospheric

chlorine levels.

• In September 2002, the first ever observed Antarctic major stratospheric warming occurred. This earlyspring warming caused a drastic reduction of the ozone hole area and resulted in a less severe ozone hole.

This warming resulted from anomalously strong dynamical wave activity in the Southern Hemisphere. The triggers

of these very unusual waves are unknown, and it is not clear whether the 2002 warming is a random event due to

internal atmospheric variability or whether it can be related to long-term changes in climate. The Antarctic winter

2004 was also dynamically very active and had less ozone mass deficit than previous years. The higher levels of

ozone in these two years were dynamically driven and not related to halogen chemical reductions.

General

• Large interannual variability in polar stratospheric temperatures complicates the interpretation of trends.

Previously reported estimates of temperature trends in polar regions have differed from assessment to assessment.

It now is evident that trends determined over a time scale of one to two decades, though they may appear statistically significant, are not robust because of large interannual and decadal variability in observed temperatures.

Therefore, changes in reported trends do not necessarily indicate systematic changes in physical or chemical

processes.

• There are indications that the chlorine monoxide (ClO) dimer cycle may be a more efficient process for polar

ozone loss than previously thought. Uncertainties in the laboratory absorption cross section of the chlorine
POLAR OZONE

4.2

monoxide dimer (ClOOCl) are large. Good overall consistency between in situ observations of ClO and the ClOOCl

and model calculations can be achieved if it is assumed that ClOOCl photolyzes faster than assumed in the previous

Assessment. Faster photolysis of ClOOCl, combined with recent laboratory studies indicating more rapid formation of ClOOCl by the ClO self-reaction, leads to more efficient ozone loss by this catalytic cycle.

• Recent measurements suggest that bromine may play a more important role in polar ozone depletion than

previously thought. Profiles of bromine monoxide (BrO) measured in the Arctic vortex suggest that inorganic

bromine levels may be 3 to 8 parts per trillion by volume larger than the amount of bromine carried to the stratosphere by methyl bromide (CH3Br) and halons. This observation indicates the BrO + ClO cycle is likely to be a

more efficient ozone loss process than considered in the previous Assessment. The BrO + ClO cycle is now estimated to contribute up to half of total chemical loss of polar ozone, even considering the more efficient ozone loss

by the ClO dimer cycle. However, the polar stratospheric bromine budget is a significant source of uncertainty.

• Calculated chemical loss rates of polar ozone substantially increased in models that assume: (1) more efficient ozone destruction by the ClO dimer cycle and (2) higher levels of bromine. These changes improve the

comparison between measured and modeled Arctic ozone loss rates for cold Januaries that was noted in the past

Assessment. These two changes also improve the comparison between theory and observation of Antarctic ozone

loss rates.

• The chemical loss of column ozone for Arctic winters exhibits a near linear relation with VPSC during each

winter. A similar relation between ozone loss and VPSC is now seen for two independent analyses of chemical

ozone loss, increasing our confidence in the robustness of this relation. Observations of the large Arctic ozone loss

in 2004/2005 are in line with the relationship established for prior winters. The slope of this relation provides an

empirical measure of the sensitivity of chemical loss of Arctic ozone to changes in stratospheric temperature (for

contemporary levels of chlorine and bromine) and provides an important metric for chemical transport models

(CTMs) and coupled Chemistry-Climate Models (CCMs) that are used in both diagnostic and prognostic studies of

polar ozone loss.

• For the first time, measurements show unambiguously that nitric acid trihydrate (NAT) polar stratospheric

cloud particles can nucleate above the ice frost point, and there is additional evidence of their widespread

occurrence. Widespread low number density NAT clouds can lead to denitrification and enhanced ozone loss.

Incorporating NAT nucleation above the ice frost point into chemical transport models has improved denitrification

simulations, but discrepancies in interannual variability remain, probably because the NAT nucleation mechanisms

are not fully understood.

• The most recent solar cycle was associated with frequent and exceptionally strong episodes of charged particle precipitation. This caused only small decreases in total column ozone in the polar region (<3%, <10

Dobson units). However, the measured ozone reduction in the mid-to-upper stratosphere (30-50 km) exceeded

30% for weeks following certain episodes of intense particle precipitation.
4.0 INTRODUCTION

Since the discovery of the Antarctic ozone hole by

Farman et al. (1985), considerable effort has been focused

on observing these ozone losses, understanding the chemical, dynamical, and radiative processes, and predicting

the future of polar ozone. This chapter builds upon a

sequence of polar ozone chapters in the World Meteorological Organization (WMO) Assessment series.

The first publication of large Antarctic ozone losses

by Farman et al. (1985) sparked scattered discussion of

polar ozone in the 1985 Ozone Assessment (WMO, 1986).

Chapter 14 of that Assessment (Ozone and Temperature

Trends) showed the Farman et al. (1985) results and included

Total Ozone Mapping Spectrometer (TOMS) total ozone

images (subsequently published in Stolarski et al., 1986).

These early satellite data confirmed the findings of Farman

et al. and moreover displayed the horizontal dimensions of

the ozone hole. The Ozone Trends Panel’s 1988 report

(WMO, 1990a) followed with an entire chapter on Antarctic

ozone (Chapter 11) including an extensive discussion of both

Antarctic observations and of the three competing theories

of the ozone hole that were proposed originally in the mid1980s. In particular, this WMO (1990a) chapter focused on

(1) the drastic reduction in nitrogen dioxide (NO2) abundance resulting from the formation of polar stratospheric

clouds (PSC) consisting of nitric acid trihydrate (NAT)

(Crutzen and Arnold, 1986; Toon et al., 1986), (2) the conversion of hydrogen chloride (HCl) and chlorine nitrate

(ClONO2) to active chlorine on PSC surfaces as a mechanism for freeing chlorine from reservoir species into radical

species (McElroy et al., 1986; Solomon et al., 1986), (3) the

two catalytic cycles involving chlorine monoxide (ClO) and

bromine monoxide (BrO) that destroy Antarctic ozone

(McElroy et al., 1986; Molina and Molina, 1987), and (4)

observations from field campaigns that supported the theory

that the ozone hole is caused by elevated ClO and BrO levels

in the lower stratosphere from human-produced chlorofluorocarbons and halons (de Zafra et al., 1987; P. Solomon et

al., 1987; S. Solomon et al., 1987; Farmer et al., 1987). This

WMO (1990a) polar ozone chapter also discussed in situ

observations of ClO and BrO (subsequently published in

Brune et al., 1989a; Brune et al., 1989b; Anderson et al.,

1989). The 1989 Ozone Assessment (WMO, 1990b) included a chapter on polar ozone (Chapter 1) that presented extensive observations from field campaigns that overwhelmingly

established human-produced chlorine and bromine chemicals as responsible for the ozone hole. The 1991 Ozone

Assessment (WMO, 1992) did not include a specific chapter

on polar ozone, but it updated polar ozone science in various

sections of the report. WMO (1995) included a Polar

Processes chapter (Chapter 3) that confirmed chlorine- and

bromine-catalyzed ozone loss in the Antarctic and the Arctic

from both in situ (Anderson et al., 1991; Toohey et al., 1993)

and satellite observations (Waters et al., 1993). WMO (1999)

included a chapter on lower stratospheric processes (Chapter

7), noting a number of very low Arctic ozone years and

stating that the Antarctic ozone losses were continuing.

This chapter provides an update to Chapter 3 of the

WMO (2003) Assessment on Polar Ozone: Past, Present,

and Future. This chapter is derived from research over the

last 4 years, other reports, and recent and updated observations from various field campaigns and new satellite

instrumentation. Discussion of the recovery of polar ozone

can be found in Chapter 6 (The Ozone Layer in the 21st

Century), and polar chemistry-climate interactions can be

found in Chapter 5 (Climate-Ozone Connections).

The WMO (2003) polar ozone chapter reported that

in the Antarctic, springtime ozone depletion has been large

throughout the 1990s, and that the area of the ozone hole

varies from one year to another, but that it was not yet possible to say whether the area of the ozone hole had maximized. It was also reported that the polar vortex and ozone

hole were persisting into early summer, increasing the

impact on ultraviolet radiation.

In the Arctic, the WMO (2003) polar ozone chapter

reported that there were a number of cold Arctic winters in

the 1990s in which maximum total column ozone losses

due to halogens had been quite large. In particular, Arctic

chemical ozone loss for the 1999/2000 winter/spring season

was extensively discussed, and an ozone loss was observed

that reached 70% near 20 kilometers (km) by early spring,

and total column ozone losses greater than 80 Dobson units

(DU) (~20 to 25%) by early spring. The large 1999/2000

ozone losses were a result of the colder than average stratospheric vortex, consistent with our understanding. It was

also reported that the magnitude of chemical loss of ozone

for all Arctic winters during the 1990s showed 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.

This chapter is divided into 4 basic sections.

Section 1 provides an update of observations of polar stratospheric conditions. Section 2 provides an update of theoretical progress on understanding polar ozone processes.

Section 2 is divided into subsections on chemical processes

and polar stratospheric clouds and aerosols. Section 3 provides a discussion of our current observations and also a

description of an unprecedented splitting of the ozone hole

that was observed in September 2002 just prior to the publishing of WMO (2003). A number of large solar particle

events have occurred in the last few years and they are

described, together with their impacts, in Section 4.

POLAR OZONE

4.3
4.1 POLAR STRATOSPHERIC OBSERVATIONS

UPDATE

4.1.1 Polar Stratospheric Dynamics and

Transport

Variability in the dynamical conditions in the

troposphere/stratosphere system results in variability of

ozone transport and temperatures in the stratosphere.

Previous Assessments and Section 4.2.1 of this document

show that on short time scales, interannual variability in

polar ozone chemistry is mainly driven by temperature

variability, which in turn is the result of variable dynamical conditions. The combined effect of dynamically

induced variability in both chemistry and transport is the

main driver of interannual variability of the abundance

of ozone in the polar stratosphere. This section gives an

update of the dynamical conditions, temperatures, and

transport of the polar winter stratosphere since the previous Assessment.

4.1.1.1 TEMPERATURES AND PSC FORMATION

POTENTIAL

WMO (1999) and WMO (2003) have dealt at length

with the observed temperature changes detected over the

polar regions using observations (radiosondes and satellite sensors) as well as objective analysis products (e.g.,

Ramaswamy et al., 2001). During the period 1979-1998,

satellite measurements showed a cooling of 3 K/decade in

the Arctic and Antarctic middle stratosphere and significant coolings of 1.8 K/decade at 70°N and 1.1 K/decade

at 70°S in the lower stratosphere during spring.

Following the mostly cold stratospheric winter and

spring seasons of the early to mid-1990s, polar regions in

both hemispheres were characterized by an alternation of

cold and warm winter seasons since the late 1990s. As

seen in Figure 4-1, Arctic lower stratospheric temperatures show considerable interannual variability. The variability is strongest during winter and spring due to the

occurrence of midwinter and final sudden stratospheric

warmings. Although several recent northern winters were

cold (e.g., 1999/2000, 2002/2003, and in particular

2004/2005; see Figure 4-1), the northern winters

1998/1999, 2000/2001, 2001/2002, 2003/2004, and

2005/2006 were disturbed by stratospheric warmings and

thus exhibited periods with high temperatures. In the

Antarctic, interannual variability is much weaker than in

the Arctic and is strongest during spring due to the timing

of the Antarctic vortex breakup. In the Antarctic, an anomalously warm spring season was observed in 2002 (see

Section 4.3.1), and relatively high temperatures also

occurred in the years 2000 and 2004 in the lower stratosphere (100 hPa) (see Figure 4-2).

Given the interannual variability of stratospheric

temperatures in winter and spring, determining statistically

significant trends requires both sufficient temporal and

spatial coverage and a careful choice of the applied trend

model. As discussed in Section 5.2.6 of Chapter 5, different temperature datasets are available to derive trends.

All of them have their own advantages and limitations.

Long-term stratospheric temperature trends (since the late

1950s) are usually derived from radiosonde measurements

or radiosonde-based gridded datasets such as the United

Kingdom Meteorological Office (UKMO) adjusted dataset

(RAOB) or the Freie Universität Berlin (FUB) stratospheric

analyses. Radiosonde limitations due to instrumental and

other operational discontinuities could be eliminated by

applying new adjustment methods, e.g., as in the radiosonde atmospheric temperature products for assessing climate (RATPAC) (Lanzante et al., 2003; Free et al., 2005)

and Hadley Centre Radiosonde Temperature 2 (HadAT2)

(Thorne et al., 2005) datasets. However, radiosonde data

are biased toward Northern Hemisphere (NH) continental

locations due to data scarcity in the Southern Hemisphere

(SH) and over extended oceanic regions. Satellite data provide better spatial and altitude coverage than radiosondes,

but are only available since 1979. Their vertical resolution

is limited, and they have to be composed from different

instruments to build continuous time series (see Section

5.2.6 of Chapter 5). Recent assessments of the uncertainties of radiosonde- and satellite-derived temperature trends

revealed disagreement in the lower stratosphere, with an

overestimation of the stratospheric cooling in the

radiosonde data (Seidel et al., 2004). The bias was confined to the tropics, and thus polar temperature trends

should not be affected (Randel and Wu, 2006).

Reanalysis products that result from the assimilation of radiosonde and satellite data provide an alternative source for long-term trend estimates. They are,

however, affected by changes in the input data and the

assimilation systems and therefore are not generally

suited for trend estimates (Bengtsson et al., 2004; Santer

et al., 2004). Labitzke and Kunze (2005) compared

North Pole temperature trends for the lower and middle

stratosphere derived from the National Centers for

Environmental Protection/National Center for Atmospheric Research (NCEP/NCAR) and European Centre

for Medium-Range Weather Forecasts (ECMWF) ReAnalysis 40 (ERA40) reanalyses with those derived

from subjective FUB stratospheric analyses that are

closely matched to direct radiosonde measurements.

They confirmed that largest discrepancies between

ERA40 and FUB data exist during the winter months in

POLAR OZONE

4.4
POLAR OZONE

4.5

170

180

190

200

210

220

230

Temperature (K)

Jul Aug Sep Oct Nov Dec Jan Feb Mar Apr May Jun

10%-90%

30%-70%

1978-2004

2004-2005

Type I PSC

HNO = 10 ppb, H O = 5 ppm

Type II PSC

170

180

190

200

210

220

230

Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec

10%-90%

30%-70%

1979-2005

2002

2005

Type I PSC

Type II PSC

SH

NH

Temperature (K)

Figure 4-1. Annual temperature cycle and variability at 50 hPa for minimum temperature between 50°N and

90°N (top), and between 50°S and 90°S (bottom) from the National Centers for Environmental Prediction/National

Center for Atmospheric Research (NCEP/NCAR) reanalysis data. The black line shows the mean annual cycle

for the periods 1978-2004 (Northern Hemisphere, NH) and 1979-2005 (Southern Hemisphere, SH). The red line

shows the evolution during 2004/2005 (NH) and 2005 (SH). In the case of the SH, the blue line shows the evolution during 2002, i.e., the year with the anomalous Antarctic vortex behavior. Dark gray shading shows the 30 to

70% probability of observations, light gray shading shows the 10 to 90% probabilities, and the lower and upper

thin black lines show the record values for the 27 years of observations. Green lines indicate the thresholds for

PSC formation (nitric acid (HNO3) = 10 ppb and H2O = 5 ppm). Updated from Figure 3-13 in WMO (2003).
the middle stratosphere, when the enhanced cold polar

bias of the pre-satellite era in ERA40 affects the longterm trends (1958-2001) (Bengtsson et al., 2004). The

use of reanalyses for polar spring temperature trends is

justified because the discrepancies between the ERA40,

NCEP/NCAR, and FUB datasets are negligible.

Figure 4-2 shows spring stratospheric temperature

trends derived from NCEP/NCAR reanalyses for the

period 1958-2005 at the altitudes of maximum cooling for

the Arctic polar cap (60°N-90°N) at 30 hPa in March

(Figure 4-2a) and for the Antarctic polar cap (60°S-90°S)

at 100 hPa in October (Figure 4-2b) (updated from

Labitzke and Kunze, 2005). Post-volcanic periods (24

months) following the three eruptions of Mt. Agung in

1963, El Chichón in 1982, and Mt. Pinatubo in 1991 were

excluded to avoid a contamination of the trend.

During the past 25 years, Arctic temperatures showed

a clear downward trend of −1.70 ± 0.20 K/decade in March.

Compared with the trend estimates for the period 1979-1998

(WMO, 2003) however, the magnitude of the stratospheric

temperature trend in Arctic spring has decreased by about

50%. For the full period 1958-2005, only a very small, nonsignificant trend of −0.25 ± 0.20 K/decade can be detected

due to a warming of the Arctic stratosphere during the 1960s

and early 1970s (Figure 4-2a), which nearly compensates

the cooling of the later decades. The dependence of the

trends on the underlying time period demonstrates the strong

impact of low-frequency dynamical variability on polar

temperatures. Multidecadal time series are thus required to

minimize the effect of low-frequency dynamical variability

in trend calculations, in particular in the Arctic. This observational result is further supported by idealized model

experiments (Nishizawa and Yoden, 2005).

In the Antarctic lower stratosphere (100 hPa), the

cooling trend since the late 1970s is substantially smaller

than that reported in WMO (2003) due to increased dynamical variability of the Antarctic polar vortex that led to

unusually high spring temperatures of the polar cap in the

years 2000, 2002, and 2004. Consistently, the temperature change in October between 1979 and 2005 is very

small at 100 hPa (−0.23 ± 0.20 K/decade) (Figure 4-2b)

and becomes even positive at higher altitudes (+1.01

± 0.25 K/decade at 50 hPa and +2.02 ± 0.24 K/decade at

30 hPa). On the long-term scale, a small warming of 0.36

± 0.18 K/decade occurred at 100 hPa.

As in Figure 4-2, long-term atmospheric changes

are usually described in terms of linear trends per decade

using multiyear time series, or the derivation of differences between averaged sub-periods, e.g., the temperature difference between 1990s and the 1970s. The

magnitude and the statistical significance of the calculated changes in both methods strongly depend on the

data period, as seen in Figure 4-2. For the purpose of

understanding the nature of the observed changes, new

statistical approaches were suggested in which trends

may occur in flat or sloped steps including breakpoints

with abrupt changes (Seidel and Lanzante, 2004) or the

use of polynomials (Malanca et al., 2005). A comparison of the methods showed that a description of stratospheric cooling by a gradual linear change since the late

1950s (without considering the singular volcanic eruptions) is as equally valid as a description by a more steplike downward trend concentrated in the two-year periods

after the three major volcanic eruptions (Seidel and

Lanzante, 2004), which was also suggested by Pawson

POLAR OZONE

4.6

1960 1970 1980 1990 2000

1960 1970 1980 1990 2000

-56

-60

-64

-68

-72

56

60

-64

-48

-52

Temperature ( C) Temperature ( C)

March, 30 hPa, 60 ̊-90 ̊N

October, 100 hPa, 60 ̊-90 ̊S

a

b

Year

Figure 4-2. Arctic 30-hPa March (a) and Antarctic

100-hPa October (b) temperatures (°C) derived from

the NCEP/NCAR reanalysis data. Trend lines were

added for the periods 1958-1979 (red), 1979-2005

(black), and 1958-2005 (turquoise). Updated from

Labitzke and Kunze (2005).
et al. (1998). Climate model simulations suggest that the

stratospheric cooling since 1979 was driven by anthropogenic factors (i.e., ozone depletion and increases in

well-mixed greenhouse gases) but modulated by natural

factors (i.e., solar variability and volcanic aerosols)

(Ramaswamy et al., 2006).

In summary, the Arctic stratosphere has cooled over

the past 27 years, however due to stronger dynamical variability in a number of recent winters and springs, the negative temperature trend is only about half the magnitude

of the estimate in WMO (2003). In the Antarctic lower

stratosphere, the cooling trend since the late 1970s has

become less negative than that reported in WMO (2003)

due to increased dynamical variability.

Estimating the overall trend of the polar stratospheric temperatures is not sufficient to assess the impact

of long-term temperature changes on ozone chemistry.

The large Arctic ozone losses are driven by the evolution

of conditions during cold winters. Warm winters are much

less relevant for ozone chemistry, because chemical loss

of ozone is very limited and the chemistry is not very sensitive to small changes in temperature. Section 4.2.1

shows that the interannual variability in Arctic chemical

ozone loss mainly follows the volume of temperatures

cold enough to activate chlorine on particle surfaces (i.e.,

VPSC). Figure 4-3 shows the long-term evolution of VPSC

derived from stratospheric temperature analyses for the

Arctic (updated from Rex et al., 2004). Over the past 40

years, the cold Arctic winters became colder, resulting in

larger values for VPSC and chemical ozone loss. The longterm data shown in Figure 4-3 relies on the FU-Berlin

subjective radiosonde analysis. Hence, the dataset is relatively independent of the introduction of satellite observations or assimilation system changes (e.g., Manney et

al., 2003b). A linear fit through the solid points in Figure

4-3, which represents maximum values of VPSC for 5-year

intervals, has a slope of 9.9 ± 1.1 × 106 km3 per decade,

with very similar results if 4- to 10-year intervals are

chosen to select the maximum values (update of Rex et

al., 2004). Rex et al. (2006) also have shown that it is

unlikely (well below 1% probability) that this estimated

tendency toward colder Arctic winters is a purely random

event or is caused by inconsistencies in the meteorological datasets.

The tendency of decreasing temperatures during

cold winters is in contrast to the clustering of a series of

warmer than average winters during recent years. Hence,

the temperature trend during the cold winters is different

from the overall temperature trend. The change in temperature during the cold winters is much larger than expected

from the direct radiative effect of increasing greenhouse

gas concentrations. The reason for the change is not clear

and it could be due to long-term natural climate variability

or an unknown dynamical feedback mechanism.

4.1.1.2 VORTEX PARAMETERS

The mean structure and seasonal evolution of the

polar vortices and their hemispheric differences were discussed in detail in WMO (2003). The Arctic polar vortex

is much more perturbed by planetary wave activity propagating upward from the troposphere than the Antarctic

vortex. It is important to assess such dynamical variations and their changes as they determine, together with

POLAR OZONE

4.7

1970 1980 1990 2000

0

20

40

60

VPSC [ 106

km3

]

ECMWF

FUB

extreme years (5-year intervals)

Year

Figure 4-3. Evolution of VPSC for

the Arctic over the past four decades, composed from ECMWF

and FUB data. The blue dots represent the maximum values of

VPSC during five-year intervals.

The dotted line is based on

radiosonde analyses of the FUB,

and the solid line is ECMWF ERA15 data extended by operational

analyses. The gray shading represents the uncertainty of VPSC

based on a 1-K uncertainty of the

long-term stability of radiosonde

temperatures. Updated from Rex

et al. (2004).
chemical processes, the ozone column at polar latitudes

during spring (Randel et al., 2002; Huck et al., 2005).

The breakup dates of the Arctic and Antarctic vortices (Figure 4-4) as determined by the wind average along

the vortex edge vary from year to year in both hemispheres

(WMO, 2003). For the Arctic vortex, no discernible longterm trend in the breakup date appears since the 1960s

(Langematz and Kunze, 2006). The period between the

mid-1980s to the mid-1990s however showed a tendency

toward later breakdown dates. These years were characterized by a reduced dynamical variability, with nine consecutive winters without major stratospheric warmings

from 1989/1990 to 1997/1998 (Labitzke et al., 2002;

Manney et al., 2005c). In contrast, since 1998/1999, the

Arctic exhibited a clustering of warmer than average and

dynamically disturbed winters. Major warmings took place

in the winters 1998/1999, 2000/2001, 2001/2002,

2002/2003, 2003/2004 and 2005/2006, leading to weak or

split vortices. Consistent with these changes in dynamical

variability, the vortex breakup date varied between midFebruary 2001 and early May 1997, with three late

breakups since 2002. However, this recent delay in the

breakup date is not statistically significant and does not

indicate a trend (Langematz and Kunze, 2006). The

described alternation of cold (1989/1990-1997/1998) and

warm (1998/1999-2005/2006) clusters may be interpreted

as shifts between warmer/weaker and colder/stronger

vortex regimes (Perlwitz and Graf, 2001; Perlwitz and

Harnik, 2003). These shifts may be caused by anthropogenic processes (i.e., ozone depletion and greenhouse

gases) that nonlinearly force natural modes of variability

(Corti et al., 1999; Pawson et al., 1998; Christiansen, 2003).

Such a clustering also could be a random occurrence, as

shown in model simulations by Taguchi and Yoden (2002).

The Antarctic vortex exhibited a delay of the

breakup date from late November into mid-December

which is consistent with an intensification of the vortex

(Huth and Canziani, 2003; Renwick, 2004) caused by additional radiative cooling following strong Antarctic ozone

loss in spring (Shine, 1986; Jones and Shanklin, 1995).

However, since the late 1990s, there is larger variability in

the breakup date over Antarctica with the process taking

place sometime between the second half of November

(1997, 2000, and 2002) and late December (1999, 2001).

The tendency toward a later vortex breakup date has

stopped and perhaps reversed (Langematz and Kunze,

2006). This means that, compared with the 1960s and

1970s, the Antarctic vortex is still more persistent; however, compared with the 1980s/early 1990s, the Antarctic

vortex broke down earlier in recent years. This is probably the result of the large interannual variability in recent

Antarctic winters and springs (see Section 4.1.2).

Summing up, since the previous Assessment

(WMO, 2003) changes in the interannual behavior of both

polar vortices have been observed. The Artic vortex continues to show large variability with a shift from a period

of dynamically undisturbed winters between 1989/1990

and 1997/1998 to a period of more dynamically disturbed

winters from 1998/1999 to 2005/2006. The Antarctic

vortex has shown an enhanced variability since the late

1990s compared with previous decades, which culminated

in the anomalous vortex split in 2002. The reasons for the

observed changes in the interannual variability in both

hemispheres are not yet understood.

POLAR OZONE

4.8

300

320

340

360

380

400

Day No

Reanalysis

Reanalysis 2

CPC

SH

Nov

Dec

Jan

1950 1960 1970 1980 1990 2000 2010

40

60

80

00

20

Day No

NH

Feb

Mar

Apr

1950 1960 1970 1980 1990 2000 2010

Year

Figure 4-4. Vortex breakup dates from 1958 to 2005

on the 500-K isentropic surface over the NH (top) and

the SH (bottom) using the method described in Nash

et al. (1996). The filled lines with dots are the dates

calculated from the NCEP/NCAR reanalysis data.

The thick lines are the same data, time-filtered to

remove interannual fluctuations. The crosses are the

breakup dates calculated from NCEP/Climate

Prediction Center analyses for the period 1979-2005.

Updated from Figure 3-10, WMO (2003).
4.1.1.3 TRANSPORT

Besides chemistry, transport is the main driver

for ozone variability at high latitudes. Chemical and

transport-related causes for polar ozone variability are

closely coupled, particularly in the Arctic. Transport

contributions to polar ozone variability are closely controlled by the strength of the middle atmospheric

residual circulation that, in turn, is driven by tropospheric (planetary and gravity) wave forcing; strong

winter planetary wave forcing leads to larger transport

of ozone to high latitudes due to a stronger residual circulation and a higher-than-average Arctic lower stratosphere temperature and a weaker vortex in early spring,

whereas weak winter forcing leads to less transport and

lower Arctic temperatures in spring and to a stronger

vortex (Fusco and Salby 1999; Newman et al., 2001).

For the Arctic, Rex et al. (2004) find that for the current halogen loading, about one-half of the year-to-year

variability in total ozone is caused by variability in

anthropogenic chemical ozone loss and one-half by

variability in transport with the variable residual circulation. Because temperatures influence chemistry, both

terms are correlated.

Although the variability is larger in the Arctic compared with the Antarctic, wave activity,has been shown to

correlate well with the observed change of polar column

ozone over the course of winter for both hemispheres

(Randel et al., 2002; Weber et al., 2003; see also Figure 4-5).

It has been shown that in the Antarctic, this is due to the same

co-action of transport and chemistry as in the Arctic.

Newman et al. (2004) and Huck et al. (2005) showed that

the year-to-year variability of Antarctic ozone depletion is

driven by the year-to-year variability of the polar vortex temperatures. Huck et al. (2005) also showed that interannual

variability in Antarctic stratospheric temperatures and ozone

loss are affected by midlatitude wave activity, while the

decadal increase in Antarctic ozone loss is driven by the

increase in halogens.

In climate models, planetary wave activity has been

shown to partially depend on sea surface temperatures,

among other forcings (Schnadt et al., 2002). It is therefore susceptible to climate change. The external factors

determining stratospheric wave forcing and its interannual

variability (i.e., stochastic tropospheric dynamics, quasibiennial oscillation (QBO), sea surface temperatures

(SSTs), solar variability, and volcanic eruptions) are difficult to quantify; Chapter 6 discusses the impact of these

uncertainties for the predictability of future changes of

polar ozone.

Besides the systematic transport of ozone to high

latitudes with the residual circulation, another important

aspect of transport is mixing. Although mixing across the

vortex edge is relatively weak, it has an effect on the distribution of ozone within the polar vortices. The Antarctic

vortex is divided into two distinct regions of approximately

POLAR OZONE

4.9

Figure 4-5. Increase of ozone over the

winter as a function of planetary wave

activity. Planetary wave activity is measured as winter eddy heat flux in the

respective hemisphere. The ozone ratio

of the zonal mean between 50° and 90°

latitudes from March over that from

September is shown for the NH, and

September over March for the SH.

Circles are NH values, and triangles represent SH values. Total ozone is from

weighting function differential optical

absorption spectroscopy (WFDOAS) V1

(Weber et al., 2005). The monthly mean

eddy heat flux at 100 hPa averaged from

40 to 75 degrees latitude is derived from

ECMWF ERA40/operational analysis,

and has been averaged from fall to

spring in each hemisphere. SH winter

eddy heat flux is negative, but is shown

here as absolute values. Updated from

Weber et al. (2003).
equal area: a strongly mixed vortex core and, separated

from the core, a broad ring of weakly mixed air extending

to the vortex boundary (Lee et al., 2001). This division

was deduced from an analysis of a mixing diagnostic (the

so-called “equivalent length”) based on transport calculations of an artificial tracer field on an isentropic surface

(480 K) for the period from July to November. A transport barrier within the Antarctic vortex between July and

November, is likewise apparent for the altitude range

between 475 and 650 K in the potential vorticity field

(Tilmes et al., 2006b) and, moreover, extends to the setup

phase of the Antarctic vortex in March and April (Tilmes

et al., 2006a). During March and April, ozone-tracer relations in the two vortex regions are clearly distinct (Tilmes

et al., 2006a). Further, measurements of ozone and volcanic aerosol show rather different vertical profiles within

the vortex core and in the edge region (Godin et al., 2001).

A better understanding of the dynamical mechanisms

responsible for the occurrence of transport barriers within

the vortex is important, because temperature variations in

the weakly mixed air mass between the vortex boundary

and the vortex core drive the year-to-year variability of

POLAR OZONE

4.10

Figure 4-6. March monthly averaged total ozone for the Arctic. The 1971 and 1972 images are from the

Nimbus-4 Backscatter Ultraviolet (BUV) instrument; the 1979 and 1980 images are from the Nimbus-7 TOMS

instrument; the 1996 image is from the NOAA-9 SBUV/2 instrument; the 1997, 2000, 2002, 2003, and 2004

images are from the Earth Probe TOMS instrument; and the 2005 and 2006 images are from the Aura Ozone

Monitoring Instrument (OMI). This figure is updated from Figure 7-21 of WMO (1999).
about 10-20% in the area of the ozone hole (measured as

the area enclosed in the 220-DU contour) (Newman et al.,

2004).

4.1.2 Polar Ozone

This section presents observations of ozone and

diagnoses of ozone destruction in both polar regions.

Section 4.1.2.1 gives an update of observations of ozone

and the primary chlorine species involved in stratospheric

ozone destruction, along with a discussion of various indicators that capture different aspects of ozone hole area and

severity. A discussion of other trace gas species such as

BrO and chlorine dioxide (OClO) can be found in Section

4.2.1 and Chapter 2 (Section 2.5.1), and nitrogen dioxide

(NO2) is discussed in Chapter 3 (Section 3.3.2). In Section

4.1.2.2, refinements since the previous Assessment

(WMO, 2003) of various methods for quantifying chemical ozone loss in the Arctic polar vortex are presented.

Near-record ozone losses during the Arctic 2004/2005

winter are discussed. Finally, recent progress in quantifying ozone loss rates in the Antarctic is discussed.

4.1.2.1 OZONE AND OTHER CONSTITUENTS

The ozone content in the Arctic stratosphere is

dependent on chemical and dynamical conditions and

shows great interannual variability. This section discusses

resulting interannual variability of ozone and other constituents, which is due to the combined effect of transport

and chemistry. Section 4.1.2.2 isolates the impact of

anthropogenic chemical loss on polar ozone variability

from the variability in transport.

As noted in the previous section, the formation, evolution, and breakup of the polar vortex and the occurrence

of stratospheric warmings are subject to large year-to-year

variations. Figure 4-3 shows that the variability in amplitude of VPSC (PSC formation potential derived from stratospheric temperature analyses) in the Arctic has clearly

increased over the last 40 years, with 1995/1996,

1999/2000 and 2004/2005 standing out as cold winters

with high PSC formation potential, whereas 1998/1999,

2001/2002, and 2003/2004 were warm winters with very

low PSC formation potential. As discussed in the previous

section, this variability in dynamical conditions has a correlated impact on ozone transport and chemistry and is

therefore reflected in the total ozone column over the polar

regions.

Total Ozone Average

Arctic total ozone has remained low over the last

few years in comparison with pre-1982 observations.

Figure 4-6 displays a series of March averages for selected

years from 1971 to 2006 (updated from Figure 7-21 in

WMO, 1999). The 60°N latitude circle generally encloses

the region of low ozone, but in some years (e.g., 2005) the

vortex and low-ozone region are displaced from the pole,

extending southward of 60°N. Nevertheless, Arctic ozone

levels during March of recent years are low in comparison

with the observations prior to 1980, as shown in the top

row of Figure 4-6.

The springtime average total ozone values in the

Arctic and Antarctic poleward of 63° latitude are shown

in Figure 4-7 in comparison with the average total ozone

for the years 1970-1982 (gray lines). The difference

between the observed values and the 1970-1982 average

indicates the combined changes in ozone due to chemistry

and dynamics. In the last 9 years, Arctic column ozone

exceeded the low values of the mid-1990s, except in the

cold and chemically active winter of 1999/2000, when a

large decrease of 63°-90° NH total ozone was observed.

The record-cold winter of 2004/2005, however, showed a

less pronounced impact on the March polar average total

ozone. Although NH polar column ozone averages are in

general a good indicator of Arctic ozone depletion (WMO,

2003), the March average in 2004/2005 reflects the strong

influence of dynamics (e.g., vortex fragments moved outside the 63°-90° region during March, see Figure 4-6) and

is consequently high relative to those of other recent cold

winters even though the magnitude of chemical ozone

POLAR OZONE

4.11

63o-90o total ozone average

1970 1980 1990 2000

Year

200

250

300

350

400

450

500

Total ozone (DU)

BUV

Nimbus-7

Meteor-3

EP-TOMS

NOAA-14

NOAA-9

Aura-OMI

SH October

NH March

Figure 4-7. Total ozone average of 63°-90° latitude

in March (NH) and October (SH). Symbols indicate

the satellite data that have been used in different

years. The horizontal gray lines represent the

average total ozone in March for the years prior to

1983 for the NH and SH. Updated from Figure 3-2,

WMO (2003).
loss in the lower stratospheric vortex was as high as or

higher than in those years (cf. Section 4.1.2.2).

In the SH polar region, the lowest average October

total ozone values were observed in the late 1990s (Figure

4-7, bottom panel). The last 6 years have shown an

increase in polar column ozone averages, with a high

degree of variability and notably higher ozone values in

2000, 2002, and 2004 resulting from greater dynamical

activity.

To aid the detection of the recovery of the Antarctic

ozone hole (discussed in detail in Chapter 6), various metrics that capture different aspects of the ozone hole are

used to describe the severity of ozone depletion, such as

ozone hole area, ozone minimum, ozone mass deficit, and

date of ozone hole appearance and disappearance. Figure

4-8 displays ozone hole area (top panel), ozone hole minimum (middle), and ozone mass deficit (bottom).

Following Newman et al. (2006), we have added a fit of

these metrics (gray line) to a quadratic function of

Antarctic equivalent effective stratospheric chlorine

(EESC) (see Box 8.1 for a description of Antarctic EESC).

Ozone Hole Area

A primary estimate of the severity of the Antarctic

ozone hole is its geographic area (WMO, 2003). Figure

4-8 (top panel) shows the ozone hole area calculated from

the area contained by total column ozone values less than

220 DU during 21-30 September from Total Ozone

Mapping Spectrometer (TOMS) and Ozone Monitoring

Instrument (OMI) observations. The value 220 DU was

chosen to define the ozone hole because it is almost always

a middle value in a strong gradient of total ozone, and

because it is lower than pre-1980 observed ozone values.

The average ozone hole area currently reaches approximately 25 million km2 each spring, while the single largest

daily value reached nearly 30 million km2 in September

2000 (Newman et al., 2004). Ozone depletion can be

enhanced by volcanic perturbations (e.g., Hofmann and

Oltmans, 1993) as could be seen in the very strong ozone

depletion (or “deep ozone holes”) observed in the 1990s

after the Mt. Pinatubo eruption (e.g., 1994). The area

growth of the ozone hole slowed during the mid-1990s,

with a large dip in 2002.

Ozone Minimum

The daily total column ozone minimum value is a

widely used measure of the state of the ozone hole.

Average minimum ozone columns calculated over

Antarctica for the period 21 September to 16 October show

a clear decrease from 1979 to the mid-1990s, with the

POLAR OZONE

4.12

0

10

20

30

40

100

150

200

0

10

20

30

1980 1990 2000 2010

Area

Minimum

Mass Deficit

Dobson Units Million Tons Million km2

Year

Figure 4-8. Top panel: Area of the Antarctic ozone

hole for 1979-2005, averaged from daily total ozone

area values contained by the 220-DU contour for 21-

30 September. The gray bars indicate the range of

values over the same time period. The gray line

shows the fit to these area values as was shown in

Newman et al. (2004), and now using EESC, as

derived in Newman et al. (2006) (see also Box 8.1 in

Chapter 8). The EESC has a mean age of 5.5 years,

an age spectrum width of 2.75 years, and a brominescaling factor of 50. The fit is quadratic in EESC.

Middle panel: An average of daily minimum ozone

values over Antarctica during the period from 21

September to 16 October. The gray bars indicate the

range of ozone values over the same time period.

The gray line is the fit as described for the top panel.

Bottom panel: Ozone mass deficit (OMD) average

over the 21-30 September period. The gray bars indicate the range of values over the same time period.

The gray line is the fit as described for the top panel.

This figure was generated using the merged ozone

data courtesy of Greg Bodeker (NIWA) and

NCEP/NCAR reanalysis 2 data. Updated from

Figures 3-5 and 3-7 from WMO (2003).
lowest minimum value observed in 1994 (Figure 4-8,

middle panel).

Ozone Mass Deficit

The ozone mass deficit (OMD) combines the effects

of changes in ozone hole area and depth, and provides a

direct measure of the deficit relative to the mass present

for a value of 220 DU (e.g., Uchino et al., 1999). Figure

4-8 (bottom panel) shows the OMD averaged over the

period 21-30 September. While the long-term evolution

of the OMD follows the halogen loading, there is higher

frequency year-to-year variability; years with anomalously

high wave activity (1988, 2002, and 2004) show weak

Antarctic ozone depletion, and years with suppressed wave

activity show severe depletion (Huck et al., 2005, and references therein).

Daily values of OMD over Antarctica for the years

2002-2005 are compared with the range of values over

the period 1990-2001 in Figure 4-9. This plot shows

that, apart from 2002, the year 2004 also stands out as

having a weak ozone hole, while the OMD in the 2003

and 2005 Antarctic winters is more comparable to those

observed during the 1990s. Although lower stratospheric (50 hPa) minimum temperatures were below

average over much of the 2004 Antarctic winter, they

increased and remained near average after mid-August

(Santee et al., 2005). The lower stratosphere warmed

rapidly in September, halting further heterogeneous processing of vortex air by the end of the month. This

resulted in a slow increase in OMD and a leveling off in

late September (see Figure 4-9). At the end of September, a large increase in mixing accompanied by a weakening vortex transport barrier signaled vortex erosion

leading to the breakup (Manney et al., 2005b).

Hoppel et al. (2005b) used measurements from both

the Polar Ozone and Aerosol Measurement (POAM)

instrument and South Pole ozonesondes to investigate

changes in the ozone mixing ratio near the top (20-22 km)

of the ozone hole, comparing the years 2001-2004 with

the previous six years of POAM observations (1994-1996;

1998-2000). They found that in this altitude range, the

year 2004 was distinguished by the highest minimum

vortex temperatures throughout August and September,

the lowest observed PSC frequency, and the smallest

photochemical ozone loss.

Chlorine Partitioning

The evolution of chlorine partitioning in the

winter polar vortices of both hemispheres is shown in

Figure 4-10. In this figure, contemporaneous measurements from two recent satellite missions are used to

update the conceptual sketch of wintertime chlorine parPOLAR OZONE

4.13

Ozone mass deficit over Antarctica (million tons)

2005

2004

2003

2002

45

0-Jul 29-Aug 8-Sep 8-Oct 27-Nov

10

15

20

25

30

35

40

27-Dec

Figure 4-9. Daily

OMD over Antarctica for the years

2002-2005 compared with the range

of values over the

period 1990-2001.

The grayed area

shows the daily

range from 1990 to

2001. Adapted from

Bodeker et al.

(2005).
titioning presented in a previous WMO Assessment

(WMO, 1995; Figure 3-1). Although several studies

based on satellite, aircraft, balloon, and ground-based

measurements over the last decade have provided insight

into both the relative abundances of the main chlorine

reservoirs prior to activation and the interhemispheric

differences in their recovery at the end of winter, the

new data from the Atmospheric Chemistry Experiment

Fourier Transform Spectrometer (ACE-FTS) (e.g.,

Dufour et al., 2006) and the Aura Microwave Limb

Sounder (MLS) (e.g., Santee et al., 2005) are of unprecedented scope for studying seasonal changes in chlorine

partitioning. In particular, the dominant chlorine reservoir before the onset of PSC processing is confirmed to

be HCl, not ClONO2 as indicated in the schematic shown

in WMO (1995).

In the Antarctic (Figure 4-10, left panel), chlorine is

converted into the active form ClO by early June; elevated

abundances of ClO then persist through mid-September.

In the severely denitrified and ozone-depleted conditions

characteristic of late Antarctic winter, HCl production is

highly favored and ClONO2 production is suppressed.

In the Arctic (Figure 4-10, right panel), the magnitude and duration of chlorine activation are much

smaller than the Antarctic, even in a relatively cold

winter. The ACE-FTS and Aura MLS measurements

clearly support the canonical picture of initial chlorine

deactivation in the Arctic, with the primary pathway the

reformation of ClONO2, followed by slow repartitioning

between ClONO2 and HCl.

4.1.2.2 POLAR OZONE LOSS STUDIES

Quantifying Arctic Ozone Loss

Diagnosing chlorine-catalyzed ozone destruction

requires distinguishing the effects of chemical processes

from those of transport and mixing, especially in the

Arctic, where ozone abundances are strongly controlled

by dynamics. WMO (2003) presented an overview of the

methods commonly used to estimate chemical ozone loss

and discussed the limitations and sources of uncertainty

of each, as well as the degree of consistency between them.

Most of these techniques have been significantly refined

since the previous Assessment, with a particular emphasis

POLAR OZONE

4.14

Jun Jul Aug Sep

0

2

3

Chlorine Species / ppbv

70-75 EqL

Dec Jan Feb Mar

490 K

MLS ClO

MLS HCl

ACE ClONO

Southern Hemisphere Northern Hemisphere

Figure 4-10. Left panel: Daily averages of ClO (red dots) and HCl (green dots) observed by Aura MLS, and

ClONO2 (cyan triangles) observed by ACE-FTS at 490 K (~20 km) during the 2005 Antarctic winter/spring

period, calculated over the 70°-75° equivalent latitude band using the Global Modeling and Assimilation Office

Goddard Earth Observing System (GEOS-4) temperatures and potential vorticity. Only daytime measurements are included in the averages for ClO; ClO data appear sparser because measurements in sunlight are

not always available at high equivalent latitudes, especially in early winter. The sampling pattern of ACE-FTS,

a solar occultation instrument, does not provide coverage of this equivalent latitude band at all times throughout

the winter. Right panel: As in the left panel, for the 2004/2005 Arctic winter/spring period.
on quantifying the uncertainties in the loss estimates. This

section presents an update of recent progress in separating

chemical and dynamical effects and quantifying ozone

destruction, mainly in the context of the 2002/2003 Arctic

winter. In 2002/2003, an unusually cold early winter promoted an earlier inception of significant ozone destruction than in most other recent cold winters (e.g., Goutail et

al., 2005; Streibel et al., 2006). The mid- to late winter

was very dynamically active and warmer than average,

but with a series of stratospheric warming events that limited cumulative ozone loss. Aircraft and ground field campaigns in the 2002/2003 winter provided an excellent

opportunity to assess understanding of polar ozone loss

processes, especially under the high solar zenith angle conditions of early winter. The different approaches to calculating chemical ozone loss generally agree well, lending

confidence in the estimates.

Ozone/Tracer Correlation Method. In this technique, changes in the compact correlations between ozone

and long-lived tracers over the winter are ascribed to

chemical loss. This approach assumes that a reference

ozone/tracer relation representative of the early-winter

polar vortex can be established and that transport and

mixing processes do not significantly alter it, assumptions that had been questioned at the time of the previous

Assessment (e.g., Harris et al., 2002; WMO, 2003; Sankey

and Shepherd, 2003). The technique was refined in recent

applications (Müller et al., 2002; Tilmes et al., 2003), and

a comprehensive re-evaluation of ozone loss in the Arctic

winters 1991-2003 was performed (Tilmes et al., 2004).

Ozone loss estimates generally agreed well with those

based on other approaches, with the smallest estimated

errors obtained in the coldest winters with greatest loss;

the major contribution to the error in the calculated loss

remained the uncertainty in the reference relation. Müller

et al. (2005) confirmed the validity of ozone-tracer relations for quantifying chemical loss, provided that a reliable early vortex reference is obtained and vortex and

midlatitude air masses are well separated by a strong

transport barrier at the vortex edge.

Match Method. “Match” is a pseudo-Lagrangian

technique involving identification (through trajectory calculations) of air parcels sampled two or more times in a

prescribed interval; differences in the observed ozone

mixing ratios are attributed to chemical loss. Match has

received substantial attention since the previous

Assessment. Morris et al. (2005) revisited the JanuaryMarch periods in 1992 and 2000 using a different trajectory code and meteorological analyses than in the original

Match studies (Rex et al., 1998, 2002). They confirmed

the statistical error bars published previously, but concluded that actual uncertainties in the calculated ozone loss

rates are larger because of additional systematic errors,

consistent with earlier discussions of systematic Match

uncertainties (e.g., Rex et al., 1998; Harris et al., 2002).

Morris et al. (2005) recommended that future Match studies

employ extended (14-day) trajectory calculations and suggested that some of the filter criteria built into the approach

may be unnecessary. In contrast, Grooß and Müller (2003)

carried out a virtual Match campaign in the Chemical

Lagrangian Model of the Stratosphere (CLaMS) threedimensional (3-D) chemical transport model and found that

these filter criteria reduce the statistical uncertainty of the

estimated ozone loss rates and remove a systematic bias

induced by mixing across the vortex edge, particularly

during periods of strong vortex disturbances.

Lehmann et al. (2005) developed a statistical

approach to take into account the error correlation between

matches that share a common ozone measurement and

found that it depends nearly linearly on the “oversampling

rate” (i.e., the average number of matches to which an

ozone measurement contributes). For ozonesonde Match

studies, the oversampling rate is low and the average

increase in the error bars is 15%. For satellite Match

studies, however, the oversampling rate is typically much

higher and uncertainties increase by more than 50% compared with those estimated assuming uncorrelated errors.

This updated approach to estimating uncertainties was

used for the 2002/2003 Match study (Streibel et al., 2006),

which also included the use of higher-resolution wind

fields for the trajectory calculations, the imposition of a

backward match radius in addition to the criterion along

the forward trajectory, and the earliest start of a Match

campaign, allowing investigation of ozone loss starting

in December.

Vortex-Average Method. In this method, the difference between an initial profile altered using diabatic

descent rates (estimated using a radiation code) and a final

observed profile yields an estimate of chemical loss.

Grooß and Müller (2003) used CLaMS simulations to

quantify the influence of intrusions of midlatitude air into

the vortex on ozone loss estimates derived with this

method and showed that, although cross-vortex transport

may often be insignificant, in some dynamically active

periods neglecting it may introduce significant error in the

estimated ozone loss rate. Christensen et al. (2005)

applied the vortex-average method, corrected to account

for transport into the vortex, to ozonesonde data from the

2002/2003 Arctic winter and compared the results with

ozone loss estimates from other techniques, finding generally good agreement. A variant on the vortex-average

method uses the evolution of a long-lived tracer, rather

than trajectory calculations and a radiation code, to deduce

diabatic descent. For example, Odin SubMillimeter

POLAR OZONE

4.15
Radiometer (SMR) measurements of nitrous oxide (N2O)

were used to account for the effects of subsidence in estimating chemical loss during the 2002/2003 winter from

both Odin/SMR (Urban et al., 2004) and ground-based

millimeter-wave (Raffalski et al., 2005) ozone observations. This approach has been greatly facilitated by the

availability of simultaneous global ozone and tracer measurements from recent satellite missions such as Odin,

Envisat, and Aura (e.g., Manney et al., 2006).

Passive Subtraction Method. In this technique,

differences between measured ozone and results from

model runs in which ozone is treated as a passive dynamical tracer are assumed to arise from chemical changes.

Recent refinements in 3-D chemical transport models (see

Section 4.2.1.3) have improved estimates of chemical loss

derived in this manner. Feng et al. (2005b) showed that

seasonal runs of the updated Single-Layer Isentropic

Model of Chemistry and Transport (SLIMCAT) model

(Chipperfield, 2006) produced realistic representations of

tracer transport and ozone loss for selected warm and cold

Arctic winters, including 2002/2003. Goutail et al. (2005)

compared ground-based total ozone columns from the

2002/2003 winter with modeled passive ozone columns

from the Reactive Processes Ruling the Ozone Budget in

the Stratosphere (REPROBUS) and updated SLIMCAT

models; although both models indicated significant earlywinter loss, full-chemistry simulations from REPROBUS

did not reproduce all of the inferred loss. Singleton et al.

(2005) compared the updated SLIMCAT with Polar

Ozone and Aerosol Measurement (POAM) III ozone

observations from 2002/2003, examining a “pure passive”

ozone advected as an inert tracer with no chemical

changes and a “pseudo passive” ozone for which only heterogeneous reactions on PSCs were switched off in order

to explore the influence of gas-phase chemistry on

inferred ozone loss. The timing, morphology, and magnitude of modeled (full chemistry) ozone loss agreed well

with that deduced from POAM observations, with the

largest differences (<1 parts per million by volume

(ppmv)) in the two passive ozone fields, attributed to

nitrogen oxides (NOx) chemistry in the “pseudo passive”

run, occurring above 450 K. Finally, Grooß et al. (2005b)

showed that ozone loss inferred using CLaMS compared

well with other estimates.

Lagrangian Transport Calculation Method. A

method frequently applied to ozone measurements from

the Upper Atmosphere Research Satellite (UARS)

Microwave Limb Sounder (MLS) involves the use of a

Lagrangian transport (LT) model advecting trace gases

passively along parcel trajectories; departures from the

observed ozone provide an estimate of chemical loss.

Previous studies covered only the late-winter period for a

subset of Arctic winters observed by UARS MLS and used

different LT models and MLS data versions. Manney et

al. (2003a) performed a comprehensive reanalysis that

encompassed the entire Arctic winter for all years with

sufficient data and used a consistent LT model and the

definitive UARS MLS dataset, facilitating evaluation of

interannual variability.

Chemical Loss in the 2004/2005 Arctic Winter

The potential for severe ozone depletion will exist

so long as stratospheric chlorine loading remains well

above natural levels. The degree of chlorine activation and

consequent ozone destruction in the Arctic is primarily controlled by dynamical variability. Winters in which the polar

vortex is cold and isolated are expected to experience large

ozone losses, whereas ozone losses should be minimal in

warmer than average, disturbed winters. The 2004/2005

Arctic winter was especially interesting in this context.

The lower stratosphere was exceptionally cold, particularly below 400 K (~15 km), with temperatures below PSC

existence thresholds on more days and over a broader

region than previously observed (Manney et al., 2006; Rex

et al., 2006). But the lower stratosphere was also dynamically active, with frequent intrusions of lower-latitude or

vortex edge air into the vortex interior throughout the

winter (Manney et al., 2006; Schoeberl et al., 2006).

Moreover, the initial morphology of ozone in early winter,

prior to the onset of chemical loss, exhibited a strong gradient in mixing ratio between low values in the vortex core

and high values at the edge (Manney et al., 2006; Rex et

al., 2006). Thus mixing processes may have masked or

mimicked chemical loss, depending on the location in the

vortex. A major final warming in early-March 2005 halted

PSC processing, terminating ozone destruction earlier than

in other recent cold winters. These effects made disentangling chemical and dynamical processes particularly difficult this year, possibly causing some analyses to be more

sensitive to sampling biases and increasing uncertainties in

the estimated losses. Nevertheless, a suite of studies using

a variety of datasets and techniques provided fairly consistent ozone loss estimates for this winter.

Manney et al. (2006) deduced chemical ozone loss

from Aura MLS measurements by using N2O to account for

the effects of diabatic descent in the vortex edge and core

regions, and estimated loss of up to 2 ppmv in the outer

vortex and ~1.5 ppmv averaged over the entire vortex, with

maximum losses in both regions between 450 and 500 K.

Manney et al. (2006) also applied the vortex-average

method, using trajectory calculations and a radiation code,

to both Aura MLS and POAM measurements (Figure 4-11)

and found maximum vortex-averaged chemical loss near

POLAR OZONE

4.16
POLAR OZONE

4.17

colors indicate year: 1996, 2000, 2005

thick lines: Manney et al., 2006

MLS

POAM

hin lines: Rex et al., 2002; 2006

ozonesondes

SAGE

POAM

Match

Ozone Loss (1012 Ozone Loss (ppmv) molec. cm-3

0 1 2 30 6 2 4

350

400

450

500

550

350

400

450

500

550

Potential Temperature (K) Potential Temperature (K)

12

4

16

8

20

22

24

12

4

16

8

20

22

24

approx Altitude (km) approx Altitude (km)

1996

2000

2005

early March early March

late March late March

Figure 4-11. Profiles of chemical ozone loss in terms of mixing ratio versus potential temperature (left panels) and

concentration versus altitude (right panels), over the intervals from 5 January to 10 March (top panels) and 5

January to 25 March (bottom panels) for three cold Arctic winters, estimated using various data sources and techniques. Although different line styles and symbols have been used to denote the different datasets and approaches,

in most cases it is not necessary to clearly distinguish the curves, as the derived ozone loss values are generally in

quite good agreement. All lines represent results from the vortex-average descent method. Thick lines without

symbols are from Manney et al. (2006), with solid lines based on POAM II/III and dotted lines based on UARS MLS

(1996) and Aura MLS (2005). Thin lines with small symbols are from Rex et al. (2002) and Rex et al. (2006), with

lines marked by filled dots based on ozonesonde data, lines marked by open circles based on SAGE III data, and

lines marked by open squares based on POAM II/III data. Large square symbols (connected by dashed lines)

show results from Match (Rex et al., 2002, 2006). Ozone loss in 2005 terminated earlier and was smaller in terms

of mixing ratios compared with 2000. But, because the 2005 loss extended to lower altitudes where ozone concentrations are high, it was larger in terms of concentration, which is the relevant quantity for total column loss.
450 K of 1.2-1.3 ppmv, slightly less than the estimates based

on MLS N2O. These results imply that, in terms of mixing

ratios, chemical loss in 2004/2005 was slightly less than

that in 1999/2000 and comparable to that in 1995/1996, but

with maximum loss occurring at a lower altitude.

Singleton et al. (2006) used the updated SLIMCAT

model to infer chemical ozone loss from POAM III, Aura

MLS, ACE-FTS, Measurement of Aerosol Extinction in

the Stratosphere and Troposphere Retrieved by Occultation (MAESTRO), and Stratospheric Aerosol and Gas

Experiment (SAGE) III measurements; all instruments

provided similar results, with a maximum inferred loss of

2-2.3 ppmv near 450 K (Figure 4-12). Jin et al. (2006)

also examined ACE-FTS measurements, using correlations between ozone and methane (CH4) (modified to

account for mixing processes), correlations between ozone

and an artificial tracer constructed to have a linear correlation with ozone in early winter, and the vortex-averaged

descent method to diagnose an average maximum loss of

~2.1 ppmv near 500 K, with an average column loss of

119 DU over the region 375-800 K (14-30 km).

Results from Match (Rex et al., 2006; Figure 4-11)

indicated chemical loss characterized by a broad peak of

~1.5 ppmv from ~400-450 K, smaller than the maximum

mixing ratio loss in 1999/2000, but with more loss occurring at lower altitudes. However, Rex et al. (2006) showed

that, in terms of concentration (rather than mixing ratio),

ozone loss in 2004/2005 was considerably more severe

than that in 1999/2000, particularly below about 16 km.

For the total column and, hence, the ultraviolet (UV)

reaching the surface, the large concentration losses at

lower altitude have the biggest impact. Therefore, total

column losses during 2004/2005 were at least on the order

of the largest losses recorded previously (Rex et al., 2006).

The 2004/2005 winter stands out as having the

largest VPSC and partial column ozone loss (∆O3 = 121 ±

20 DU over 380-550 K) of the last 14 years, although the

differences between column ozone losses in 1995/1996,

1999/2000, and 2004/2005 are within the uncertainties in

the calculation. The compact and nearly linear relationship between the influence of chemical ozone loss on ∆O3

and the winter average of VPSC reported by Rex et al.

(2004) and confirmed by Tilmes et al. (2004) was maintained in 2004/2005 (Figure 4-13; Rex et al., 2006). In

contrast, comparisons of ground-based total ozone measurements with passive ozone from REPROBUS (e.g.,

Goutail et al., 2005) indicated that the total column loss in

2004/2005 rivaled that in 1999/2000 but was smaller than

that in 1995/1996. The various partial and total column

ozone loss estimates for the 2004/2005 Arctic winter discussed here are compared with those available from previous years in Figure 4-14.

The March 2005 averaged total ozone in Figure 4-6

clearly shows a distinct minimum (near 60°N) that is comparable to that in other years with strong depletions. As

discussed in Section 4.1.2.1, total column ozone (data from

the OMI on Aura; Figure 4-7) averaged over the Arctic

region in March 2005 was below normal but comparable

to that in other recent winters, even though the magnitude

of chemical ozone destruction in the lower stratospheric

vortex was substantially larger. The March 2005 average

in Figure 4-6 clearly shows a distinct minimum (near 60°N)

that is comparable to other years with strong depletions.

Several factors account for these apparent discrepancies.

POLAR OZONE

4.18

 D J F M

400

450

500

550

600

650

700

Potential Temperature (K)

 D J F M

400

450

500

550

600

650

700

Potential Temperature (K)

POAM III 

 D J F M

SAGE III

 D J F M

EOS MLS

 D J F M

ACE-FTS

 D J F M

MAESTRO

-2.3

-2.0

-1.8

-1.5

-1.3

-1.0

-0.7

-0.5

-0.2

0.0

0.3

∆ ppmv

Figure 4-12. Daily average chemical ozone loss (ppmv) inside the vortex during the 2004/2005 Arctic winter

inferred by differencing passive ozone calculated by the updated SLIMCAT chemical transport model, and

ozone measured by the POAM III, SAGE III, Aura MLS, ACE-FTS, and MAESTRO instruments. Days on which

data are missing or an instrument did not sample the vortex have been interpolated in time. Data have been

smoothed with a seven-day running average. The solid black line denotes the zero contour.
In previous years with later final warmings, the March

average reflected primarily winter vortex values that, for

dynamical reasons, were much lower than the spring values

dominating the 2005 average. Even before the final

warming, the 2005 vortex had become highly distorted and

shifted off the pole (Manney et al., 2006). The monthly

average over the entire Arctic region, therefore, encompassed air from both inside and outside the vortex; together

with the vigorous mixing following the vortex breakup,

this effect muted the signature of chemical depletion in the

average ozone north of 63°N. As a consequence, the

average total column ozone over the Arctic in March 2005

was strongly influenced by dynamics and was not representative of chemical loss inside the polar vortex remnants.

As a result of dynamical processes during the

warming in combination with the large chemical depletion, total ozone over large parts of Europe was very low

(reaching values below 250 DU) for several days during

late March of 2005. This led to elevated levels of UV at

the ground over Europe during those days.

Quantifying Antarctic Ozone Loss

Over the past two decades, numerous studies have

compared modeled and measured ozone in the Antarctic

ozone hole region. Although in most cases the models

were shown to be successful in simulating the general features of ozone hole development, such as the timing of

the onset and the cumulative amount of ozone loss, these

POLAR OZONE

4.19

0

20

40

60

80

100

120

40

Ozone loss [DU]

992

1994

996

1998

2000

2002

2004

992

1994

996

1998

2000

2002

2004

VPSC (106 k m3

Year

0 10 20 30 40 50

1992 1994 1996 1998 2000 2002 2004 2006

0

50

100

150

Year

Ozone column loss [DU]

Total column:

SAOZ

column above ~380K:

ozonesondes

HALOE

ACE-FTS

Match

Figure 4-13. Scatter plot of vortex-average chemical loss of column ozone (∆O3, calculated over the

range 380 to 550 K) versus VPSC inferred from ozonesonde observations for the 1991/1992 to 2004/2005

Arctic winters (update from Rex et al., 2004). Colored

squares and the red fit line show results based on

ozonesonde analyses; colored circles and the green

fit line show results from tracer correlation studies

based on Halogen Occultation Experiment (HALOE)

data (update from Tilmes et al., 2004). Also shown

are results from SLIMCAT model simulations. Open

black circles and the dotted fit line show results from

an older version of the model, run at a resolution of

7.5° for the 1990/1991 to 1999/2000 winters (same

model results as shown in Figure 2 of Rex et al.,

2004). Closed black circles and the dashed fit line

show results from a new version of the model, run at

higher resolution (2.8°) for the 1994/1995 to

2003/2004 winters (same model results as shown in

Figure 1 of Chipperfield et al., 2005). See Section

4.2.1 for descriptions of the old and newly updated

versions of the SLIMCAT model. Adapted from Rex

et al. (2006), Rex et al. (2004), Tilmes et al. (2004),

and Chipperfield et al. (2005).

Figure 4-14. Interannual variation of the ozone

column losses in the Arctic since the early 1990s

based on Système d’ Analyse par Observation

Zénithale (SAOZ) satellite data (blue, update of

Goutail et al., 2005); ozonesonde analyses (red, typical uncertainty illustrated by the single error bar,

update from Rex et al., 2004); HALOE data (green,

Tilmes et al., 2004); Match (purple, Rex et al., 2002,

2006); and ACE-FTS data (orange, Jin et al., 2006).

The sonde, HALOE, and Match analyses represent

partial column losses between 380 and 550 K potential temperature. The other studies represent estimates of the total column losses, which can result in

slightly larger estimates in some winters. Therefore,

the loss estimates from the SAOZ data have been

plotted with a dashed line to signify that they should

not be directly compared with the results from the

other two long-term studies.
did not provide a stringent test of theoretical understanding. Until recently, few studies have undertaken the

kind of detailed quantitative estimations of ozone loss

that have become routine in the Arctic. The inability to

accurately model ozone loss during cold Arctic winters

(e.g., WMO, 2003; Rex et al., 2003), however, has

renewed interest in the Antarctic, where ozone losses are

larger, the vortex is more isolated, and, consequently,

transport and mixing processes are less important in controlling ozone abundances. Several recent studies have

applied methods described above to quantify ozone loss

using measurements from satellites and the first coordinated Antarctic ozonesonde Match campaign, conducted

in 2003.

Hoppel et al. (2005b) used the vortex-average

descent method with 10 years of ozone measurements

from the POAM II and III instruments to deduce chemical loss, with a particular emphasis on the 20-22 km

region near the top of the ozone hole (see 4.1.2.1 for further details). Hoppel et al. (2005a) used the satellite

Match technique to calculate ozone loss rates at four

potential temperature levels over the range ~450-520 K

from five years of POAM III data and compared them

with rates from a photochemical box model. Measured

loss rates at the high solar zenith angles characteristic of

the POAM-Match trajectories were found to increase

slowly from late June to early August, and then increase

rapidly until mid-September. The Match loss rates were

also found to be highly sensitive to the meteorological

analyses used for the trajectory calculations. Frieler et al.

(2006) compared modeled ozone loss rates with those

estimated for the 2003 ozonesonde Match campaign, as

well as several Arctic winters. Detailed discussion of the

implications of both the Hoppel et al. (2005a) and Frieler

et al. (2006) studies for various model parameters is given

in Section 4.2.1.3.

Tilmes et al. (2006b) used the tracer/tracer correlation technique with Improved Limb Atmospheric

Sounder (ILAS)-II measurements to examine the temporal evolution of ozone loss throughout the 2003

Antarctic winter. Chemical loss began in July for all altitudes considered (380-620 K), and the accumulated

ozone loss and loss rates were shown to be highly

dependent on altitude. Ozone loss rates increased

strongly during September throughout the lower stratosphere; half of the entire column ozone loss of 157 DU

occurred during September, with virtually all ozone

between 380 and 470 K destroyed by the end of the

month. Simulations by the CLaMS box model confirmed

that increasing solar illumination and persistent low temperatures led to enhanced ozone loss rates in midSeptember.

4.2 PROGRESS IN OUR UNDERSTANDING OF

THE PHYSICAL AND CHEMICAL

PROCESSES

4.2.1 Polar Ozone Chemistry

The chemical loss of polar ozone during winter and

spring occurs primarily by two gas-phase catalytic cycles

that involve chlorine oxide radicals (Molina and Molina,

1987) and bromine and chlorine oxides (McElroy et al.,

1986):

Cycle 1

ClO + ClO + M → ClOOCl + M (1a)

ClOOCl + hν → 2 Cl + O2 (1b)

2 [Cl + O3 → ClO + O2] (1c)

Net: 2 O3 → 3 O2

Cycle 2

BrO + ClO + hν → Br + Cl + O2 (2a)

Br + O3 → BrO + O2 (2b)

Cl + O3 → ClO + O2 (2c)

Net: 2 O3 → 3 O2

If loss of ClOOCl occurs by thermal decomposition:

ClOOCl + M → ClO + ClO + M (1d)

rather than photolysis (1b), a null cycle results that recycles ClO but leads to no change in ozone. Small contributions to polar ozone loss also occur due to cycles limited

by the reactions ClO + O and ClO + HO2.

Abundances of ClO in the polar vortex are greatly

elevated by reactions of inactive chlorine reservoir species

on various types of PSCs that form when temperatures

drop below about 195 K (see Section 4.2.2). Abundances

of BrO determine the removal rate by cycle (2), which may

contribute almost half of the total chemical loss rate of

Antarctic and Arctic ozone (e.g., Frieler et al., 2006). The

abundance of BrO, in contrast to ClO, is not strongly

affected by reactions involving PSCs (Pundt et al., 2002).

The emerging issue involving BrO, important also for our

understanding of polar ozone loss, is quantification of total

inorganic bromine (Bry) supplied to the stratosphere by

very short-lived (VSL) bromocarbons (see Chapter 2).

Since the previous Assessment, numerous field

and laboratory studies have provided advances in our

POLAR OZONE

4.20
quantitative understanding of ozone destruction by

cycles (1) and (2). These advances include the first

measurements of the abundance of the chlorine

monoxide dimer (ClOOCl) (Stimpfle et al., 2004; von

Hobe et al., 2005); observations of BrO in the polar vortices (Fitzenberger, 2000; Pfeilsticker et al., 2000;

Pundt et al., 2002; Canty et al., 2005; Giovanelli et al.,

2005; Dorf et al., 2006; Sinnhuber et al., 2005b; Frieler

et al., 2006; Schofield et al., 2006; Sioris et al., 2006);

better representation of denitrification in models (Grooß

et al., 2005b; Chipperfield et al., 2005); recognition that

formation of ClOOCl likely occurs faster than previously assumed (Boakes et al., 2005); and better constraints on the equilibrium constant between ClO and

ClOOCl from both laboratory studies (Plenge et al.,

2005) and field observations (von Hobe et al., 2005;

Stimpfle et al., 2004). As detailed below, these findings lead to better model representation of two key

processes: (1) the apparent discrepancies between measured and modeled chemical loss rates of Arctic ozone

during cold Januaries of specific winters (e.g., Rex et

al., 2003) that was noted in Section 3.3 of the previous

Assessment (Frieler et al., 2006), and (2) the relation

between chemical ozone loss and the volume of air

exposed to PSCs, based on more than a decade of observations, that was introduced in the previous Assessment

(Rex et al., 2004; Chipperfield et al., 2005). These

advances suggest better predictive capability for future

ozone loss due to anthropogenic halogens is achievable

in general circulation models (GCMs) and coupled

Chemistry-Climate Models (CCMs), provided certain

fundamental characteristics of the atmosphere (e.g.,

chlorine and bromine loading; photolysis and thermal

decomposition of ClOOCl; denitrification) are well represented. However, as discussed below, certain aspects

of polar chemistry are still subject to considerable

uncertainty, due to differences in various laboratory

measurements of the absorption cross section of

ClOOCl (σClOOCl) (Section 4.2.1.1) and uncertainty in

the atmospheric abundance of BrO (Chapter 2).

4.2.1.1 CHLORINE

Attention since the previous Assessment has continued to be focused on reducing uncertainties in the

parameters that control ozone loss by Cycle 1, such as the

rate constant for formation of ClOOCl (k1a), the photolysis rate of ClOOCl (J1b), and the equilibrium constant

between the concentrations of ClO and ClOOCl (denoted

[ClO] and [ClOOCl], respectively) established if loss of

ClOOCl occurs via thermal decomposition:

KEQ = k1a/k1d = [ClOOCl] / [ClO]2 (3)

This section begins with a description of recent laboratory measurements of reactions that affect polar ozone

loss by chlorine. Then, we describe analyses of field studies

that focus on consistency between measured and modeled

representation of ClO and ClOOCl photochemistry.

Laboratory Studies

Boakes et al. (2005) reported a laboratory measurement of k1a that is ~20% larger, over the temperature

range 206 to 298 K, than the measurements of Bloss et al.

(2001) and the Jet Propulsion Laboratory (JPL)

Publication 02-25 compendium (Sander et al., 2003; hereafter referred to as JPL 02-25) recommendation for this

rate constant. This new study suggests that ozone loss by

Cycle 1 is faster than previously thought, which has implications for the partitioning between ClO and ClOOCl (see

below). Boakes et al. (2005) support the upward revision

in the recommended value of k1a by JPL 02-25 that followed the study of Bloss et al. (2001). The Bloss et al.

(2001) and Boakes et al. (2005) values for k1a are fast

enough to pose a challenge to the theoretical understanding

of this reaction (Golden, 2003).

Plenge et al. (2004) is the only published laboratory

study of ClOOCl photolysis since the previous Assessment.

They found complete production of chlorine atoms and

chloroperoxy radicals (ClOO) upon photolysis of ClOOCl

at 250 and 308 nm, leading to 2Cl + O2 upon the rapid

thermal decomposition of ClOO at polar temperatures,

resulting in catalytic loss of ozone by Cycle 1. This contrasts with an earlier study, described in the previous

Assessment, that indicated a ~10% yield of 2 ClO at 308

nm (Moore et al., 1999) (this branch leads to a null cycle

for ozone). Most ozone depletion models assume production of only ClOO + Cl upon photolysis of ClOOCl. Thus,

Plenge et al. (2004) increase our confidence in the accuracy

of this assumption. However, the primary contribution to

photolysis of ClOOCl occurs longward of 308 nm, and there

remains a need to define the product yields of ClOOCl photolysis for the atmospherically relevant spectral region.

Many of the ozone loss simulations described

below use a value for σClOOCl based on measurements by

Burkholder et al. (1990) out to 410 nm, with a log-linear

extrapolation to 450 nm using a formula given by Stimpfle

et al. (2004). This Burkholder et al. (1990) photolysis rate

of ClOOCl is 40 to 50% faster than J1b based on the JPL

02-25 cross section and results in a good overall description of measured ClO and ClOOCl, if the JPL 02-25 value

of k1a is also used (Stimpfle et al., 2004; see below). The

JPL Publication 06-2 compendium (Sander et al., 2006;

POLAR OZONE

4.21
hereafter referred to as JPL 06-2) recommendation for

σClOOCl is the same as in the previous recommendation.

The correctness of a log-linear extrapolation relies

on the assumption of a single dissociative state in the

wavelength region of interest. The primary photolytic

pathway for ClOOCl is excitation to a singlet electronic

excited state. Peterson and Francisco (2004) suggested

the existence of a weakly absorbing, dissociative triplet

state, predicted to have an absorption maximum at about

385 ± 25 nm, based on an ab initio calculation of the structure of the ClOOCl molecule. This study calls into question the appropriateness of log-linear extrapolations of

ClOOCl cross-section data.

Present uncertainty in J1b underscores the need for

laboratory studies able to quantify σClOOCl to wavelengths

as high as ~450 nm. Existing laboratory studies are based

on spectra measured in the presence of other molecules,

such as molecular chlorine (Cl2), ozone, and/or dichlorine

monoxide (Cl2O). The ClOOCl cross sections must be

determined by analysis of a composite spectrum that

includes absorptions from multiple species. Large uncertainties in laboratory determinations of σClOOCl persist at

atmospherically important wavelengths because the procedures to correct for interfering species can be qualitative in

nature and prone to error (e.g., Huder and DeMore, 1995).

Plenge et al. (2005) recently reported values of KEQ

based on the bond strength of ClOOCl determined using

photoionization mass spectrometry. Their values of KEQ

are smaller than the JPL 02-25 recommendation and the

earlier laboratory study of Nickolaisen et al. (1994), but

agree well with the laboratory measurement of Cox and

Hayman (1988). The range of uncertainty for the JPL 02-

25 value of KEQ encompasses results from all of these laboratory studies. A lower value of KEQ implies that loss of

ClOOCl by thermal decomposition occurs faster than previously thought, leading to higher levels of ClO during

nighttime. Bröske and Zabel (2006) report a value for KEQ,

determined from a laboratory kinetics study over the temperature range 243 to 261 K, that lies between the values

of Cox and Hayman (1988) and Plenge et al. (2005) but

also exhibits a possible pressure dependence. The highpressure (p > 30 hPa) value for KEQ of Bröske and Zabel

(2006) is similar to the JPL 06-2 recommendation. The

low-pressure result, considered to be more reliable by

Bröske and Zabel (2006), leads to a value of KEQ quite

similar to the Plenge et al. (2005) value.

Ultimately, consistency between laboratory thermodynamic determinations of the heat of formation of ClOOCl

and kinetic measurements of k1a and the reverse of k1a is

desired. Determinations of KEQ from kinetics studies are

affected by uncertainties in k1a. Bröske and Zabel (2006)

note that the apparent pressure dependence of KEQ points to

inconsistencies between dissociation and recombination

data, since KEQ is independent of pressure on theoretical

grounds. Consistency between thermodynamic and kinetic

measurements has yet to be achieved, especially for temperatures relevant for the winter polar stratosphere.

Field Observations

Stimpfle et al. (2004) reported the first measurements of ClOOCl, acquired from an instrument aboard

the NASA ER-2 aircraft in the Arctic stratosphere during

the winter of 1999/2000. These observations, analyzed

with simultaneous measurements of ClO from the same

instrument, were used to test the understanding of k1a,

J1b, and KEQ. Assuming the JPL 02-25 recommendation

for k1a, measurements made during daylight, over a wide

range of solar zenith angles, indicate best agreement with

a value for J1b based on absorption cross sections from

Burkholder et al. (1990) (Figure 4-15). The JPL

Publication 00-3 (Sander et al., 2000; hereafter referred

to as JPL 00-3) rate for k1a is most consistent with a value

of J1b based on ClOOCl cross sections that lie between

those given by JPL 02-25 and Burkholder et al. (1990),

while the Boakes et al. (2005) rate for k1a implies faster

photolysis of ClOOCl than given by any of the available

cross section measurements. Figure 4-15 highlights the

fact, emphasized by Stimpfle et al. (2004), that daytime

measurements of ClO and ClOOCl constrain only the

ratio k1a/J1b. The value of J1b decreases by more than an

order of magnitude as solar zenith angle (SZA) increases

from 70° to 92° (Figure 4-15), so analysis of [ClO]2/ -

[ClOOCl] provides a test of the partitioning of these

species for a wide range of photolytic conditions

(Stimpfle et al., 2004). This ratio is important because

models adopting J1b based on Burkholder as well as k1a

from JPL 02-25 will calculate more rapid ozone loss than

models using JPL 02-25 kinetics, since photolysis of

ClOOCl is the rate-limiting step for ozone loss by Cycle

1. Considering the ±25% (1σ) uncertainty in the observations used to define β (see caption of Figure 4-15), a

model based solely on JPL 02-25 kinetics is not consistent with the observations at the 1σ uncertainty level

(Stimpfle et al., 2004).

Field observations of nighttime ClO and in some

cases ClOOCl abundances by several groups suggest that

the equilibrium constant KEQ may be smaller than the

JPL 02-25 recommendation. Observations of ClO and

ClOOCl by Stimpfle et al. (2004), obtained during nighttime when thermal equilibrium (e.g., equation (3)) can

be assumed, indicate better overall agreement with values

of KEQ from the laboratory study of Cox and Hayman

POLAR OZONE

4.22
(1988) compared with values from either JPL 02-25 or

Nickolaisen et al. (1994). Nighttime observations of ClO

and ClOOCl abundances in the Arctic stratosphere during

the winter of 2002/2003 reported by von Hobe et al.

(2005) were used to determine an empirical value for KEQ

that implies faster thermal decomposition for ClOOCl

than the JPL 02-25 value, and somewhat faster decomposition than the Cox and Hayman (1988) and Plenge et

al. (2005) values. Berthet et al. (2005) analyzed nighttime observations of ClO in the Arctic vortex during the

winter of 2002/2003, obtained from a microwave

radiometer onboard the Odin satellite. They concluded

that KEQ lies between the laboratory determination of

Cox and Hayman (1988) and the empirical value of von

Hobe et al. (2005) and that their data are not consistent

with the JPL 02-25 value for KEQ. Finally, an analysis of

daytime and nighttime field measurements together with

thermodynamic calculations and unimolecular rate

theory reveals that overall best consistency between

theory and observation is obtained using the Plenge et al.

(2005) value for KEQ, the Nickolaisen et al. (1994) value

for k1a, and σClOOCl that lies between JPL 02-25 and the

value given by Burkholder et al. (1990) (von Hobe et al.,

2006). The uncertainty in the JPL 02-25 recommendation for KEQ encompasses all of the field observations as

well as the recent laboratory determination based on

measurement of the ClOOCl bond dissociation energy,

so none of the recent results are outside of the range of

prior expectation.

All of these recent results are in basic agreement

with earlier analyses of nighttime atmospheric measurements of ClO (Avallone and Toohey, 2001, and references therein). Vogel et al. (2006) examined chemistry

related to ClO-radical complexes, and concluded that

this chemistry, while still highly uncertain, might be

important for regulating the ratio of ClO and ClOOCl at

night, which might complicate atmospheric, empirical

determinations of KEQ. However, the effect of uncertainty in KEQ on ozone loss is small compared with

other uncertainties, such as the abundance of BrO and

POLAR OZONE

4.23

SZA (deg.)

Figure 4-15. Daytime analysis of all SAGE III

Ozone Loss and Validation Experiment (SOLVE)

flights, β versus solar zenith angle (SZA), where β =

([ClOOCl] / [ClO]2)OBSERVED / ([ClOOCl] / [ClO]2)MODEL.

Data selected for SZA < 92° and M < 3.0 x 1018 molec

cm-3. Average values of β are shown for 2°-wide SZA

bins for 70° < SZA < 86°, and for 1°-wide bins for SZA

> 86°. The top panel shows calculated values of J1b

versus SZA at 60 hPa for a model albedo of 0.24,

using values of the ClOOCl cross section from

Burkholder et al. (1990), JPL 02-25, and Huder and

DeMore (1995), as indicated. Observed albedos are

used for the calculations of β. The bottom three

panels show results for using three values of the

absorption cross section of ClOOCl: Huder and

DeMore (1995), JPL 02-25, and Burkholder et al.

(1990). Within each panel, results are shown for

model simulations using five values of the rate constant of ClO + ClO + M: Trolier et al. (1990), JPL

00-3, Bloss et al. (2001), JPL 02-25, and Boakes et

al. (2005). Error bars depict the ± 25% uncertainty

(1σ) in β attributable to the uncertainties in the observations of ClO and ClOOCl; error bars are shown only

for simulations using the Trolier et al. (1990) and

Boakes et al. (2005) rate constants. Points associated with these two model runs have been displaced

slightly with respect to the actual mean SZA, for clarity

of the error bars. After Stimpfle et al. (2004).
the photolysis rate of ClOOCl. Perhaps the most important reason to reduce uncertainties in KEQ is that accurate knowledge of this parameter will enable better

determination of ClOx (ClO + 2 × ClOOCl) from nighttime observations of ClO.

4.2.1.2 BROMINE

Measurements of BrO profiles in the Arctic vortex

have been obtained for a number of winters using a balloonborne spectrometer (e.g., Pfeilsticker et al., 2000;

Fitzenberger, 2000; Pundt et al., 2002; Dorf et al., 2006).

Model studies that deduce BrOx (BrO + BrCl) or Bry from

these BrO profiles estimate a total bromine loading of 20

to 24 parts per trillion by volume (pptv) (Canty et al.,

2005; Dorf et al., 2006; Frieler et al., 2006), significantly

larger than the abundance of bromine that can be delivered to the stratosphere by methyl bromide (CH3Br) +

halons alone (see Chapter 2). Most chemical transport

models (CTMs), GCMs, and CCMs used to assess polar

ozone loss use estimates of Bry based only on supply of

bromine from CH3Br + halons. A value of bromine radicals (BrOx) inferred from BrO in the Arctic vortex on 18

February 2000 is 40 to 60% higher (e.g., 6 to 9 pptv

larger) than values of BrOx (BrO + BrCl) used in typical

simulations of the Modèle Isentropique de transport

Mésoéchelle de l’Ozone Stratosphérique par Advection

avec CHIMie (MIMOSA-CHIM) 3-D CTM and the

Atmospheric and Environmental Research, Inc. (AER)

two-dimensional (2-D) model (Figure 4-16), resulting in

a larger relative contribution from bromine to polar ozone

loss (Frieler et al., 2006). The BrOx profile shown in

Figure 4-16 provides additional and independent support

to the results to the analyses shown in Chapter 2 because

the chemistry used to infer BrOx from BrO at high latitude (involving only bromine chloride (BrCl) and BrO) is

distinct from the chemistry used to infer Bry from BrO at

midlatitudes (involving mainly BrO and bromine nitrate

(BrONO2), and hypobromous acid (HOBr) and hydrogen

bromide (HBr) to a lesser degree).

Retrievals of BrO profiles from Scanning Imaging

Absorption Spectrometer for Atmospheric Chartography

(SCIAMACHY) limb scattered radiances, which provide

near-global coverage and hence include the polar vortices,

have been conducted by two research groups. Sinnhuber

et al. (2005b) report a modest contribution, 3 ± 3 pptv, to

Bry from VSL compounds, whereas Sioris et al. (2006)

deduce a much larger contribution, 8.4 ± 2 pptv. As discussed in Chapter 2, these discordant findings are driven

by differences in the retrievals of BrO. Schofield et al.

(2006) report ground-based observations of BrO over

Arrival Heights, Antarctica (77.8°S), that are consistent

with Bry of ~21.2 pptv at 20 km, suggesting again a significant role for the presence of more bromine in the polar

regions than can be supplied solely by CH3Br and halons.

Giovanelli et al. (2005) obtained aircraft measurements of

BrO over Antarctica that range from 4 to 10.3 pptv.

Chapter 2 provides an extensive discussion on the quantification of the role of VSL compounds on the stratospheric bromine budget.

The reaction of BrO + ClO has three product channels:

BrO + ClO → ClOO + Br (34%) (4a)

→ BrCl + O2 (7%) (4b)

→ OClO + Br (59%) (4c)

The percentage yields for the three branches at 195 K using

JPL 02-25 kinetics are noted. Channels 4a and 4b lead to

POLAR OZONE

4.24

Figure 4-16. Profile of BrOx (= BrO + BrCl) inferred

from a DOAS measurement of BrO over Kiruna,

Sweden (68°N), on 18 February 2000 (Canty et al.,

2005; Fitzenberger, 2000) compared with a profile

of BrOx found for this region by the AER twodimensional model assuming supply of stratospheric bromine from only CH3Br and halons

(Salawitch et al., 2005). Also shown is an estimate

of Bry from the MIMOSA-CHIM three-dimensional

chemical transport model, assuming again stratospheric supply of Bry only by CH3Br + halons.

Adapted from Frieler et al. (2006).

350

400

450

500

550

0 4 8 12 16 20 24 28

BrO [pptv]

Theta [K]

AER 2-D Model

Mimosa-CHIM

3-D CTM

DOAS

February 2000

68 No

x
ozone loss (following thermal decomposition of ClOO or

photolysis of BrCl), which is generically represented by

process 2a in the depiction of Cycle 2. Channel 4c, the

most efficient branch, results in a null cycle, since the OO bond is not reformed upon loss of OClO.

The only direct determination of the yield of BrCl

from BrO + ClO, a laboratory study conducted at room

temperature (Poulet et al., 1990), resulted in a yield of

~12%, larger than the JPL 02-25 room temperature recommendation of 9 ± 2%. Considerable uncertainty exists

in the branching ratios at temperatures characteristic of

the polar vortices, owing mainly to the lack of direct

laboratory determinations at low temperature. Nighttime abundances of OClO are particularly sensitive to

the branching ratios of BrO + ClO, since the twilight formation of BrCl as a nighttime reservoir of bromine affects

the efficiency of OClO production. An analysis of nighttime OClO measurements, obtained by lunar and stellar

occultation, suggests a higher yield for BrCl and a smaller

yield for OClO from the BrO + ClO reaction at low temperature than is given by JPL 02-25 (Canty et al., 2005).

This result could imply a ~10% increase in the rate of

ozone loss by the BrO + ClO cycle (Canty et al., 2005).

While numerous uncertainties complicate the interpretation of nighttime OClO, the study of Canty et al. (2005)

underscores the need for better definition of the branching

ratios of the BrO + ClO reaction.

The precise level of BrO and Bry in the polar vortices remains an area of active research. Many measurements are emerging from satellite, balloon, and in

situ aircraft instruments. Quantification of the contribution of VSL bromocarbons to the Bry budget of the

polar vortices and of the branching ratios of the BrO +

ClO reaction at low temperature are needed to improve

future quantification of the polar ozone loss rates and to

better quantify the relative contribution of bromine to

polar ozone loss.

4.2.1.3 OZONE LOSS RATES

This section focuses on studies that examine the consistency of modeled and measured chemical loss rates of

polar ozone. The discussion is organized around two key

concepts: (1) that models, using rate constants from JPL 02-

25 and estimates of bromine based on supply of bromine

from only CH3Br + halons, tend to underestimate measured

chemical ozone loss rates, particularly during cold Arctic

Januaries (e.g., Rex et al., 2003 and references therein); and

(2) that the overall amount of chemical loss of column ozone,

for specific Arctic winters, exhibits a tight correlation with

the volume of air in the vortex that has been exposed to PSCs

throughout winter (Rex et al., 2004; Tilmes et al., 2004).

Ozone Loss Rates: Measurement Overview

The estimates of ozone loss rates in Rex et al. (2003)

are based on the Match technique applied to a series of ozonesonde measurements. This study showed that measured

chemical loss rates of Arctic ozone exceed the maximum

possible calculated rates, assuming complete chlorine activation and JPL 00-3 kinetics, for four cold Januaries (1992,

1995, 1996, and 2000). A number of other studies, some of

which have been discussed in previous assessments, also

have documented similar discrepancies between measured

and calculated chemical ozone loss rates during cold Arctic

Januaries (e.g., Hansen et al., 1997; Becker et al., 1998). The

estimates of ozone loss rates in Rex et al. (2004), which

examine processes over longer time scales than Rex et al.

(2003), are based on the “vortex-averaged descent” technique applied to temporal averages of ozonesonde profiles

inside the polar vortex. Rex et al. (2004) showed a tight,

near-linear relation between chemical loss of column ozone

(∆O3) and the volume of air exposed to PSC conditions

(VPSC) in the Arctic vortex for data collected during 10 Arctic

winters that span 12 years (Figure 4-13, previous section).

A preliminary version of this relation was presented in the

previous Assessment. The relation between ∆O3 and VPSC

was poorly simulated by the 3-D CTM simulations of

SLIMCAT shown in Rex et al. (2004). These SLIMCAT

simulations exhibit a slope that is considerably less steep

than the observations (Figure 4-13), implying the model

might be underestimating the sensitivity of future polar

ozone depletion to climate change. However, as shown

below, updated runs of SLIMCAT result in an improved

simulation of the relation between ∆O3 and VPSC.

Tilmes et al. (2004) conducted an independent

analysis of the relation between ∆O3 and VPSC that was

based on chemical loss rates deduced from analyses of

Halogen Occultation Experiment (HALOE) data. Their

study supports the linear relationship between ∆O3 and

VPSC reported by Rex et al. (2004). They also showed that,

if VPSC is averaged over the same time period as ∆O3, then

solar illumination of the cold vortex is a factor that impacts

ozone loss rates.

Ozone Loss Rates: Theory and Observation

A number of studies using various models have

shown considerable improvement in our understanding of

polar ozone loss rates since the previous Assessment. Feng

et al. (2005b) reported 3-D CTM calculations using the

SLIMCAT model of Arctic ozone for the winters of

1999/2000, 2002/2003, and 2003/2004. In their model,

they assumed an extra 100 pptv of chlorine and an extra 6

pptv of bromine reaches the stratosphere, due to shortPOLAR OZONE

4.25
lived halocarbons (these values are consistent with estimates given in Chapter 2). They used JPL 02-25 kinetics,

except the value for σClOOCl was based on Burkholder et

al. (1990), extrapolated to 450 nm as described by Stimpfle

et al. (2004). They used a simple nitric acid trihydrate

(NAT)-based denitrification scheme, described by Davies

et al. (2002). Overall, the model calculations provided a

realistic representation of chemical ozone loss for the three

winters. Some discrepancies between observed and calculated ozone are noted during times of vortex disturbance,

but overall these simulations provide a quantitative

advance in the ability to accurately simulate chemical

ozone loss rates using the SLIMCAT model. Grooß et al.

(2005b) reported improvements in the agreement between

modeled and measured ozone loss rates upon use of a more

realistic denitrification scheme within the CLaMS

Lagrangian trajectory model for the Arctic winter of

2002/2003. A study using the MIMOSA-CHIM CTM

found that denitrification contributed about 23% and 17%

more ozone depletion for the Arctic winters of 1999-2000

and 2002-2003, respectively, compared with model calculations that did not include denitrification (Tripathi et al.,

2006). They also found best agreement between measured and modeled chemical loss of column ozone for

Arctic winters of 1999/2000 and 2002/2003 for a simulation that used the Burkholder et al. (1990) cross section,

allowed for supply of bromine from CH3Br and halons as

well as short-lived species methylene bromide (CH2Br2)

and bromochloromethane (CH2BrCl), and an adjustment

to cooling rates that improved agreement with measured

N2O (i.e., that led to more realistic vortex descent).

Chipperfield et al. (2005) examined the relation

between ∆O3 and VPSC using runs of the SLIMCAT CTM

updated relative to results shown in Rex et al. (2004).

These runs used the Burkholder et al. (1990) extrapolated

value for σClOOCl, an extra 100 pptv of inorganic chlorine

(Cly) and an extra 6 pptv of Bry due to VSL halocarbons

(Chapter 2), and a treatment of denitrification by large

nitric acid trihydrate (NAT) particles (Davies et al., 2002)

that results in denitrification for the cold winters that was

largely absent from earlier ice-based denitrification

schemes. The new denitrification scheme increased modeled ozone depletion by ~30% for cold winters, such as

1999/2000 (Davies et al., 2002; Chipperfield et al., 2005).

Also, the new SLIMCAT runs were at higher spatial resolution (2.8° × 2.8°) than the older simulations (7.5° × 7.5°)

shown in Rex et al. (2004) and used a new radiation

scheme; these improvements lead to stronger descent and

a more isolated vortex. The new SLIMCAT runs simulate

quite well the empirical relation between ∆O3 and VPSC

(Figure 4-13). Improvements relative to the SLIMCAT

results shown in Rex et al. (2004) are due to all three chemical factors as well as the improved model representation

of vortex descent and isolation.

Douglass et al. (2006) simulated ∆O3 versus VPSC

using a CTM driven by winds from a 50-year integration of

the Goddard Earth Observing System (GEOS-4) GCM.

Their calculations were conducted at a resolution of 2.5°

(longitude) by 2.0° (latitude). They concluded that use of

winds from a GCM tends to result in a more realistic simulation of vortex isolation than found using assimilated winds.

Douglass et al. (2006) were able to simulate well the empirical relation between ∆O3 and VPSC for standard photochemistry (e.g., JPL 02-25) and for Bry assuming no contribution

from VSL halocarbons. They also reported that use of the

Burkholder et al. (1990) value for σClOOCl and extra bromine

from VSL compounds resulted in excess ozone loss compared with observations. Hence, it is presently unclear

whether changes in ClOx photochemistry and representation

of Bry from VSL compounds are needed to achieve accurate

representation of the ∆O3 versus VPSC relation within CTMs.

Finally, Tripathi et al. (2006) examined the sensitivity of

ozone loss to CTM resolution, and concluded that their

model required a resolution of 1.0° × 1.0° to properly represent ozone loss at the edge of the vortex at low solar illumination (i.e., during January in the Arctic).

Hoppel et al. (2005a) compared measured ozone

loss rates in the Antarctic vortex, found using the Match

approach applied to POAM III satellite data, with

Lagrangian box model estimates of calculated ozone loss

rates. They reported that agreement between modeled and

measured ozone loss rates is improved if the model

employs larger values of σClOOCl (e.g., Burkholder et al.,

1990) and a total bromine loading of 24 pptv, reflecting an

~8 pptv contribution from VSL bromocarbons. This

amount of bromine from VSL species is near the upper

limit of estimates based on satellite and aircraft data

(Section 2.5, Chapter 2). Tripathi et al. (2006) concluded

that a CTM simulation using JPL 02-25 kinetics and

bromine from CH3Br, halons, CH2Br2, and CH2BrCl

tended to underestimate observed Antarctic ozone loss

rates for several winters. Slightly better agreement was

obtained using Burkholder et al. (1990) cross sections.

Frieler et al. (2006) examined measured and modeled ozone loss rates for the Arctic and Antarctic vortices,

with the measured loss rates based on Match estimates

applied to ozonesonde data, and the modeled loss rates

based on Lagrangian box model calculations. Their reference simulation used JPL 02-25 kinetics and a value for

BrOx based on supply of Bry from only CH3Br + halons.

Using the maximum amount of ClOx available, the reference simulation underestimates observed ozone loss rates

for several Arctic winters, particularly during cold Januaries, such as in 2000 (Figure 4-17, top panel), consistent

POLAR OZONE

4.26
POLAR OZONE

4.27

0

5

0 30 60 90

OZONE LOSS PER

SUNLIT HOUR [ppbv/h]

500 K

0

5 Arctic 2000

NECESSARY

ClO [ppbv]

0

5

0 30 60 90

OZONE LOSS PER

SUNLIT HOUR [ppbv/h]

450 K

DAY OF THE YEAR

0

5

ClO [ppbv]

0

20

40

60

80

100

RELATIVE CONTRIBUTION TO

TOTAL OZONE LOSS [%]

Arctic 1995

Arctic 1996

Arctic 2000

Arctic 2000

Arctic 2001

Arctic 2003

Antarctic 2003

490 K

475 K

500 K

450 K

475 K

475 K

500 K

Figure 4-17. Top two panels: All results are for the Arctic winter 1999/2000. Gray diamonds represent the

mean value of ClOx measured at 450 ± 10 K potential temperature for 8 ER-2 flights that remained entirely

inside the vortex; vertical bars represent the range between maximum and minimum ClOx (Stimpfle et al.,

2004). The lines represent the amount of ClOx necessary to account for the measured ozone loss rate and the

modeled ozone loss assuming ClOx = 3.7 ppbv (“maximum possible ozone loss”), for the following four model

runs:

• Run 1 (black): JPL 02-25 kinetics and modeled BrOx from CH3Br and halons

• Run 2 (blue dashed): Burkholder et al. (1990) σClOOCl and modeled BrOx from CH3Br + halons

• Run 3 (blue dotted): JPL 02-25 kinetics and measured BrOx

• Run 4 (blue solid): Burkholder et al. (1990) σClOOCl and measured BrOx

Red boxes show the observed chemical ozone loss rate per sunlit hour from Match (error bars denote 1σ

uncertainty), used to constrain model estimates of necessary ClOx. Bottom panel: Relative importance of

ozone loss due to the BrO + ClO cycle (gray), the HOCl cycle (orange), the ClO + O cycle (red), and the ClO +

ClO cycle (blue), for various Arctic winters (as indicated). Each grouping shows model results, in order, for the

four model runs described above (e.g., run 4 has highest contribution of BrO + ClO to ozone loss). Adapted

from Frieler et al. (2006).
with previous results (e.g., Rex et al., 2003). They show

that the discrepancy between measured and maximum possible ozone loss rates in January (i.e., rates based on an

assumption of complete chlorine activation) is largely

resolved in a second model run that assumed a profile for

BrOx based on differential optical absorption spectroscopy

(DOAS) balloon BrO that is considerably higher than BrOx

found assuming supply of Bry from only CH3Br + halons

(Figure 4-16), and a value for J1b based on the Burkholder

et al. (1990) value for σClOOCl (extrapolated to 450 nm).

The largest effects on ozone removal rates were the faster

value of J1b (~20% effect) and the higher value of BrOx

(~15% effect). The relative contribution of Cycle 2 (BrO

+ ClO) to overall chemical loss for various winters

increases from 17 to 33% for the reference run to 27 to

48% for the model run that agrees best with overall

measured ozone loss rates (Figure 4-17, bottom panel).

Most important, Frieler et al. (2006) were able to

examine ClOx (Stimpfle et al., 2004), BrOx (Fitzenberger, 2000; Canty et al., 2005), and measured and modeled ozone loss rates for the Arctic winter of 1999/2000.

They showed good overall consistency between all of

these quantities for the second model run (Figure 4-17,

middle panel), suggesting a good quantitative explanation of the chemical loss rates of Arctic ozone for this

winter at 450 K, the highest level where sufficient in situ

measurements of ClOx species are available. This

finding relies on the validity of the larger value of J1b

and higher levels of stratospheric bromine. These

aspects of polar chemistry are still subject to considerable uncertainty, due to differences in various laboratory measurements of σClOOCl (Section 4.2.1.1) and

uncertainty in the atmospheric abundance of BrO

(Chapter 2). Also, quantitative consistency between

observation of polar ozone loss rates and measured ClOx

and BrO has yet to be demonstrated at other altitudes

and for other years.

In summary, numerous recent studies have shown

that models are better able to capture the observed

degree of chemical loss of polar ozone for a wide range

of meteorological conditions that occurred during the

past decade. This is an important advancement since the

previous Assessment. The improvements of simulated

ozone depletion are driven in part by the use of global

models that are able to better capture the isolation of the

polar vortices, due to improved horizontal resolution

and/or winds that minimize excessive horizontal mixing

across vortex boundaries. Previously noted discrepancies between measured and modeled ozone loss rates are

largely resolved, within Lagrangian trajectory box

models, assuming a faster rate of ClOOCl photolysis,

higher levels of BrO that may be consistent with a significant source from VSL bromocarbons, and improved

representation of denitrification that appears to be consistent with available observations (see Section 4.2.2).

While the faster rate of ClOOCl photolysis appears to

be consistent with atmospheric observations of ClO and

ClOOCl, it is important to stress that daytime observations of these species constrain only the ratio k1a / J1b,

and that considerable uncertainty exists in present laboratory determinations of k1a and the cross sections used

to calculate J1b. Nighttime observations of ClO provide

an important test of the consistency of our understanding

of polar ozone photochemistry. However, field observations reveal inconsistencies with many laboratory

studies. Also, we have yet to achieve consistency

between thermodynamic and kinetic laboratory studies

of ClO/ClOOCl photochemistry, in part due to the challenges associated with laboratory work at temperatures

near 200 K. Finally, quantification of BrO and Bry levels

in the polar vortices, and the role of VSL bromocarbons,

presents an additional source of uncertainty that is an

active area of research.

4.2.2 PSC Processes

Polar stratospheric clouds (PSCs) influence

ozone loss through two main processes: (1) chlorine

activation on PSC particles leading to ozone losses and

(2) sedimentation of PSCs causing denitrification and

exacerbating ozone loss. Our understanding of denitrification in particular continues to evolve, mainly in

response to research into the characteristics and formation of solid-phase nitric acid hydrate particles. WMO

(2003) described the unexpectedly low (~10-4 cm-3) concentrations of large nitric acid particles that were first

observed during the 1999/2000 Arctic winter with the

introduction of improved instrumentation (Fahey et al.,

2001). Several studies have now concluded that such

particles are widespread, both from in situ observations

during the 2002/2003 Arctic winter (Larsen et al., 2004;

Voigt et al., 2005) and from updated analyses of remote

measurements (Poole et al., 2003; Adriani et al., 2004).

Using multiple simultaneous PSC measurements, Larsen

et al. (2004) have demonstrated that a low concentration

of nitric acid trihydrate (NAT) particles can be masked

when the coexisting liquid particles become optically

dominant. Therefore, a background population of solid

particles may be more pervasive than is apparent in

remote sensing observations, requiring the introduction

of new PSC formation mechanisms.

Synoptic-scale ice formation below the frost

point (Tice), which was considered to be a primary mechanism for solid nitric acid formation, is much too infrePOLAR OZONE

4.28
quent to account for the widespread formation of solid

nitric acid particles in the Arctic. In contrast, mesoscale

temperatures below Tice may provide a NAT formation

mechanism, through nucleation of nitric acid hydrates

on mountain wave-induced ice clouds. High concentrations of mesoscale solid nitric acid particles have been

unambiguously identified in multiple studies (Carslaw

et al., 1999; Wirth et al., 1999; Hu et al., 2002;

Fueglistaler et al., 2003; Luo et al., 2003). Modeling

studies have demonstrated that sedimentation from the

base of a NAT cloud with high particle concentrations

(the “mother cloud”) allows selective growth of a few

NAT particles in the underlying supersaturated, cloudfree air (Fueglistaler et al., 2002a; Dhaniyala et al.,

2002). Over several days, low concentrations (10-2-10-5

cm-3) of large NAT particles are generated over a wide

vertical and horizontal extent. This mechanism can successfully explain observations of low NAT particle concentrations from clouds assumed to initially contain

much higher solid particle concentrations (Fueglistaler

et al., 2002b). Mann et al. (2005) have shown that

including this mechanism in a 3-D model can produce

large NAT particles in up to 60% of the NAT supersaturated region in the Arctic. Most studies have focused on

the Arctic, but some Antarctic NAT PSC observations

have also been attributed to mesoscale nucleation

(Höpfner et al., 2006).

Other Arctic PSC observations, however, cannot

be explained by mesoscale cloud formation (Pagan et

al., 2004; Larsen et al., 2004; Voigt et al., 2005). Pagan

et al. (2004) examined the role of mountain wave ice

clouds in generating solid-phase nitric acid PSCs that

were observed on three dates. Using satellite observations and model results to identify any regions with ice

cloud formation, Pagan et al. (2004) concluded that none

of the observed PSCs could have originated in an ice

cloud. In situ measurements of large nitric acid particles at low concentrations made by Voigt et al. (2005) in

the Arctic also cannot be attributed to ice formation, at

either the synoptic scale or mesoscale. The observed

particles formed within one day of the temperature

dropping below the NAT temperature, TNAT, with temperatures at most 3.1 K below TNAT (implying NAT

supersaturation ratios less than ten). The narrow range

of environmental conditions strictly constrains the formation mechanism and the NAT freezing rate. Voigt et

al. (2005) show that heterogeneous freezing of NAT,

specifically triggered by meteoritic particles, could

explain the observations. Homogeneous freezing of

NAT is unlikely, because laboratory data show that such

freezing is very slow under the observed conditions

(Knopf et al., 2002). Alternative interpretations of the

homogeneous freezing rates have been proposed

(Tabazadeh, 2003). However, in a study of several proposed homogeneous freezing rates, Drdla and Browell

(2004) concluded that none yielded model results consistent with observations of denitrification and PSC

onset for the 1999/2000 Arctic winter. Heterogeneous

freezing currently appears to be the most likely mechanism for producing nitric acid hydrate particles at synoptic scales in the Arctic. However, the specifics remain

poorly defined, such as the nuclei involved, the factors

that control the freezing rate, and the extent to which

heterogeneous freezing also occurs in the Antarctic.

Based on these PSC formation mechanisms, new

denitrification schemes for CTMs have been developed.

Previous denitrification schemes that required synoptic

scale temperatures below Tice frequently failed to get

any Arctic denitrification; the new schemes produce

much more widespread denitrification. Revised denitrification has contributed to improvements in modeling

ozone loss (see Section 4.2.1.3). Synoptic-scale particle

formation has been approximated by applying a slow,

uniform NAT particle-formation rate whenever NAT is

supersaturated. Using this approach and an assumed

rate of 8 × 10-10 cm-3s-1, Davies et al. (2005) have simulated several Arctic winters with the SLIMCAT CTM.

Generally good agreement was found between model

denitrification and observations (both satellite and in

situ), but interannual variability in denitrification was

not fully captured. The CLaMS has introduced a similar denitrification scheme (Grooß et al., 2005b); for the

Arctic winter studied by Grooß et al. (2005b), a

prescribed rate of 2 × 10-9 cm-3s-1 best reproduced

observed denitrification. A single, constant particle formation rate appears insufficient to simulate interannual

variability, but refinements will require a more detailed

understanding of PSC formation mechanisms and their

rates.

Alternatively, denitrification may be caused by

mesoscale NAT PSCs. This mechanism has been studied

with the SLIMCAT CTM (Mann et al., 2005) by introducing mountain wave-induced ice clouds that generate

NAT “mother clouds.” They concluded that approximately 80% of the denitrification observed in the

1999/2000 Arctic winter could have been caused by sedimentation of particles out of mountain wave-induced NAT

mother clouds. Further quantification of this mechanism

is limited by the requirement for accurate mountain wave

information spanning the polar vortex. Therefore, the relative importance of synoptic scale and mesoscale

processes for denitrification remains uncertain.

In summary, the mechanisms responsible for NAT

PSC formation and thus denitrification have been refined

POLAR OZONE

4.29
over the last four years, with increasing evidence that key

processes occur above the ice frost point. However,

specifics of PSC formation remain uncertain, forcing

model denitrification to rely upon empirical freezing rates

that are unable to fully capture interannual variability.

Recent findings are primarily based on Arctic observations. Whether these PSC formation processes and denitrification mechanisms are equally applicable in the

Antarctic remains to be evaluated.

4.3 RECENT POLAR OZONE CHANGES

4.3.1 2002 Antarctic Ozone Hole

In September 2002, a SH major stratospheric

warming split the polar vortex and ozone hole for the first

time in the observational record. Major warmings cause

dramatic stratospheric circulation changes. Warmings are

caused by planetary waves propagating up from the troposphere. These waves decelerate the polar night jet,

increase polar temperatures, and distort and/or split the

vortex. Warmings are classified as major if the 10 hPa

zonal-mean temperatures increase poleward of 60° and 10

hPa zonal-mean winds turn easterly (Julian, 1967;

Labitzke, 1968). Major warmings are distinguished from

final warmings, where final warmings mark the transition

from the cold vortex in winter to the warm anticyclone in

summer.

The September 2002 major warming was puzzling

because it was previously thought they only happened in

the NH, where the tropospheric forced planetary wave

activity is much stronger (see WMO, 1986, Section 6.1.7).

The Arctic polar vortex is regularly disturbed by waves,

with major warmings occurring every two to three years.

The difference in planetary wave activity between the

hemispheres is due to various factors: less orographic

forcing and weaker longitudinal land-sea contrast in the

SH, and the presence of the cold elevated Antarctic continent at the pole. As a consequence, Antarctic winter stratospheric temperatures are much colder than the Arctic and

exhibit much less variability (see midwinter period in

Figure 4-1). Furthermore, temperature records from 1940

do not show evidence of any major Antarctic warmings

(Roscoe et al., 2005; Naujokat and Roscoe, 2005).

This unprecedented event induced a dramatic reduction of the ozone hole area to less than 5 million km2 as

compared with more than 20 million km2 in the previous

years (Stolarski et al., 2005). Its occurrence triggered

numerous investigations using meteorological analyses,

observations of chemical species from various satellite and

ground-based instruments, and model simulations.

4.3.1.1 OBSERVATIONS

Chemistry

The major warming had a dramatic impact on total

ozone (Stolarski et al., 2005). On 23 September, the

ozone hole elongated and split in two pieces (Figure 4-

18). One piece drifted over South America and dissipated, while the other drifted back over the pole as a

significantly weakened ozone hole. The 2002 total

ozone daily minimum value did not reach values lower

than 150 DU, as compared with around 100 DU in the

preceding decade. Higher total ozone values were

observed in the polar region from mid-September to

mid-October. Ozone hole metrics (Figures 4-7 and 4-8)

all show remarkable deviations from averages over the

last decade.

Ozone profiles showed that the vortex was not

vertically aligned during the major warming. Solar

Backscatter Ultraviolet (SBUV) measurements show an

increase of ozone in the 70°-80°S region of about 200

POLAR OZONE

4.30

Sept 18, 2002 Sept 23, 2002

Sept 28, 2002 Oct 3, 2002

100160220280340400460520

Total Ozone (DU)

Figure 4-18. TOMS total ozone maps for four

days during September and October 2002. The

white space around the South Pole is polar night,

where no measurements are made. Adapted

from Stolarski et al. (2005).
DU due to the influx of low latitude air (Kondragunta et

al., 2005). Upper levels of the vortex were displaced

from the pole so that the ozone-depleted lower stratosphere region was overlaid by ozone-rich extra-vortex

air (Allen et al., 2003; Randall et al., 2005a; Yela et al.,

2005). POAM measurements showed that below 22 km,

the chemical ozone loss was similar to previous years

up to the major warming (Hoppel et al., 2003).

Measurements of minor chemical species provide

further insights into the 2002 ozone hole. Most measurements showed that prior to the warming, the vortex

was denitrified, with near-complete chlorine activation

and extensive PSC areas up to 24 km (Ricaud et al.,

2005; von Savigny et al., 2005). The large temperature

increase during the warming induced rapid PSC disappearance and deactivation of chlorine radicals. Faster

decay of the OClO slant column was measured by satellite and ground observations in September as compared

with the same period in 1996-2001 (Richter et al., 2005;

Frieß et al., 2005). The re-establishment in October of a

weak vortex in the lowermost polar stratosphere was

confirmed by satellite observations that showed low

ozone mixing ratios below 20 km (Ricaud et al., 2005;

von Savigny et al., 2005). As the middle stratosphere

remained strongly perturbed, rapid NO2 recovery above

25 km was observed, inducing larger than usual total

NO2 columns in the polar regions in October (Frieß et

al., 2005; Richter et al., 2005; Yela et al., 2005).

Dynamics

Meteorological conditions in 2002 showed that

the early winter was already unusually disturbed (Hio

and Yoden, 2005; Newman and Nash, 2005; Allen et al.,

2003). Figure 4-19 (second panel) displays the vertical

distribution of the 2002 zonal-mean temperature departure from the 1979-2001 mean in the 55°-75°S latitude

band. From June, temperatures steadily increased with

respect to the climatological average until the September

major warming. Temperature increases are associated

with decreases of the zonal-mean wind intensity at 60°S

and 10hPa (Figure 4-19, panels 1 and 3). In August, the

zonal wind dropped below the 1979-2001 range of

values and turned easterly during the warming. The temperature increase in the polar vortex collar is directly

controlled by the planetary waves propagating upward

from the troposphere to the stratosphere. The black line

in the bottom panel of Figure 4-19 is the time series of

midlatitude eddy heat flux in 2002 at 100 hPa for waves

1-3 (heat flux is proportional to the vertically propagating wave activity). The 2002 times series is

compared with the climatological average over the 1979-

2001 period (gray shading). There are several significant wave events from May to October. After each wave

event, the stratosphere warmed by a few degrees until

the major warming in late September. During this event,

the eddy heat flux reached twice the largest value in the

1979-2001 period.

The synoptic development of the vortex split in

September 2002 is described in several studies (e.g.,

Charlton et al., 2005; Krüger et al., 2005). The vortex

began to split on 24 September. By 26 September, it had

split completely at 10 hPa into two vortices of similar

intensity with a tongue of anticyclonic circulation

stretching across the pole and a region of high temperature between each vortex and the anticyclone. By 8

October, a single weakened vortex was re-established

over the pole. The vortex split extended up to 1 hPa. In

the lower stratosphere, the vortex did not split but

formed two distinct cyclonic systems. During the major

warming, polar temperatures south of 60° increased by

about 25 K at 10 hPa and local temperature at several

Antarctic stations showed an increase of up to 50 K at

30 and 10 hPa. The 60°S mean zonal wind reversed from

~60 m/s westerly to ~15 m/s easterly at 10 hPa until 30

September, when it switched back to westerlies. Much

weaker westerly winds were re-established in October

as one of the cyclonic remnants of the polar vortex

moved back toward the pole. The final warming of the

vortex occurred in late October, much earlier than in

previous years.

4.3.1.2 MODELING OF THE WARMING

The 2002 warming was simulated by mechanistic

models for reproducing the event’s dynamics, and by

CTMs for a detailed evaluation of the vortex chemical evolution. The UK stratosphere-mesosphere model (USMM)

reproduced the warming’s dynamical features (Manney et

al., 2005a). This simulation was initialized on 14

September 2002 and forced at 100 hPa by analyzed geopotential heights. The model produced a good simulation of

the vortex recovery phase, suggesting that the warming’s

evolution was largely determined by the initial conditions

and the prescribed 100 hPa geopotential heights. The

modeled transport during the warming showed enhanced

diabatic descent in the vortex below ~700 K and strong

poleward transport and mixing in mid- to high-latitude

regions. The strong vortex vertical tilt during the warming

was well reproduced, showing low-latitude air surrounding and overlying the vortices after the split in the

middle stratosphere. Sensitivity tests indicate a strong

dependence on the boundary forcing, especially the amplitude and upward propagation of planetary wave-2.

POLAR OZONE

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POLAR OZONE

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Apr May Jun Jul Aug Sep Oct

Apr May Jun Jul Aug Sep Oct

Apr May Jun Jul Aug Sep Oct

2002 U* (20 ̊S-90 ̊S)

2002 T* (55 ̊S-75 ̊S)

2002 U (10 hPa, 60 ̊S)

2002 v'T' (100 hPa, 40 ̊S-70 ̊S)

Figure 4-19. From top to bottom: (1) Daily zonal-mean zonal wind departures from the 1979-2001 mean for 1

April to 31 October 2002, averaged for 20°S to 90°S. Contour intervals are 2 m s-1 with dashed negative contours. (2) Daily zonal-mean temperature departures, averaged for 55°S to 75°S. Contour intervals are 2 K

with dashed negative contours. (3) Daily zonal-mean zonal wind at 10 hPa and 60°S. Units are in m s-1. (4)

Daily zonal-mean eddy heat flux at 100 hPa, averaged for 40°S to 70°S. Units are in K m s-1. In the two bottom

panels, the black line shows 2002 values, which have been smoothed with a 1–2–1 filter applied three times.

The white line displays the 23-year average (1979-2001), smoothed with a 15-day boxcar. The gray shading

indicates the range of values (also 1–2–1 filtered) observed between 1979 and 2001. Adapted from Newman

and Nash (2005) and Allen et al. (2003).
Several CTMs simulated the chemical polar ozone

loss during the 2002 winter and spring. Using ECMWF

or United Kingdom Meteorological Office (UKMO)

analyses, the simulations successfully reproduced the main

features of the 2002 Antarctic ozone hole (e.g., Ricaud et

al., 2005; Konopka et al., 2005; Grooß et al., 2005a;

Sinnhuber et al., 2003). Feng et al. (2005a) compared simulations of the 2002 Antarctic ozone hole with calculations for 2000. While large ozone losses were observed in

2002, the warmer than average 2002 vortex temperatures

led to smaller amounts of ClOx that induced smaller chemical ozone loss (Richter et al., 2005; Frieß et al., 2005).

The accumulated chemical loss in the polar lower stratosphere was about 20 DU less than in 2000. Significant

contributions to the larger September 2002 ozone column

were caused by enhanced descent at the vortex edge and

increased horizontal transport associated with the distorted

vortex. Grooß et al. (2005a) showed that the rapid chlorine deactivation after the major warming depended on the

ozone mixing ratio. Very low ozone mixing ratios favored

formation of HCl, while larger ozone mixing ratios and

less denitrification led to formation of ClONO2. At higher

altitudes (above 700 K), the midlatitude air masses transported to the polar region during the major warming experienced very large ozone depletion rates caused by NOx

catalytic cycles typical of polar summer chemistry.

Several model simulations showed enhanced ozone loss

in the midlatitudes due to the dilution of ozone-poor vortex

air during the major warming (Grooß et al., 2005a; Feng

et al., 2005a; Marchand et al., 2005). The polar vortex

remnant that survived the major warming was strongly

isolated from the extra vortex air until late November and

did not experience any significant dilution with the midlatitude air (Konopka et al., 2005).

4.3.1.3 THEORETICAL UNDERSTANDING

Several studies examined the conditions that led to

the major warming in 2002 in order to explain this

unprecedented event. Scaife et al. (2005) and Newman

and Nash (2005) both argued that the stratosphere had been

preconditioned throughout the preceding months by stratospheric vacillations in the zonal-mean winds. Scaife et

al. (2005) further argued that the large September tropospheric wave pulse that caused the warming was also

dependent on this preconditioning. Numerical simulations

of the stratospheric flow show distinct stratospheric

regimes that are either steady or vacillating. Vacillations

are found for high levels of planetary wave forcing emanating from the troposphere (Scaife and James, 2000). The

SH stratosphere flow is generally in quasi-steady state or

in vacillating regimes for a short period of time. The 2002

winter was in a vacillating regime beginning in June. This

vacillation induced a systematic weakening of the polarnight jet that ultimately allowed a strong pulse of planetary wave to propagate into the stratosphere.

Gray et al. (2005) examined the influence of the

QBO on the major warming. Contrary to expectations,

the warming event occurred during the QBO west phase

in the lower stratosphere. NH major warmings are likelier

in the QBO easterly phase when a zero wind line confines

planetary wave propagation to higher latitudes, closer to

the polar vortex. However, meteorological analyses indicated the presence of anomalously strong easterly equatorial winds above 10 hPa from January to September 2002.

The influence of these anomalous winds was investigated

in idealized model experiments where equatorial winds

were relaxed to observations obtained in several years. It

was shown that the 2002 equatorial winds hurried the simulated warming event. Harnik et al. (2005) also suggested

that the low latitude, middle stratosphere easterly wind

was a main factor in the warming. They showed that the

zonal wind at 30°S and 10 hPa turned easterly in May

2002, due to a large burst of upward wave activity into the

stratosphere. The primary effect of this event was the

deceleration of the low-latitude winds in the upper stratosphere. The zero wind line was shifted significantly farther poleward than normal and the resulting enhanced

poleward wave focusing is likely to have contributed to

the erosion of the vortex.

Moreover the vortex preconditioning was connected with the persistence of large wave-2 amplitudes

through the winter. The warming itself occurred when a

large wave-1 disturbance combined with a traveling wave-2.

Such a simultaneous combination was not observed during

the period 1979-2001. Through a detailed study of the

wave activity in both the troposphere and stratosphere in

the winter 2002, Newman and Nash (2005) confirmed

that the major warming could be explained by two main

factors: (1) stronger than usual tropospheric wave forcing

that propagated upward into the stratosphere and (2) a

propagation state in the stratosphere that favored upward

propagation of waves. They showed that for AprilSeptember, the heat flux at 200 hPa for wave 1-3 between

40°S and 70°S was 50% larger than climatology. The

stronger wave-1 in the lower stratosphere was found to

be statistically related to wave-1 in the lower troposphere,

itself highly correlated with wave-1 in the tropics. The

2002 winter average wave-1 amplitudes in both the midlatitude lower troposphere and tropical upper troposphere

were the largest observed over the 1979-2002 record.

During the 2002 winter, larger-than-normal

planetary-scale wave events occurred regularly in the

POLAR OZONE

4.33
midlatitude troposphere. The zonal-mean flow at the

tropopause and in the tropical upper stratosphere was

conducive to wave propagation, and these waves in the

troposphere were efficiently refracted upward into the

stratosphere. Each wave event warmed the polar lower

stratosphere and weakened the jet stream. By midSeptember the cumulative effect of wave activity had

efficiently preconditioned the stratosphere. The major

warming was preceded by an extraordinary strong pulse

of eddy heat flux in the upper troposphere/lower stratosphere, much stronger than what is typically observed

in the SH and even in a NH major warming (Harnik et

al., 2005; Manney et al., 2005a).

4.3.2 Is Polar Ozone Getting Worse?

Observed total ozone columns in the polar regions

have been larger in recent winters than in the 1990s, as

emphasized in Section 4.1.2. The stabilization of ozonedepleting substances (ODSs) in the stratosphere has raised

the question of its impact on observed ozone levels in the

polar regions. This topic is discussed throughout the

Assessment; see Chapter 1 for past and present ODS

levels, Section 6.5.3 of Chapter 6 for attribution of recent

changes in polar ozone as they pertain to changes in

halogen loadings, Section 6.6.4 of Chapter 6 for projected

future changes in polar ozone, and Chapter 8 for future

ODS levels. The effect of anthropogenic elevated halogen

levels in the stratosphere on polar ozone is different in

both hemispheres due to differences in meteorological

conditions. The Antarctic winter stratosphere is characterized by cold and stable meteorological conditions. The

large chlorine activation induced by the very low temperatures in the vortex core leads to a nearly complete ozone

destruction in the lower stratosphere by the end of

September, which indicates that the Antarctic ozone depletion is saturated. The effects of ODS leveling off and saturation are difficult to discriminate. In contrast, Arctic

chemical ozone loss is not saturated, due to much warmer

conditions in the winter stratosphere. Arctic year-to-year

variation is primarily controlled by the year-to-year variability in stratospheric temperatures. These temperature

variations are much larger in the NH than in the SH. The

large NH variability makes it difficult to establish longterm trends in Arctic ozone.

4.3.2.1 ARCTIC

Recent winters (since 1997/1998) have been characterized by higher than average temperatures than in the

1990s, although 1999/2000 and 2004/2005 stand out as

being particularly cold, with record ozone chemical depletion (see Section 4.1.2). Major stratospheric warmings

were more frequent prior to 1990, while the 1990s were

remarkable by the absence of such events (except during

the winter 1998/1999). Studies of Arctic chemical ozone

loss over this range of years and meteorological conditions have shown that a near-linear relationship can be

found between the vortex average chemical ozone loss and

VPSC, the stratospheric volume of potential PSC as shown

in Figure 4-13 (Rex et al., 2004, 2006). Arctic ozone losses

since 2000 are in line with the relationship based on data

collected prior to 2000, with very low (as in 2001/2002)

and very large (as in 2004/2005) chemical ozone loss that

is clearly due to differences in VPSC for these years. No

specific change in the behavior of Arctic ozone loss can

be detected from these observations, which is consistent

with present understanding given the similar halogen

loadings over the time period for which this relationship

can be defined from observations.

4.3.2.2 ANTARCTIC

Several studies have analyzed the evolution of the

ozone hole since its appearance in the early 1980s.

Newman et al. (2004) investigated the ozone hole area

(OHA), while Bodeker et al. (2005) investigated OHA,

annual minimum, and ozone mass deficit (OMD) (see

Section 4.1.2 and Figure 4-8). The OHA and minimum

ozone showed modest increases in severity in the 1990s,

while OMD increased greatly in severity in the 1990s. The

trend of monthly averaged OHA was analyzed by AlvarezMadrigal and Pérez-Peraza (2005) in the period 1982-

2003. After continuous growth since the eighties, the trend

in October and November over 6-year subsets became negative since 2002. However, this trend did not account for

temperature effects, or for dynamical effects during the

breakup phase of the ozone hole.

Other studies investigated the relationship between

the evolution of the Antarctic ozone hole and stratospheric

chlorine and bromine levels over the last two decades.

Newman et al. (2004) have shown that OHA is mainly controlled by catalytic ozone loss with stratospheric chlorine

and bromine, and secondarily by the year-to-year variations in temperature in the collar region, near the edge of

the polar vortex. Newman et al. (2004) fit the OHA to

equivalent effective stratospheric chlorine (EESC, see

Chapter 8, Box 8.1) with a 6-year time lag, and showed

that OHA is currently decreasing at a very slow rate of

0.4% per year in response to halogen decreases. Huck et

al. (2005) also fit the annual mean OMD to EESC with a

3-year time lag, the eddy heat flux anomalies at 20 hPa

and 60°S, and South Pole 100 hPa temperature anomalies.

In both studies, EESC and temperature explain almost all

POLAR OZONE

4.34
of the variability of ozone-hole diagnostics. The residuals

of the OHA fit to EESC and temperature have year-to-year

variations of a few millions km2 (compared with a current

area of 23-24 millions km2

). This variance masks the area

response to the small halogen decreases over Antarctica

since 2000.

An increase of ozone due to ODS decreases can also

be looked for in regions where the chemical loss is still

the main factor affecting ozone variations but without saturation, as in the upper edge (above 20 km) or in the collar

region (60°S-70°S) of the Antarctic ozone hole. Using

measurements from ozonesondes and the satellite-borne

POAM instrument, Hoppel et al. (2005b) showed that the

ozone mixing ratio near the top of the ozone hole at 20-22

km in early October was higher in the period 2001-2004

than in the previous years of POAM measurements (1994-

1996; 1998-2000). However, this increase was accompanied by higher temperatures and reduced PSC occurrence

frequency, which indicates that the ozone changes were

linked to temperature variability associated with changes

in dynamical processes rather than a decline in ODSs. The

relationship between ozone mixing ratios and temperature

in recent years also was noted by Solomon et al. (2005)

from the analysis of four decades of ozonesondes over

Antarctica. Yang et al. (2005) studied the evolution of

ozone in the extra-vortex collar region (60°S-70°S), but

this study is not completely representative of ozone hole

conditions.

The latest studies on polar ozone thus show no further increase of the severity of the polar ozone depletion

both in the Arctic and the Antarctic stratosphere. The

Antarctic ozone hole shows a clear leveling off since the

beginning of the century. The leveling off is primarily due

to saturation of losses because of ODSs. Both 2002 and

2004 were weak ozone holes, giving the appearance of a

downward trend. However, these weak holes directly

resulted primarily from active SH dynamical forcing.

Interannual variability of the SH is currently masking the

expected ozone hole improvement as ODSs decrease.

Arctic losses also appear to be decreasing, but the very

large interannual variability of the Arctic masks the

improvement as ODSs decrease.

4.4 THE INFLUENCE OF PRECIPITATING

CHARGED PARTICLES ON POLAR OZONE

Precipitating charged particles influence ozone and

other constituents. Processes on the Sun and in Earth’s

magnetosphere, the interaction of the solar wind with

Earth’s magnetosphere, and the interplanetary magnetic

field ultimately drive the flux of charged particles into the

atmosphere and have led to observed and predicted stratospheric and mesospheric constituent changes in the past

few years. Some of these natural charged-particle-driven

variations have been measured to cause significant

increases in odd nitrogen and odd hydrogen constituents,

with corresponding significant ozone decreases. These

effects are transient as ozone recovers to quiescent levels

within days (middle to upper mesosphere) to several years

(middle and lower stratosphere) after these natural

impacts. Since the previous Ozone Assessment, a wealth

of new satellite measurements have become available that

have helped to better quantify the charged particle atmospheric influences.

The charged particle effects that influence the

atmosphere can roughly be grouped into three types of

perturbations: (1) solar particle events, which are primarily protons entering the polar regions and thereby often

are referred to as solar proton events (SPEs); (2) energetic

electrons precipitating in the auroral zone and lower latitudes; and (3) galactic cosmic rays. Galactic cosmic rays

(GCRs) continually create odd nitrogen and odd hydrogen

constituents in the lower stratosphere and upper troposphere but play a small role in polar ozone variations.

SPEs and energetic precipitating electrons influence polar ozone levels. The most recent solar cycle period

(solar cycle 23) was very active, with both SPEs and

periods of enhanced fluxes of energetic electrons. In particular, six of the nine largest SPEs in the past 40 years

occurred in the most recent solar cycle. The precipitating

particles associated with these solar-driven events produced ionizations, excitations, dissociations, and dissociative ionizations of the background constituents in both the

polar cap (reaching to 60° geomagnetic latitude or lower

during strong geomagnetic storms) and lower latitudes,

including the auroral oval region. The solar protons primarily deposited their energy in the mesosphere and

stratosphere, whereas the energetic electrons primarily

deposited their energy in the thermosphere and upper

mesosphere.

HOx (odd hydrogen, atomic hydrogen (H),

hydroxyl radical (OH), hydroperoxyl radical (HO2)) constituents are created by a series of ion chemistry reactions

(e.g., Solomon et al., 1981) as a result of the charged particle precipitation. The solar particles and associated secondary electrons (produced in ionization events) also

create atomic nitrogen through dissociation-producing

collisions with molecular nitrogen (N2). Atomic nitrogen

then leads to the production of other odd nitrogen constituents, NOy (total reactive nitrogen; usually includes

atomic nitrogen (N), nitric oxide (NO), NO2, nitrogen trioxide (NO3), dinitrogen radical (N2O5), nitric acid

(HNO3), peroxynitric acid (HNO4), BrONO2, ClONO2)

through chemical reactions.

POLAR OZONE

4.35
The HOx increases cause short-lived ozone

decreases in the polar mesosphere and upper stratosphere

due to the short lifetimes of the HOx constituents. This

HOx-caused ozone effect was first observed in Weeks et

al. (1972) and explained in Swider and Keneshea (1973).

The NOy increases lead to both short- and long-lived polar

stratospheric ozone changes because of the long lifetime

of the NOy family in this region. This NOy-caused ozone

impact was first suggested by Crutzen et al. (1975) and

observed in Heath et al. (1977). Since these early works,

a number of papers have been published that document

these solar-caused polar changes (recently reviewed in

Jackman and McPeters, 2004). Although HOx-driven

short-lived decreases are useful for understanding the

mesosphere and upper stratosphere, the NOy enhancements cause the more important charged-particle-induced

polar ozone decreases.

4.4.1 Odd Nitrogen (NOy) Enhancements

Substantial increases in odd nitrogen (NOy) constituents in the mesosphere and upper stratosphere as a

direct result of SPEs have been measured by several satellite instruments during solar cycle 23. Very large fluxes

of solar protons in July 2000 produced huge increases (>50

parts per billion by volume (ppbv) in the mesosphere) in

Arctic NOx (NO + NO2) (Jackman et al., 2001). The large

SPEs in late October and early November 2003 also caused

very large proton fluxes that created substantial amounts

of NOx (Jackman et al., 2005a; López-Puertas et al., 2005a;

Seppälä et al., 2004). Other NOy constituents also were

elevated as a result of huge SPEs that occurred in

October/November 2003, including HNO3(Orsolini et al.,

2005; López-Puertas et al., 2005b) and N2O5 and ClONO2

(López-Puertas et al., 2005b).

Solar protons also caused long-term NOy enhancements. For example, Randall et al. (2001) showed that

Antarctic middle stratospheric NOx in September 2000

was enhanced as a result of the transport of huge NOx

enhancements from the lower mesosphere and upper stratosphere, which were originally produced by the solar protons during the July 2000 SPE. Large enhancements of

polar stratospheric NO2 through November and into early

December of 2003 probably were caused by the solar protons from the October/November 2003 SPEs (Seppälä et

al., 2004; López-Puertas et al., 2005a).

High Arctic NOx was observed in January to July

2004 by several satellite instruments (Natarajan et al.,

2004; Randall et al., 2005b; Rinsland et al., 2005; LópezPuertas et al., 2005a, 2006). Figure 4-20 (top, adapted

from Randall et al., 2005b) shows NO2 at 40 km for years

1994-1996 and 1998-2004. Very significant NO2 enhancePOLAR OZONE

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2004

Figure 4-20. Seven-day running means of POAM II

(1994-1996) and POAM III (1998-2004) NO2 (top)

and ozone (middle) mixing ratios in the NH at 40 km.

The NO2 enhancements and ozone reductions in the

spring of 2004 resulted from transport to the stratosphere of NOx produced by energetic particles precipitating in the mesosphere and thermosphere

during the winter. The bottom panel shows 7-day

running means of POAM partial column ozone over

the vertical region from 35 to 50 km. Black lines represent the average (thick) ± 1σ standard deviation

(thin) for column measurements from 1994 to 1996

and 1998 to 2004 (gray dots), excluding 2004. The

red curve shows results for 2004 in all three panels.

POAM II measurements have been corrected for

instrument biases as determined by comparisons

with independent satellite measurements. Adapted

from Randall et al. (2005b).
ments are observed from March through July in 2004. It

is unclear when these enhanced Northern Hemisphere NOx

levels were generated, although it is likely that energetic

particle precipitation was the cause. The energetic electron fluxes associated with the geomagnetic storms of

October/November 2003 were large and probably generated significant mesospheric and thermospheric NOx

(Natarajan et al., 2004), which could have been transported

to lower levels. Other active geomagnetic periods in late

November and December 2003 and even January 2004

could have been the main sources of the NOx observed in

the stratosphere and mesosphere (Randall et al., 2005b;

López-Puertas et al., 2005a, 2006).

Funke et al. (2006) found high levels of Antarctic

upper stratospheric NOx during May to August 2003. This

NOx enhancement was attributed to precipitating electrons

in the lower thermosphere and subsequent descent during

polar night.

4.4.2 Ozone Decreases

Substantial mesospheric and upper stratospheric

ozone decreases during and shortly after the July 2000,

October/November 2003, and January 2005 SPEs were

measured by several satellite instruments (see Jackman et

al., 2001, 2005a, b; Seppälä et al., 2004, 2006; Verronen

et al., 2005; Degenstein et al., 2005; López-Puertas et al.,

2005a; Rohen et al., 2005). The transported NOy enhancements from SPEs also resulted in observed ozone

decreases. Randall et al. (2001) used POAM III observations to show middle stratospheric ozone decreases in

September 2000 up to 45% (at 33 km) as a result of the

solar proton precipitation in July 2000.

The long-lasting polar upper stratospheric ozone

depletion was measured to be >30% for late November

through December 2003 (Seppälä et al., 2004) and conjectured to be from the SPEs of October/November 2003.

The polar upper stratospheric ozone depletion of more than

60% measured in the spring of 2004 (Natarajan et al.,

2004; Randall et al., 2005b; López-Puertas et al., 2005a)

is likely connected with the huge enhancements of geomagnetic storm-generated NOx in November/December

2003 and January 2004 (see Section 4.4.1). Figure 4-20

(middle and bottom, adapted from Randall et al., 2005b)

shows the measured variations in ozone at 40 km (middle

plot) and partial column ozone between 35 and 50 km

(bottom plot). Total ozone decreases were measured to be

~0.5-1% (~1.5-3 DU, compared with unperturbed levels

of ~300 DU) in the spring and early summer of 2004 and

are correlated with the NO2 enhancements.

Models have been used to predict the impact of

the solar protons in solar cycle 23. Krivolutsky et al.

(2005) predicted significant short-term ozone changes as

a result of SPEs in July and November 2000, November

2001, and October/November 2003 with the use of a onedimensional (1-D) photochemical model. Jackman et al.

(2005b) simulated the influences of all the solar proton

fluxes between 2000 and 2003 with a two-dimensional

(2-D) CTM and predicted polar total ozone decreases of

~0.5-3% (~2-10 DU). These computed depletions

resulted from the large solar proton fluxes in 2000, 2001,

and 2003. The ozone decreases last beyond the SPEs

and gradually diminish over several years as the NOy

(mainly in the form of nitric acid (HNO3)) is transported

to the troposphere and rained out from the atmosphere

through wet deposition or sedimentation.

The electron impacts associated with geomagnetic

storms are more difficult to simulate, since the magnitude

of the precipitating electrons is hard to measure. However,

Rozanov et al. (2005) and Langematz et al. (2005) simulated the impact of energetic electron precipitation with

3-D CCMs and suggest that the solar cycle variation of

the electron flux could cause polar middle stratospheric

ozone changes of 20% or more. Sinnhuber et al. (2005a)

analyzed polar ozonesonde measurements with the use of

a 3-D CTM, which includes measured temperatures and

wind fields. They find that the residual ozone (modeled

minus measured) at 800 K correlates very well with

Geostationary Operational Environmental Satellite

(GOES)-measured electron fluxes greater than 2 MeV,

which are not in phase with 10.7-cm solar radio flux.

Precipitating electrons may be somewhat important in

determining long-term stratospheric ozone variability in

both the SH and in the NH during years when there is a

strong polar vortex accompanied by significant downward

transport from the mesosphere to the stratosphere (e.g.,

early 2004 in the NH) (Manney et al., 2005a; Semeniuk et

al., 2005).

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CHAPTER 5

Climate-Ozone Connections

Lead Authors:

M. Baldwin

M. Dameris

Coauthors:

J. Austin

S. Bekki

B. Bregman

N. Butchart

E. Cordero

N. Gillett

H.-F. Graf

C. Granier

D. Kinnison

S. Lal

T. Peter

W. Randel

J. Scinocca

D. Shindell

H. Struthers

M. Takahashi

D. Thompson

Contributors:

D. Battisti

P. Braesicke

R. Garcia

P. Haynes

E. Manzini

K. Matthes

G. Pitari

V. Ramaswamy

K. Rosenlof

B. Santer

R. Scott

A. Stenke

C. Timmreck

Final Release: February 2007

From Scientific Assesment of Ozone Depletion: 2006

CHAPTER 5

CLIMATE-OZONE CONNECTIONS

Contents

SCIENTIFIC SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.1

5.1 INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3

5.2 COUPLING OF THE STRATOSPHERE AND TROPOSPHERE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3

5.2.1 Radiation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3

5.2.2 Dynamics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.6

5.2.2.1 Role of Waves . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.6

5.2.2.2 How the Stratosphere Affects Its Own Variability . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.6

5.2.2.3 Transport of Air into the Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.7

5.2.2.4 Modeling and Parameterization of Small-Scale Waves . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.7

5.2.2.5 Annular Modes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.8

5.2.2.6 Effects of Stratospheric Variability on the Troposphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.8

5.2.3 Tropospheric Composition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.8

5.2.4 Stratospheric Aerosols . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.10

5.2.5 Past Changes in Stratospheric Water Vapor . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.12

5.2.6 Past Changes in Stratospheric Temperature . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.13

5.2.7 Impact of Ozone Changes on Surface Climate . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.16

5.2.7.1 Southern Hemisphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .5.16

5.2.7.2 Northern Hemisphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.18

5.3 EFFECTS OF ANTHROPOGENIC CLIMATE CHANGE AND OF EMISSIONS 

ON STRATOSPHERIC OZONE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.18

5.3.1 Stratospheric Temperature Changes Due to Shifts in Radiative Forcing . . . . . . . . . . . . . . . . . . . . . . 5.18

5.3.2 Stratospheric Temperature Changes Due to Additional Effects . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.23

5.3.3 Troposphere-Stratosphere Exchange and the Brewer-Dobson Circulation . . . . . . . . . . . . . . . . . . . . 5.26

5.3.4 Changes in the Tropical and Extratropical Tropopause Layer . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.29

5.3.4.1 The Tropical Tropopause Layer (TTL) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.29

5.3.4.2 The Extratropical Tropopause Layer (ExTL) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.31

5.3.5 Impact of Future Water Vapor Changes on Ozone Chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.31

5.3.6 Impact of Temperature Changes on Ozone Chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.32

5.3.6.1 Upper Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.32

5.3.6.2 Polar Lower Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.33

5.3.7 Impact of Climate Change on Ozone Recovery . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.34

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.38

APPENDIX 5A: AOGCMs Used in Chapter 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.49

CLIMATE-OZONE CONNECTIONS

5.1

SCIENTIFIC SUMMARY

Climate change will affect the evolution of the ozone layer through changes in transport, chemical composition,

and temperature. In turn, changes to the ozone layer will affect climate through radiative processes, and consequential

changes in temperature gradients will affect atmospheric dynamics. Therefore, climate change and the evolution of the

ozone layer are coupled. Understanding all of the processes involved is made more complex by the fact that many of the

interactions are nonlinear.

Impact of Climate Change

• The stratospheric cooling observed during the past two decades has slowed in recent years. Satellite and

radiosonde measurements reveal an overall cooling trend in the global-mean lower stratosphere of approximately

0.5 K/decade over the 1979-2005 period, with a slowdown in the temperature decline since the late 1990s. The

overall temperature decrease is punctuated by transient warmings of the stratosphere associated with the major volcanic eruptions in 1982 and 1991. Model calculations suggest that the observed ozone loss is the predominant

cause of the cooling observed over this period. The lower stratospheric cooling is evident at all latitudes, in particular in both Arctic and Antarctic winter/spring lower stratosphere but with considerable interannual variability.

Satellite observations show larger temperature trends in the upper stratosphere, with values of –1 to –2 K/decade,

but little additional decline since the middle 1990s. Model calculations suggest that the upper stratosphere trends

are due, roughly equally, to decreases in ozone and increases in CO2.

• Future increases of greenhouse gas concentrations will contribute to the average cooling in the stratosphere.

Estimates derived from climate models (AOGCMs, coupled ocean-atmosphere general circulation models) and

Chemistry-Climate Models (CCMs) with interactive ozone consistently predict continued cooling of the global

average stratosphere. The predicted cooling rate within the next two decades is dependent on the prescribed scenario and the type of model used for the assessment. At 50 hPa an average of all AOGCMs gives approximately 0.1

K/decade, while CCMs predict a larger cooling of about 0.25 K/decade caused by the interactive consideration of

ozone changes. All models calculate a stronger cooling at 10 hPa, averaging approximately 0.5 K/decade. Polar

temperatures in the future are less certain than global mean temperatures because of greater natural variability. 

• Chemical reaction rates in the atmosphere are dependent on temperature, and thus the concentration of

ozone is sensitive to temperature changes. Decreases in upper stratospheric temperature slow the rate of photochemical ozone destruction in this region. Hence the concentration of upper stratospheric ozone increases in

response to cooling. Cooling of the polar lower stratosphere would lead to more efficient chlorine activation on

aerosol and polar stratospheric clouds and enhanced ozone destruction. Therefore, the concentration of ozone in the

springtime polar lower stratosphere would decrease in response to cooling.

• Greenhouse-gas-induced temperature and circulation changes are expected to accelerate global ozone

increases in the next decades. Two-dimensional latitude-height models, as well as CCMs, show that 1980 global

mean total ozone values will be reached several years earlier than in a constant-temperature stratospheric environment. 

Impact on the Troposphere

• Changes to the temperature and circulation of the stratosphere affect weather and climate of the troposphere. The response is seen largely as changes to the strength of the surface westerly winds in midlatitudes, and is

found in both observations and model results. The strongest evidence for coupling is seen in the Northern

Hemisphere during winter and in the Southern Hemisphere during spring. We do not have a complete understanding

of the mechanisms that cause the stratosphere to affect the troposphere.

• Stratospheric ozone depletion in the Southern Hemisphere appears to have caused circulation changes not

only in the stratosphere, but in the troposphere as well. The observed cooling of the Antarctic lower stratosphere has led to an increase in the speed of the stratospheric westerly winds and an associated delay in the seasonal

breakdown of the stratospheric polar vortex. Observations and model results suggest that the changes to the lower
CLIMATE-OZONE CONNECTIONS

5.2

stratosphere have contributed to the observed strengthening of midlatitude tropospheric winds and to cooling over

the interior of Antarctica during December-February. As ozone recovers, tropospheric changes due to ozone loss

are expected to reverse. However, temperature changes due to increasing greenhouse gas concentrations may offset

this reversal.

Importance of Tropospheric Changes

• Human activities are expected to affect stratospheric ozone through changes in emissions of trace gases.

Enhanced methane (CH4) emission (from wetter and warmer soils) is expected to enhance ozone production in the

lower stratosphere, whereas a climate-driven increase in nitrous oxide (N2O) emission is expected to reduce ozone

in the middle and high stratosphere. Also, changes in nonmethane hydrocarbons and nitrogen oxide (NOx) emissions are expected to affect the tropospheric concentrations of hydroxyl radical (OH) and, hence, impact the lifetime

and concentration of stratospheric trace gases such as CH4 and organic halogen species.

• The exchange of air between the troposphere and the stratosphere is predicted to increase due to climate

change. Model studies predict that the annual mean troposphere-to-stratosphere mass exchange rate is expected to

increase significantly, which will also decrease the average time that air remains within the stratosphere. Another

possible consequence is a counterbalance of the stratospheric cooling associated with increasing greenhouse gases

by an increase in the descent and adiabatic heating in the polar stratosphere during winter and spring. The net effect

could result in local stratospheric temperature increases confined to high northern latitudes during winter and spring. 

Importance of Water Vapor

• Updated datasets of stratospheric water vapor concentrations now show differences in long-term behavior.

Recent trend analyses, which are based on only two available multiyear datasets, casts doubt on the positive stratospheric water vapor trend that was noted in the previous Assessment. Balloonborne measurements at Boulder,

Colorado, for the period 1980-2005 show a significant increase of 5-10% per decade over altitudes of 15-28 km.

Global measurements from the Halogen Occultation Experiment (HALOE) satellite instrument for 1991-2005 do

not show corresponding positive lower stratospheric trends. Interannual water vapor changes derived from HALOE

data exhibit quantitative agreement with temperature variations near the tropical tropopause. In contrast, the longterm increases inferred from the Boulder data are larger than can be explained by observed tropopause temperature

changes or past increases in tropospheric methane. 

• Future changes of stratospheric water vapor concentrations are uncertain. If water vapor concentration

increases in the future, there will be both radiative and chemical effects. Modeling studies suggest increased water

vapor concentrations will enhance odd hydrogen (HOx) in the stratosphere and subsequently influence ozone depletion. Increases in water vapor in the polar regions would raise the temperature threshold for the formation of polar

stratospheric clouds, potentially increasing springtime ozone depletion. 

Importance of Volcanoes

• If a major volcanic eruption occurs while stratospheric halogen loading is elevated, ozone will temporarily be

depleted. Strong volcanic eruptions enhance stratospheric aerosol loading for two to three years. A global average

increase of lower stratospheric temperature (about 1 K at 50 hPa) was observed following the eruptions of El

Chichón and Mt. Pinatubo, and globally averaged total ozone significantly decreased by about 2% before recovering after about two to three years. Ozone destruction via heterogeneous reactions depends on halogen loading, so

the effects on ozone of a major eruption are expected to decrease in the coming decades. For sufficiently low

halogen loading, a large volcanic eruption would temporarily increase ozone. Long-term ozone recovery would not

be significantly affected. 
5.1 INTRODUCTION

The purpose of this chapter is to assess the effects

of human-induced climate change and greenhouse gases

on stratospheric ozone. Investigations of the relationships

and feedbacks between ozone depletion and climate

change processes have demonstrated that it is not possible

to achieve a complete understanding of ozone changes

without the consideration of climate change. This chapter

primarily concentrates on how climate change affects stratospheric ozone. The effect of stratospheric ozone depletion on climate was a focus of the Intergovernmental Panel

on Climate Change/Technology and Economic Assessment Panel special report (IPCC/TEAP, 2005).

An increase of well-mixed greenhouse gas concentrations in the atmosphere leads to higher tropospheric

temperatures (the greenhouse effect) and lower stratospheric temperatures. The rates of many chemical reactions are temperature dependent, and these reaction rates

affect the chemical composition of the atmosphere.

Reduced stratospheric temperatures lead to a slowing of

some gas-phase reactions that destroy ozone, but also lead

to intensified depletion of ozone in the lower polar stratosphere due to increased activation of halogens on polar

stratospheric clouds (PSCs). 

Since climate change processes influence the

dynamics of the troposphere and the stratosphere, dynamically induced temperature changes could locally reinforce

or oppose the temperature changes caused by radiative

processes. These future changes are highly uncertain, with

some models projecting that temperature will increase in

the polar regions during northern winter and spring. The

net effect of radiative, chemical, and dynamical interactions and feedbacks (many of which are nonlinear) is

poorly understood and quantified at present. 

Results of investigations presented in this chapter

are based on observations and numerical modeling studies.

Although atmospheric models have improved in recent

years, they are still subject to uncertainties due to an

incomplete description of atmospheric processes, their

forcing, and their feedbacks. Weaknesses of models must

be considered when evaluating calculated future changes,

in particular for the assessment of the future evolution of

the stratospheric ozone layer, as discussed in Chapter 6. 

A summary description of the types of models used

in this Assessment is given in Box 5-1. Section 5.2 provides an overview of stratospheric processes and a basic

description of the coupling of the stratosphere and the troposphere. It contains background information about

mechanisms and key processes that are relevant to describe

and explain climate-ozone connections and feedbacks. In

addition, it describes the influences of water vapor and of

its changes, sulfate aerosol, changes in source gases, temperature trends, their feedbacks on both chemistry and

dynamics, and comparisons to results derived from numerical models describing atmospheric processes. Section

5.3 focuses on interactions between human-induced climate change and ozone depletion. At the end a detailed

discussion of results derived from Chemistry-Climate

Models is presented, which specifically addresses the

question of how climate change will affect the evolution

of the ozone layer.

5.2 COUPLING OF THE STRATOSPHERE AND

TROPOSPHERE

5.2.1 Radiation

Greenhouse gases (GHGs), mainly carbon dioxide

(CO2) and water vapor, warm the troposphere by absorbing

outgoing infrared (IR) radiation from the Earth in the wellknown greenhouse effect. The dominant balance in the

troposphere is between latent heating and radiative cooling

by greenhouse gases. In the stratosphere, however,

increased greenhouse gases lead to a net cooling as they

emit more IR radiation out to space than they absorb. IR

emission increases with local temperature, so the cooling

effect increases with altitude, maximizing near the

stratopause, where stratospheric temperatures are highest.

The stratospheric cooling effect of greenhouse gases varies

with latitude, as it depends on the balance between absorption of IR from below and local emission. The net cooling

effect of greenhouse gases extends to lower levels at high

latitudes, roughly following the tropopause. 

Any change in radiatively active gas concentrations

will change the balance between incoming solar (shortwave) and outgoing terrestrial (longwave) radiation in the

atmosphere. The change of this balance due solely to the

species in question, keeping other climate variables fixed,

is termed radiative forcing (WMO, 2003; IPCC/TEAP,

2005). Radiative forcing is conventionally given as the

net change in radiative fluxes at the tropopause, which can

be a reasonable indicator of the surface temperature

response.

Ozone absorbs both shortwave and longwave radiation. To determine radiative forcing from stratospheric

ozone changes, it is important to distinguish between

instantaneous effects and the effects after the stratospheric

temperature has adjusted. Depletion of ozone in the lower

stratosphere causes an instantaneous increase in the shortwave solar flux at the tropopause and a slight reduction of

the downwelling longwave radiation. The net instantaneous effect is a positive radiative forcing. However, the

CLIMATE-OZONE CONNECTIONS

5.3
CLIMATE-OZONE CONNECTIONS

5.4

Box 5-1. Atmospheric Models

Numerical models are useful for investigations of the composition and the thermal and dynamical structure of

Earth’s atmosphere. They allow evaluation of different processes and mechanisms as well as feedbacks. Scientific

progress can be achieved by understanding the discrepancies between observations and results derived from model

simulations. Assessments of the future development of atmospheric dynamics and chemistry are typically based on

scenario simulations and sensitivity studies. In this 2006 Assessment, results of the following model systems have

been used: 

• Two-Dimensional (2-D) Photochemical Model: Zonally averaged representation of the atmosphere, with

detailed chemistry but simplified transport and mixing. Chemical reactions are included in the model according

to the physical characteristics: pressure, temperature and incident solar radiation. In each model box, the

movement of air into and out of each box is simulated, representing advection and dispersion. Advection by

three-dimensional (3-D) motion and sub-grid scale mixing are parameterized. Some models include emissions

from different sources, particularly for tropospheric pollutants, otherwise they use imposed tropospheric concentrations. In an “interactive” model, changes in the chemical composition of the atmosphere cause changes

in temperatures and hence transport, whereas in a “non-interactive” model, this feedback is missing and temperatures are unaffected by changes in chemical composition.

• Chemical Transport Model (CTM): Simulation of chemical processes in the atmosphere employing meteorological analyses derived from observations or climate models. A CTM is a non-interactive model that does not

consider the feedback of chemistry to dynamical and radiative processes. It uses winds and temperatures from

meteorological analyses or predictions to specify the atmospheric transport and temperatures and to calculate

the abundances of chemical species in the troposphere and stratosphere. A CTM can be used to simulate the

evolution of atmospheric composition and help interpret observations. 

• Atmospheric General Circulation Model (AGCM): Three-dimensional model of large-scale (spatial resolution of a few hundred km) physical, radiative, and dynamical processes in the atmosphere over years and

decades. An AGCM is used to study changes in natural variability of the atmosphere and for investigations of

climate effects of radiatively active trace gases (greenhouse gases) and aerosols (natural and anthropogenic),

along with their interactions and feedbacks. Usually, AGCM calculations employ prescribed concentrations of

radiatively active gases, e.g., carbon dioxide (CO2), methane (CH4), nitrous oxide (N2O), chlorofluorocarbons

(CFCs), and ozone (O3). Changes of water vapor (H2O) concentrations due to the hydrological cycle are

directly simulated by an AGCM. Sea surface temperatures (SSTs) are prescribed. An AGCM coupled to an

ocean model, commonly referred to as an AOGCM or a climate model, is used for investigation of climate

change. More recently, climate models may also include other feedback processes (e.g., carbon cycle, interaction with the biosphere). 

• Chemistry-Climate Model (CCM): An AGCM that is interactively coupled to a detailed chemistry module (see

Figure 5-1). In a CCM, the simulated concentrations of the radiatively active gases are used in the calculations of net heating rates. Changes in the abundance of these gases due to chemistry and advection influence

heating rates and, consequently, variables describing atmospheric dynamics such as temperature and wind.

This gives rise to a dynamical-chemical coupling in which the chemistry influences the dynamics (via radiative heating) and vice versa (via temperature and advection). Not all CCMs have full coupling for all chemical

constituents; some radiatively active gases are specified in either the climate or chemistry modules. Ozone is

always fully coupled, as it represents the overwhelmingly dominant radiative-chemical feedback in the stratosphere. Transient simulations consider observed or predicted gradual changes in concentrations of radiatively

active gases and other boundary conditions (e.g., emissions). The temporal development of source gas emissions and SSTs are prescribed for a specific episode (years to decades). In time-slice simulations, the internal

variability of a CCM can be investigated under fixed conditions, e.g., for greenhouse gas (GHG) concentrations and SSTs, to estimate the significance of specific changes. 
decrease in ozone causes less absorption of solar and longwave radiation, leading to a local cooling. After the stratosphere has adjusted, the net effect of ozone depletion in

the lower stratosphere is a negative radiative forcing

(IPCC/TEAP, 2005). In contrast, ozone depletion in the

middle and upper stratosphere causes a slight positive

radiative forcing (IPCC/TEAP, 2005). The maximum sensitivity of radiative forcing for ozone changes is found in

the tropopause region, and the maximum sensitivity of

surface temperatures to ozone changes also peaks near the

tropopause (Forster and Shine, 1997). 

Quantifying the impact of stratospheric ozone

changes on surface temperatures is less straightforward

than estimating radiative forcing (RF). This is because a

climate sensitivity term, λ, has to be introduced to translate radiative forcing to changes in surface temperatures

(Tsurf). The relationship between λ and Tsurf changes

is generally given by ¢Tsurf = λ RF (see Box 1.3 in

IPCC/TEAP, 2005), where λ is in units of K (W m-2)-1.

¢Tsurf is the equilibrium response of global mean surface

temperature and RF is the radiative perturbation. The

assumption of linearity between radiative forcing and the

surface temperature change made in this equation is found

to hold well for ozone depletion (Forster and Shine, 1999).

λ is poorly constrained by observations and is conventionally evaluated in climate models as the equilibrium

global mean temperature response to a radiative forcing

change equivalent to a doubling of CO2. λ is dependent

on the strength of climate feedbacks, such as those associated with clouds, water vapor, and ice albedo, and its magnitude varies considerably from model to model. Its value

is likely to lie in the range 1.5 to 4.5 K (W m-2)-1 (IPCC,

2001). λ can depend significantly on the nature of the

forcing, particularly in the case of stratospheric ozone

changes, and can differ from that for a doubling of CO2.

Section 5.2.7 discusses dynamic responses in more detail. 

Eleven-year solar ultraviolet (UV) irradiance variations have a direct impact on the radiation and ozone

budget of the middle atmosphere (e.g., Haigh, 1994; see

also Chapter 3, Section 3.4.4). During years with maximum solar activity, the solar UV irradiance is enhanced,

which leads to additional ozone production and heating in

the stratosphere and above. By modifying the meridional

temperature gradient, the heating can alter the propagation of planetary and smaller-scale waves that drive the

global circulation. Although the direct radiative forcing

of the solar cycle in the upper stratosphere is relatively

weak, it could lead to a large indirect dynamical response

in the lower atmosphere through a modulation of the polar

night jet and the Brewer-Dobson circulation (Kodera and

Kuroda, 2002). Such dynamical changes can feed back

on the chemical budget of the atmosphere because of the

temperature dependence of both the chemical reaction

rates and the transport of chemical species. 

The Arctic lower and middle stratosphere tend to

be cold and undisturbed during west-wind phases of the

equatorial quasi-biennial oscillation (QBO; Section

5.2.2.1), while they are warm and disturbed during QBO

east-wind phases (Holton and Tan, 1980; 1982). Further

analysis (Labitzke, 1987; Labitzke and van Loon, 1988)

showed that this relationship is strong during solar minimum conditions, while during solar maximum years the

relationship does not hold. This solar-QBO interaction

has remained robust in the observations since its discovery.

Equatorial upper stratospheric winds during the early

winter appear to be important for the evolution of the

Northern Hemisphere winter, especially the timing of stratospheric sudden warmings (Gray et al., 2001a; 2001b;

Gray, 2003; Gray et al., 2004). 

CLIMATE-OZONE CONNECTIONS

5.5

Emissions of

Natural and

Anthropogenic

Gases

Volcanic and

Nonvolcanic

Aerosols

Solar Cycle

Sea Surface

Temperatures

Concentrations

o Radiatively

Active Gases

Chemistr y Dynamics

Temperature

and Wind

Radiation Photolysis and

Heating Rates

Figure 5-1. Schematic of a Chemistry-Climate Model

(CCM). The core of a CCM (oval symbols) consists

of an atmospheric general circulation model (AGCM)

that includes calculation of the heating and cooling

rates and a detailed chemistry module. They are

interactively coupled. Photolysis rates are calculated

online or are determined from a lookup table. Arrows

indicate the direction of effect. Rectangular boxes

denote external impacts. In current CCMs, sea surface temperatures (SSTs) are prescribed based on

observations or are adopted from calculations with a

climate model. Natural and anthropogenic emissions

of gases are considered. Tropospheric and stratospheric aerosol loading (especially after volcanic

eruptions) can be taken into account. CCMs often

consider the changes of solar radiation caused by

the 11-year activity cycle of the Sun.
Some modeling studies have confirmed the modulation of the polar night jet and the Brewer-Dobson circulation by the solar cycle (e.g., Matthes et al., 2004). The

transfer of the solar signal from the stratosphere to the troposphere is the subject of ongoing research and includes

the possibility of the modulation of the Northern

Hemisphere Annular Mode (NAM; Section 5.2.2.5)

(Kodera, 2002; Matthes et al., 2006) and changes in vertical motion and precipitation in the tropics (e.g., Kodera,

2004; Haigh et al., 2005; Matthes et al., 2006). It is also

possible that a “wave-ozone feedback” mechanism communicates the solar signal to the QBO (Cordero and

Nathan, 2005), although Mayr et al. (2006) found a solar

modulation of the QBO even though there is no waveozone feedback in their model.

5.2.2 Dynamics 

Although the stratosphere and troposphere are in

many ways distinct, the atmosphere is continuous,

allowing vertical wave propagation and a variety of other

dynamical interactions between these regions. A complete

description of atmospheric dynamics requires an understanding of both of these layers. The dynamical coupling

of the stratosphere and troposphere is primarily mediated

by wave dynamics. A variety of waves originate in the

troposphere, propagate upward into the stratosphere and

above, and then dissipate, shaping the spatial and temporal

structure of the stratospheric flow. This traditional view

of a passive stratosphere has more recently given way to a

greater appreciation of the stratosphere’s ability to shape

not only its own evolution but that of the troposphere as

well.

5.2.2.1 ROLE OF WAVES

The climatological temperature structure of the stratosphere, as well as its seasonal cycle and variability, depend

crucially on the dynamics of waves that are generated in

the troposphere. Wave dynamics can be divided broadly

into three processes: generation mechanisms, propagation

characteristics, and dissipation (primarily due to wave

breaking and thermal damping, with thermal damping

important for planetary wave dissipation at high latitudes,

outside the surf zones; it is also likely to be important for

the dissipation of equatorial waves). In the extratropics,

the temperature structure of the stratosphere depends on a

balance between diabatic radiative heating and adiabatic

heating from induced vertical motion due to planetary wave

dissipation (Andrews et al., 1987). Planetary wave

breaking in the winter stratosphere (and mesosphere) generally produces a westward force that decelerates the polar

stratospheric jet, resulting in ascent (adiabatic cooling) in

the tropics, and descent (adiabatic warming) over the poles

(Holton et al., 1995). This response pattern describes a

meridional mass circulation that is called the BrewerDobson circulation. 

The basic climatology of the extratropical stratosphere is broadly understood in terms of wave dynamics

together with the seasonal cycle of radiative heating. For

example, the easterly winds of the summer stratosphere

inhibit upward propagation of planetary waves (Charney

and Drazin, 1961) and so the summer stratosphere is much

less disturbed than the winter stratosphere. Asymmetries

in the continental land mass between the Northern

Hemisphere (NH) and Southern Hemisphere (SH) imply

asymmetries in the efficiency of planetary wave generation mechanisms. Consequently in winter, planetary wave

disturbances in the stratosphere of the NH are significantly

larger than those in the SH. In the tropical stratosphere,

the dominant form of variability is a quasi-periodic (2-3

year) wave-driven descending zonal mean wind reversal,

called the quasi-biennial oscillation (QBO). The peak-topeak amplitude of the wind QBO is ~55 m/s at 25-30 km

(Baldwin and Gray, 2005), while the temperature QBO

amplitude is ~8 K. The QBO affects the global stratospheric circulation, and extends to ~20° north and south

with an amplitude of about 10 m/s (Dunkerton and Delisi,

1985). It affects a variety of extratropical phenomena

including the strength and stability of the wintertime polar

vortex, and the distribution of ozone and other gases (see

Baldwin et al., 2001 for a review). The QBO is driven by

the dissipation of a variety of equatorial waves (Lindzen

and Holton, 1968; Dunkerton, 2001) that are primarily

forced by deep cumulus convection in the tropics. 

5.2.2.2 HOW THE STRATOSPHERE AFFECTS ITS OWN

VARIABILITY

Although stratospheric variability has long been

viewed as being caused directly by variability in tropospheric wave sources, it is by now widely accepted that

the configuration of the stratosphere itself also plays an

important role in determining the vertical flux of wave

activity from the troposphere. The original theory of

Charney and Drazin (1961) states that it is only when the

winds are westerly that the longest waves (mainly waves

1, 2, and 3) can propagate vertically. This theory has been

extended to account for the strongly inhomogeneous

nature of the stratospheric background state and the steep

potential vorticity gradients at the polar vortex edge (e.g.,

Scott et al., 2004). Given a steady source of waves in the

troposphere, any changes in the stratospheric background

potential vorticity (PV) gradient will change the vertical

CLIMATE-OZONE CONNECTIONS

5.6
wave fluxes, giving rise to the possibility of internally

driven variability of the stratosphere, as modeled by

Holton and Mass (1976). Similar internal variability has

also been demonstrated in more comprehensive AGCMs

(see Box 5-1) (e.g., Christiansen, 1999). Modeling studies

suggest that realistic stratospheric variability can arise in

the absence of tropospheric variability (Scott and Polvani,

2004; 2006).

The modulation of vertical wave flux into the stratosphere by the stratospheric configuration may be related

to the extent to which the stratosphere can act as a resonant cavity, involving downward reflection of stationary

planetary waves (McIntyre, 1982; Smith, 1989). Further,

through these processes, the tropospheric circulation itself

is also influenced by the stratospheric configuration.

Reflection of stationary planetary wave energy takes place

when the polar vortex exceeds a critical threshold in the

lower stratosphere, leading to structural changes of the

leading tropospheric variability patterns (Perlwitz and

Graf, 2001; Castanheira and Graf, 2003; Walter and Graf,

2005).

There exists also an external or barotropic mode

whose potential impact on the stratospheric circulation (in

terms of its deceleration of the polar night vortex) is significantly larger than that of upward propagating waves.

Esler and Scott (2005) demonstrated the relevance of this

mode in wavenumber-2 major warmings in which the

vortex is split throughout the full depth of the stratosphere. 

5.2.2.3 TRANSPORT OF AIR INTO THE STRATOSPHERE

Planetary waves breaking in the stratosphere are

also important for the transport of species from the troposphere to and within the stratosphere. The mean circulation of the stratosphere is essentially a “wave-driven

pump” (see Holton et al., 1995, for a review) in which

stratospheric wave drag moves air poleward and downward over the polar cap. As a consequence, the air in the

tropical lower stratosphere rises slowly (0.2 to 0.3 mm/s)

and carries ozone-poor air from the troposphere higher

into the stratosphere. There, with increasing altitude, photochemical production becomes more effective. The

upwelling in the tropics is modulated by the seasonal cycle

and the tropical QBO phase (Baldwin et al., 2001). When

the QBO is westerly at 40-50 hPa, the ascent rate is lower,

there is more time for ozone production, and the tropical

ozone column is enhanced. In the subtropics and extratropics, transport of chemical species from the troposphere

to the lowermost stratosphere occurs through quasi-isentropic motion associated with synoptic-scale and

mesoscale circulations (e.g., baroclinic eddies, frontal circulations). In the same circulations, there is substantial

transport from stratosphere to troposphere. Quantification of this two-way (troposphere-to-stratosphere and

stratosphere-to-troposphere) transport has improved significantly over the last few years through modeling and

observations. (See, e.g., Stohl et al., 2003 for a review.)

However, significant quantitative uncertainty remains over

the role of small-scale circulations, e.g., convective systems, in transport from troposphere to stratosphere or vice

versa. The transport of air in the tropopause region is discussed in greater detail in Chapter 2, Section 2.4. 

5.2.2.4 MODELING AND PARAMETERIZATION OF SMALLSCALE WAVES

The consideration of wave dynamics in determining

the climatology of the stratosphere is very important. Any

systematic change in the generation, propagation, or dissipation of waves (both resolved and parameterized) will

result in systematic changes in the temperature structure

of the stratosphere. The capability to simulate the climatology and space-time changes of stratospheric properties

hinges critically on our ability to simulate highly nonlinear

wave dynamics in a reliable way. Atmospheric models

generally have inadequate horizontal resolution (for baroclinic eddies, interaction with ocean and land surface) and

vertical resolution (for the planetary wave propagation

characteristics and cross-tropopause transport). Another

issue is that deep convection (an important excitation

mechanism for waves that propagate into the stratosphere)

is a sub-grid-scale process that must be parameterized.

One of the most challenging aspects of modeling the

dynamical coupling of the troposphere and stratosphere is

the parameterization of unresolved waves (in particular,

non-orographic gravity waves) and their feedback on the

resolved flow. There are limited observational data to

constrain the tropospheric sources and basic middleatmosphere climatology of gravity waves (e.g., see reviews

by Fritts and Alexander, 2003 and Kim et al., 2003).

Parameterizations of ever-increasing complexity are being

developed to more realistically model the dynamics of these

waves, and the free parameters are used to reproduce

present-day climate. This raises a credibility issue when

these gravity-wave parameterizations are employed for the

purpose of climate change simulations (but see Section

5.3.2). With the advent of newer satellite temperature and

wind observations of global extent and higher spatial resolution (e.g., Wu, 2004; Eckermann et al., 2006), it is anticipated that current parameterizations of gravity waves will

be better constrained and more objectively validated. There

are also current efforts to specify gravity wave source

spectra in terms of fields calculated by the underlying

CLIMATE-OZONE CONNECTIONS

5.7
AGCM, such as frontal zones and convective heating (e.g.,

Charron and Manzini, 2002; Beres et al., 2005).

5.2.2.5 ANNULAR MODES

Annular modes are hemispheric spatial patterns of

climate variability characterized by north-south shifts in

mass between polar and lower latitudes (Thompson and

Wallace, 2000). Tropospheric signatures of stratospheric

variability are often well described by annular mode patterns (e.g., Baldwin and Dunkerton, 2001; Gillett and

Thompson, 2003). In both the stratosphere and troposphere, the annular modes explain a larger fraction of variance than any other pattern of climate variability in their

respective hemisphere. On month-to-month time scales,

annular variability at tropospheric levels is strongly coupled with annular variability at stratospheric levels

(Baldwin and Dunkerton, 1999; Thompson and Wallace,

2000). Thus, as noted in Section 5.2.2.6, time series of

the annular modes provide a convenient way to describe

some aspects of stratosphere-troposphere coupling. The

Northern Hemisphere Annular Mode (NAM) near Earth’s

surface is alternatively known as the Arctic Oscillation

(AO; Thompson and Wallace, 1998) and the North Atlantic

Oscillation (NAO; Hurrell, 1995). The Southern Hemisphere Annular Mode (SAM) is also referred to as the

Antarctic Oscillation and High Latitude Mode.

Recent studies suggest the annular modes reflect

feedbacks between the eddies and the zonal flow at middle

latitudes (Lorenz and Hartmann, 2001; 2003). Other

studies imply the annular modes are expected in any

rotating planetary fluid system that conserves momentum

and mass, and that has some smoothness property (Gerber

and Vallis, 2005). The key dynamics that underlie the

annular modes are still under investigation.

5.2.2.6 EFFECTS OF STRATOSPHERIC VARIABILITY ON

THE TROPOSPHERE

Observational analysis suggests that stratospheric

processes affect surface weather and climate (e.g., Scaife

et al., 2005). Figure 5-2 shows composites of indices of

the Northern Hemisphere Annular Mode (NAM) during

periods when the stratospheric vortex rapidly changes

strength. It reveals that within the winter season when the

stratospheric flow is westerly, changes to the strength of

the northern polar vortex are, on average, accompanied by

similarly signed and similarly persistent changes to the

tropospheric flow (Baldwin and Dunkerton, 1999; 2001).

The illustration would be similar if high-latitude zonal

winds (~60°N) were used instead of the annular mode

index. The figure thus suggests that changes to the strength

of the polar vortex—especially in the lowermost stratosphere—may affect the tropospheric flow.

Despite the apparent robustness of the above evidence, the principal mechanisms whereby stratospheric

variability may influence the tropospheric circulation

remain unclear. A complete explanation of the observed

coupling likely lies in one or more of the following physical processes:

1) Geostrophic and hydrostatic adjustment of the tropospheric flow to anomalous wave drag (Haynes et al.,

1991; Thompson et al., 2006) and anomalous diabatic

heating at stratospheric levels (Thompson et al.,

2006);

2) The impact of anomalous shear in the lower stratospheric zonal flow on the momentum flux by baroclinic eddies (Shepherd, 2002; Kushner and Polvani,

2004; Wittman et al., 2004);

3) Amplification due to internal tropospheric dynamics

(Song and Robinson, 2004);

4) The impact of anomalous shear at the tropopause level

on vertically propagating waves (Chen and Robinson,

1992; Shindell et al., 1999; Limpasuvan and Hartmann, 2000);

5) The reflection of planetary waves (Hartmann et al.,

2000; Perlwitz and Harnik, 2004). 

Through such mechanisms, long-term changes in

temperature and circulation of the stratosphere (e.g., radiative heating anomalies due to a GHG increase or changes

to the ozone distribution) may affect surface weather patterns, at least during winter, spring, and early summer.

5.2.3 Tropospheric Composition

Many of the chemical constituents present in the

stratosphere have sources that originate in the troposphere.

Any changes in the chemical composition of the troposphere can affect the composition of the stratosphere. The

chemical constituents are either directly emitted in the troposphere, mostly at or near the surface, or they are oxidation products of emitted species. The predominant source

gases for stratospheric hydrogen, halogen, and nonvolcanic sulfur are long-lived species (water vapor, methane

(CH4), nitrous oxide (N2O), organic halogen gases such as

chlorofluorocarbons (CFCs), halons, and carbonyl sulfide

(COS)). Surface emissions of short-lived species (sulfur

dioxide (SO2), dimethyl sulfide (DMS)) are also important sources of sulfur to the stratosphere, as are occasional

large volcanic eruptions. 

CLIMATE-OZONE CONNECTIONS

5.8
Long-term increases in CH4, N2O, and CFCs

brought about by increasing anthropogenic emissions are

discussed in Chapter 1. However, with the exception of

CFCs and halons, emissions of most stratospheric source

gases have a substantial natural component. For instance,

natural emissions represent more than a third of the CH4

source, more than half of the N2O source (IPCC, 2001),

and are the dominant source of carbonyl sulfide (COS)

(SPARC, 2006). As natural emissions are very likely to

be affected by climate changes and, in particular, changes

in precipitation, vegetation, and temperature, it is important to assess their sensitivity to climate changes, taking

into account this effect when forecasting their future evolution and their overall impact on ozone recovery. Natural

emissions also represent a large source of key short-lived

species, such as nitric oxide (NO) emissions from soil and

lightning, or emissions of nonmethane hydrocarbons such

as isoprene. Sources, transport, and stratospheric impact

of halogenated short-lived species, many of which have

predominantly natural origins, are dealt with in Chapter 2.

The lifetimes of such short-lived compounds in relation to

transport time scales are such that only a fraction of surface emissions reach the stratosphere (see Chapter 2).

They may also affect the chemical composition of the troposphere, and hence potentially the lifetimes of other trace

gases, through changes in atmospheric hydroxyl radical

(OH). 

Some progress has been made in estimating the sensitivity of natural sources to various climate parameters.

For example, process modeling suggests that CH4 emissions from wetlands could increase by 20% for a temperature increase of 1 K (Walter and Heimann, 2000). The

CLIMATE-OZONE CONNECTIONS

5.9

Composite of Weak Vortex Events

-90 -60 -30 0 30 60 90

Lag (Days)

hPa

a 10

30

100

300

1000

km

0 

10

20

30

Composite of Strong Vortex Events

-90 -60 -30 0 30 60 90

Lag (Days)

hPa

b 10

30

100

300

1000

km

0 

10

20

30

Figure 5-2. Composites of time-height development of the Northern Hemisphere Annular Mode (NAM) for: (a)

22 weak vortex events; and (b) 35 strong vortex events during 1958 to 2006. Updated from Baldwin and

Dunkerton (2001). The events are determined by the dates on which the 10 hPa NAM values crossed –3.0

and +1.5, respectively. The indices are non-dimensional; the contour interval for the color shading is 0.25, with

values between –0.25 and 0.25 unshaded. The white contours begin at ±1.25 with a contour interval of 0.5.

The thin horizontal lines indicate the approximate tropopause. The diagrams illustrate that large anomalies in

the strength of the polar vortex at 10 hPa tend to descend to the lowermost stratosphere, where they last, on

average, more than two months. After the stratospheric events occur, the tropospheric NAM anomaly is of the

same sign as the stratospheric anomaly.
emissions of short-lived species such as biogenic hydrocarbons also increase with temperature. However, it is

difficult to quantify climate-driven changes in natural

sources because the temperature is not the only driving

factor. Other factors such as water table level, soil moisture, vegetation cover, photosynthetically active radiation,

biogenic productivity, or exposure to atmospheric pollutants can play a role, depending on the emitting substrate

and the emitted species. Ignoring changes in land use,

most natural emissions are expected to increase as Earth’s

surface warms. The changes that are most relevant to the

stratosphere are climate-driven increases in CH4 and N2O

emissions. An increase in CH4 would accelerate the ozone

recovery whereas an increase in N2O would delay it

(Randeniya et al., 2002; Chipperfield and Feng, 2003). 

Climate changes can also alter other key processes

in the exchange of chemical constituents between the troposphere and stratosphere. The vertical transport of surface emissions to the tropopause is largely dependent on

the intensity of convective activity, and the flux of tropospheric air into the stratosphere is mostly determined by

the strength of the upwelling from the troposphere that is

linked to the strength of the Brewer-Dobson circulation

(see Section 5.2.2.3). 

5.2.4 Stratospheric Aerosols 

Non-explosive volcanic surface emissions of sulfur

species (SO2, COS) are important sources of stratospheric

aerosol loading mainly through tropical stratosphere/troposphere exchange (Notholt et al., 2005). Stratospheric

aerosols have a direct radiative impact that decreases the

surface temperature, since more shortwave radiation is

reflected. Although there has been no significant change

in the background (nonvolcanic) stratospheric aerosols for

the period 1970 to 2004 (Deshler et al., 2006; see also

SPARC, 2006), the non-volcanic aerosol loading could

increase in the future if convection increases (Pitari et al.,

2002), because convection is a key process that transports

the short-lived species SO2 from the surface to the upper

troposphere and lower stratosphere. 

Volcanic SO2 injected into the stratosphere is oxidized to sulfuric acid that condenses and forms aerosols,

on which heterogeneous reactions occur. Volcanic eruptions have strong impacts on the lower stratospheric

thermal structure because the volcanic aerosols scatter

back incoming solar radiation and absorb solar nearinfrared and terrestrial infrared radiation (e.g., Stenchikov

et al., 1998; Al-Saadi et al., 2001; Rozanov et al., 2002;

Timmreck et al., 2003). Heterogeneous chemistry occurring on aerosol surfaces affects ozone concentrations, producing an additional indirect radiative impact depending

on the concentration of atmospheric chlorine. In addition,

the modified meridional temperature profile in the stratosphere may result in colder polar vortices in winter

(Chapter 3 in WMO, 2003).

The ozone impact of a given volcano depends on

the amount of material, in particular sulfur, injected by the

volcanic eruption and on whether the material reaches the

stratosphere, as well as the phase of the QBO and the latitude of the eruption. The height reached by ejecta depends

on the explosivity of the eruption, not on its location.

Ejecta from tropical eruptions will be carried upward and

poleward by the Brewer-Dobson circulation and, therefore, they will spread throughout much of the stratosphere

with a long residence time, whereas ejecta from mid- to

high-latitude eruptions will more quickly be returned to

the troposphere by the descending branch of the BrewerDobson circulation.

The effect of a future major volcanic eruption will

depend on chlorine levels. For low-chlorine conditions,

heterogeneous chemistry can lead to ozone increases in

the stratosphere, whereas for high chlorine conditions, as

observed in recent years, volcanic aerosols lead to additional ozone depletion (Tie and Brasseur, 1995).

Moreover, volcanic aerosol affects photolysis rates and

therefore ozone concentrations (Timmreck et al., 2003). 

The results of several CCM simulations (see Box

5-1) are consistent with the Tie and Brasseur study. After

the eruption of Mt. Agung in 1963, the total amount of

ozone was reduced, particularly in the tropics (see Chapter

3, Figure 3-4). Since the chlorine loading of the atmosphere was low at that time, the solar cycle influences were

of particular importance (see Chapter 3, Section 3.4.4).

For example, the CCM E39C simulated reasonably well

the observed decrease of tropical ozone around 1965,

since the influence of the 11-year solar cycle was considered, which was minimal around 1965 (Dameris et al.,

2005). Following the major eruptions of El Chichón

(1982) and Mt. Pinatubo (1991), when atmospheric chlorine amounts were much higher and solar activity was

near maximum, the total ozone significantly decreased in

subsequent months in both observations and models,

before recovering after two to three years. The simulated

globally averaged total ozone decreases for many models

in the Eyring et al. (2006) assessment were about 2%,

similar to what was observed (cf. Fioletov et al., 2002;

WMO, 2003). 

In AOGCM and CCM simulations, the temperature

perturbations after the eruptions of El Chichón and Mt.

Pinatubo are often larger than 1 K (annually averaged) in

the lower stratosphere (see Figures 5-3 and 5-11), whereas

observations indicate an increase of about 1 K. The temperature impact is important since this will determine to a

large extent the water vapor perturbation. In turn, this will

CLIMATE-OZONE CONNECTIONS

5.10
influence the ozone chemistry in the lower stratosphere

(Section 5.3.5) and lower stratospheric cooling rates

(Section 5.2.6). Figure 5-3 shows the globally averaged

temperature from a range of tropospheric climate models

(Santer et al., 2006; see also Ramaswamy et al., 2006),

vertically averaged by the Microwave Sounding Unit

MSU4 weighting function, together with MSU4 data. The

variations due to volcanic eruptions are clear for several

years, with an overall slow cooling. The CCM and climate model studies indicate that the strength of the volcanic signal varies substantially between the models (see

also Figures 3-26 and 5-11, and Eyring et al., 2006). 

The sulfate aerosols and ash injected into the stratosphere from volcanic eruptions can cause tropospheric

cooling (e.g., Hansen et al., 1992; Robock, 2000; Santer et

al., 2001; Wigley et al., 2005; Yokohata et al., 2005),

although ash particles have a short residence time because

of gravitational settling. The tropospheric cooling would

be expected to change the tropospheric circulation, as well

as the interaction between the stratosphere and the troposphere (Stenchikov et al., 2002). 

Based on the historical volcanic record, a major

eruption of similar atmospheric impact to that of Mt.

Pinatubo is likely to occur during the next 30 years

(Roscoe, 2001). The previous 50-year period, with three

such eruptions, appears to be unusual. Much larger eruptions, such as Toba 74,000 years ago, are a remote possibility, with recovery times of a decade (Bekki et al., 1996).

Model studies by Rosenfield (2003), assuming a Pinatubosized eruption every 10 years from 2010 to 2050, suggest

that the long-term recovery of ozone would not be strongly

affected by infrequent large volcanic eruptions. The

impacts on ozone of all except the largest eruptions would

be expected to last only a few years, and not significantly

CLIMATE-OZONE CONNECTIONS

5.11

1980 1985 1990 1995 2000

Time (years)

0.5

0

0.5

1.5

2

Temperature anomaly ( C)

CCSM3

GFDL CM2.0

GFDL CM2.1

GISS-EH

GISS-ER

MIROC3.2(medres)

MIROC3.2(hires)

PCM

UKMO-HadGEM1

Observations (RSS)

Observations (UAH)

Simulated and Observed Stratospheric Temperature Changes

Figure 5-3. Globally averaged temperature anomalies in the lower stratosphere from a range of climate

models (Santer et al., 2006) in comparison with Microwave Sounding Unit MSU4 channel temperatures. The

MSU4 observations are from two independent analyses by the Remote Sensing Systems (RSS) and the

University of Alabama at Huntsville (UAH) groups. See Appendix 5A for information about the models.
affect long-term ozone recovery, because ozone recovery

depends primarily on halogen levels.

5.2.5 Past Changes in Stratospheric Water

Vapor

Water vapor is the most important greenhouse gas,

and it plays an important role in chemistry-climate interactions. Previous studies reported in WMO (2003) showed

that ozone perturbations are amplified in the upper troposphere and lower stratosphere as a result of feedback

processes with water vapor. Any increase in stratospheric

water vapor could lead to an increase in the level of oddhydrogen radicals (HOx), which could affect the nitrogen

oxide radicals (NOx) and chlorine oxide radicals (ClOx)

cycles, leading to ozone depletion (see Section 5.3.5). A

change in water vapor concentration could also change the

temperature thresholds for polar stratospheric cloud formation over the polar caps (see Section 5.3.5).

Air enters the stratosphere mostly in the tropics, and

stratospheric water vapor is primarily controlled by temperatures near the tropical tropopause. The associated

processes have been studied in detail over the past few

years, for example using trajectory studies (Bonazzola and

Haynes, 2004; Fueglistaler et al., 2005). Water vapor is

also produced in the stratosphere by the photochemical

oxidation of methane, producing approximately two molecules of water vapor per molecule of methane. The

increase in the concentration of tropospheric methane

since the 1950s (0.55 parts per million by volume (ppmv))

is responsible for part of the increase in stratospheric water

vapor over this time period (SPARC, 2000). 

Measurements of stratospheric water vapor content

are available from ground-based instruments and aircraft

observations, plus balloonborne and satellite datasets. The

longest continuous dataset is from a single location

(Boulder, Colorado, USA), based on balloonborne frost

point hygrometer measurements (approximately one per

month), beginning in 1980 (Oltmans et al., 2000; see

updated data in Figure 5-4). Over the period 1980-2005,

a statistically significant linear trend of ~5-10% per decade

is observed at all levels between approximately 15 and 26

km. However, although a linear trend can be fitted to this

25-year long record, there is a high degree of variability in

the infrequent sampling, and the increases seen are neither

continuous nor steady. In particular, stratospheric water

vapor concentrations have decreased since 2001. Longterm increases in stratospheric water vapor content are

also inferred from a number of other datasets covering the

years 1980-2000 (Rosenlof et al., 2001), although the time

series are short and the sampling uncertainty is high in

many cases.

Recent work has focused on the interannual and longterm evolution of the water vapor distribution using nearglobal observations from the satellite-based Halogen

Occultation Experiment (HALOE). These measurements

span more than a decade (late 1991 to 2005). Interannual

changes in water vapor derived from HALOE data show

excellent agreement with the Polar Ozone and Aerosol

Measurement (POAM) satellite data (Randel et al., 2004b)

and also with the Stratospheric Aerosol and Gas Experiment

II (SAGE II) water vapor data (Thomason et al., 2004; Chiou

et al., 2006). An updated comparison of the HALOE measurements with the Boulder balloon data for the period 1992-

2005 is shown in Figure 5-4. The Boulder and HALOE data

show reasonable agreement for the early part of the record

(1992-1996), but the Boulder data are about 5% to 8%

higher after 1997. These differences are within the accuracies of both types of observations (reported to be around

10-20% in SPARC, 2000). Year-to-year water vapor

changes in each dataset appear to be correlated, and both

time series show the persistent decreases after 2001.

However, as a result of the differences after 1997, changes

CLIMATE-OZONE CONNECTIONS

5.12

Figure 5-4. Evolution of stratospheric water vapor

mixing ratio (in ppmv, averaged over 17-22 km) at

Boulder, Colorado (40°N, 105°W), derived from balloonborne frost point hygrometer measurements covering 1980-2005. The thin line shows a smooth fit

through the data points, using a running Gaussian

window with a half-width of three months. The heavy

line shows HALOE satellite water vapor data during

1992-2005 for the same altitude region, using measurements near Boulder (over latitudes 35°N-45°N,

and longitudes 80°W-130°W). Note the difference

between the two datasets after about 1997. Updated

from Randel et al., 2004a.
derived from the two datasets over the (short) overlap period

1992-2005 are very different, with net decreases in the

HALOE data but not in the Boulder record. These differences seem statistically significant, as Randel et al. (2004b)

report a statistical uncertainty of linear fits of the 1992-2002

record of less than ±0.5%/year. The reason for the differences between the balloon and satellite datasets (for the

same time period and location) is unclear. 

Interannual changes in the HALOE stratospheric

water vapor data during 1992-2005 are in quantitative

agreement with observed changes in tropical tropopause

temperatures for this period (Randel et al., 2004a;

Fueglistaler and Haynes, 2005). Tropopause temperature variations associated with the QBO (and to a lesser

degree the El Niño-Southern Oscillation, ENSO) are

echoed in observed water vapor changes for this period,

consistent with the modeling studies of Giorgetta and

Bengtsson (1999) and Geller et al. (2002). Furthermore,

the persistent decreases in stratospheric water vapor concentrations since 2001 are associated with anomalously

low tropopause temperatures (Randel et al., 2004a, 2006;

Fueglistaler and Haynes, 2005; see also Figure 5-21).

This agreement suggests a reasonable level of understanding for interannual water vapor changes during the

HALOE time period (and ability to project future values

based on tropical tropopause temperatures). In contrast,

the long-term water vapor increases inferred from the

Boulder balloon data since 1980 (and from combined

datasets beginning in the 1960s; Rosenlof et al., 2001)

are difficult to reconcile with observed long-term

decreases in tropical tropopause temperatures (e.g.,

Seidel et al., 2001). Only a fraction of the changes can

be attributed to increasing tropospheric methane, as discussed in detail in Fueglistaler and Haynes (2005). Thus,

while the HALOE record appears quantitatively well

understood, the long-term increases inferred from the

Boulder data over 1980-2005 (and combined datasets

since the 1960s) are larger than can be explained by

observed tropopause temperature changes or past

increases in tropospheric methane. 

5.2.6 Past Changes in Stratospheric

Temperature

Stratospheric temperature changes are closely coupled to ozone changes. Ozone is a key radiatively active

constituent in the stratosphere, and it is important to assess

the consistency between observed changes in ozone and

temperature. Also, the halogen-related ozone destruction

rate is generally reduced by lower temperatures in the

upper stratosphere, but increased by lower temperatures

in the polar lower stratosphere. 

Estimates of past temperature changes in the stratosphere have been derived from several different types of

data. Most of these datasets were not designed for climate monitoring purposes, and each has strengths and

limitations that require careful evaluation and scrutiny.

An important advance during the last several years is

increased quantification of trend uncertainties, accomplished by comparisons of independent datasets and

analyses. Temperature trends can also be derived from

meteorological analysis and reanalysis datasets, but evidence suggests these can be influenced by artificial

changes related to data inhomogeneity effects (e.g., Santer

et al., 2004; Randel et al., 2004a; Birner et al., 2006), and

hence may not be reliable in all applications.

Optimal detection and attribution techniques have

been widely used to attribute observed changes in surface and tropospheric temperatures to particular external

climate influences (Chapter 12 in IPCC, 2001). However, while some detection and attribution studies have

identified an anthropogenic influence in variables incorporating stratospheric temperature, such as radiosonde

temperature trends in the troposphere and stratosphere

(e.g., Tett et al., 2002; Thorne et al., 2002), or tropospheric height (Santer et al., 2003, 2004), none has

focused exclusively on stratospheric temperature. One

reason may be that the coupled ocean-atmosphere

models (AOGCMs; see Box 5-1) required for such

studies generally have limited stratospheric resolution,

and they underestimate stratospheric variability (Tett et

al., 2002). A realistic estimate of internal variability is

required in order to distinguish an externally forced

response. 

A nearly continuous record of stratospheric temperature measurements from satellites is available from

the series of operational NOAA satellites beginning in

1979. These are based on the Microwave Sounding Unit

(MSU) and Stratospheric Sounding Unit (SSU) instruments, which have flown on ten individual operational

satellites over 1979-2005. These data represent mean

temperatures for 10- to 15-km thick layers covering the

lower to upper stratosphere. The SSU data are available for three fundamental channels (25, 26, and 27),

together with several synthetic channels derived by differencing nadir and off-nadir measurements, which

provide increased vertical resolution (Nash, 1988).

Records of stratospheric temperatures for 1979-2005

are derived by combining data from the individual

instruments, adjusted for calibration effects using

periods of overlap between adjacent satellites. Effects

of orbital drift and decay, and the influence of aliasing

atmospheric tides also need to be considered in constructing long-term stratospheric datasets. There are

CLIMATE-OZONE CONNECTIONS

5.13
several independent analyses of the MSU Channel 4

data, which covers the lower stratosphere, including

results from the University of Alabama at Huntsville

(UAH, Christy et al., 2003) and Remote Sensing

Systems (RSS, Mears et al., 2003). At present there are

two analyses of the combined SSU dataset, with data

details discussed in Scaife et al. (2000) and Ramaswamy

et al. (2001).

Long-term temperature changes can also be evaluated from historical radiosonde data, for which stratospheric measurements (up to ~25-30 km) are available

since approximately 1960. Uncertainties in radiosondebased temperature trends are associated with spatial

sampling (the majority of measurements occur over NH

midlatitudes), and more importantly with changes

(improvements) in instrumentation over time, which can

result in artificial cooling biases (Gaffen, 1994; Luers and

Eskridge, 1998). Temperature trends calculated from

ensemble radiosonde datasets exhibit strong cooling in

the lower stratosphere (Lanzante et al., 2003b; Thompson

and Solomon, 2005; Free et al., 2005). However, these

trends are substantially larger than corresponding trends

derived from satellite measurements (Seidel et al., 2004)

or estimates from current model simulations (Santer et

al., 2005), and a recent U.S. Climate Change Science

Program Assessment (CCSP, 2006) concludes that the

ensemble radiosonde trends are probably influenced by

artificial cooling biases.

There is strong evidence for a large and significant

cooling in most of the stratosphere since 1980. Figure 5-

5 shows near-global average temperature anomalies

derived from satellite datasets for 1979-2005, spanning a

range of altitudes from the lower to upper stratosphere.

The vertical profile of near-global temperature trends

during 1979-2005 derived from these satellite data are

shown in Figure 5-6, in addition to trends derived from

radiosonde data, and a synthesis of model results (taken

from Shine et al., 2003). Both the satellite and radiosonde

datasets reveal an overall cooling of the stratosphere, with

trend values of about 0.5 K/decade in the lower stratosphere, increasing to larger values of about 1 to 2

K/decade in the upper stratosphere. There is reasonable

CLIMATE-OZONE CONNECTIONS

5.14

° °

Figure 5-5. Time series of nearglobal mean (60°N-60°S) temperature anomalies for 1979-2005

derived from satellite radiance

measurements. Results are

shown for MSU Channel 4 (representing the layer mean temperature over ~13-22 km), and SSU

channels 25 (~18-37 km), 26 (~25-

44 km), 27 (~34-51 km), and 47x

(~44-58 km). Vertical dashed lines

mark the occurrence of volcanic

eruptions of El Chichón (E) and Mt.

Pinatubo (P).
agreement between the lower stratospheric trends derived

from satellite data and radiosondes, although note that the

radiosonde results in Figure 5-6 are derived from a subset

of one particular homogenized dataset (see Figure 5-6

caption). The radiosonde data suggest there has been significant cooling of the lower stratosphere (70-30 hPa)

over most of the globe for 1979-2005, including the

tropics.

Temperature trends derived from both the MSU4

(Randel and Wu, 2006) and radiosonde datasets

(Thompson and Solomon, 2005; Free et al., 2005; Randel

and Wu, 2006) suggest the cooling of the lower stratosphere spans tropical and extratropical latitudes.

However, temperature trends derived from the SSU data

(as shown in WMO, 2003) suggest the cooling of the

lower stratosphere is restricted to the extratropics. The

differences between temperature trends derived from the

SSU, MSU4, and radiosonde data are currently under

investigation.

The time series in Figure 5-5 show that the observed

cooling is not a simple linear trend. A strong imprint of

transient warming (for 1-2 years) is observed in the lower

CLIMATE-OZONE CONNECTIONS

5.15

Figure 5-6. Vertical profile of temperature trends derived from satellite and radiosonde data over 60°N-60°S for

the period 1979-2005, together with a synthesis of model results taken from Shine et al. (2003). The satellite

results are shown for each individual SSU channel (as noted) and MSU Channel 4. For each channel the vertical

bar denotes the approximate altitude range sampled by that channel, and the horizontal bar denotes the (two

sigma) statistical trend uncertainty. Radiosonde results are shown for individual pressure levels (200-20 hPa),

and are derived from a subset of stations described in Lanzante et al. (2003a), updated as described in Randel

and Wu (2006). The subset of stations is chosen to omit stations with large artificial cooling biases, in particular

stations where MSU4 minus radiosonde trends are greater than 0.3 K/decade (taken from Table 1 of Randel and

Wu, 2006). Note that the Shine et al. (2003) results refer to model calculations and observations for the period

1979-1997, whereas the observed trends in the figure are for the longer period 1979-2005.
and middle stratosphere following the volcanic eruptions

of El Chichón (1982) and Mt. Pinatubo (1991). In the lower

stratosphere, the long-term cooling manifests itself as more

of a step-like change following the volcanic warming

events (Pawson et al., 1998; Seidel and Lanzante, 2004).

The overall lower stratospheric cooling is primarily a

response to ozone decreases (Shine et al., 2003; Langematz

et al., 2003), with a possible but much less certain contribution from changes in stratospheric water vapor.

Ramaswamy et al. (2006) suggest that the step-like time

series behavior is due to a combination of volcanic, solar

cycle, and ozone influences. There is a substantial flattening of these trends evident in Figure 5-5 after approximately 1995; the latter aspect agrees with small global

trends in the upper stratosphere and lower mesosphere

observed in HALOE data for 1992-2004 (Remsberg et al.,

2005). Although some flattening might be expected in

response to changes in the ozone trends, the strength of this

behavior is curious in light of continued increases in wellmixed greenhouse gases during this decade.

5.2.7 Impact of Ozone Changes on Surface

Climate

5.2.7.1 SOUTHERN HEMISPHERE

The largest stratospheric ozone depletion is

observed in the austral spring in the Antarctic stratosphere, and therefore it is here that one might expect any

effect of stratospheric ozone depletion on tropospheric

climate to be largest. While the largest stratospheric

temperature and geopotential height trends over the

Antarctic have been observed in November, coincident

in space with the maximum ozone depletion, significant

decreases in geopotential height also extend to the troposphere 1-2 months later (Thompson and Solomon,

2002) (Figure 5-7). 

Several modeling studies have examined the tropospheric response to prescribed changes in stratospheric ozone. Two studies using atmospheric models

with prescribed sea surface temperatures demonstrated

that a tropospheric response similar to the positive phase

of the SAM is simulated in the austral summer in

response to prescribed changes in stratospheric ozone

(Kindem and Christiansen, 2001; Sexton, 2001). Using

a model with high vertical resolution in the stratosphere

and a mixed-layer ocean, Gillett and Thompson (2003)

simulated a significant tropospheric geopotential height

response to prescribed stratospheric ozone changes

(Figure 5-7) with no greenhouse gas (GHG) changes.

The simulated and observed geopotential height and

temperature changes agreed well, both in magnitude and

in seasonality, supporting the hypothesis that the

observed trends were largely induced by stratospheric

ozone depletion. 

Transient simulations with coupled oceanatmosphere models (AOGCMs) also indicate that ozone

depletion has played an important role in inducing SAM

trends, particularly in the summer (Marshall et al., 2004;

Shindell and Schmidt, 2004; Arblaster and Meehl, 2006),

CLIMATE-OZONE CONNECTIONS

5.16

30 y li ea t e ds i Z

Month

30 y li ea t e ds i T

P ess e

20

260

100

100

20

180

340

20

2.5

0.5

4.5

6.5

0.5 0.5

-20

-260

-100

-420

-180

-340

Month

P ess e

-2.5

-0.5

-4.5

-6.5

Model Observations

from Thompson and Solomon 2002)

-8.5

Figure 5-7. Antarctic geopotential height (in m, top)

and temperature trends (in

K, bottom) by month and

pressure level in radiosonde

observations over the period 1969-1998 (right) and a

simulation of the response

to stratospheric ozone depletion (left). Shading indicates regions of significant

change based on (left

panels) exceedance of one

standard deviation of the

monthly anomalies and

(right panels) a two-sample

t-test at the 5% level. From

Gillett and Thompson, 2003.
although all three studies found that GHG changes have

played at least as large a role in inducing trends in the

annual mean SAM. 

The observed trend toward the positive phase of the

SAM in December to May has been associated with a surface cooling of the Antarctic interior of ~1 K, and a

warming of the Antarctic Peninsula, the Scotia Sea, and

the southern tip of South America (Thompson and

Solomon, 2002) (Figure 5-8). A similar pattern of warming

and cooling has been simulated in response to stratospheric

ozone depletion (Gillett and Thompson, 2003) and combined stratospheric ozone depletion and GHG increases

(Shindell and Schmidt, 2004; Arblaster and Meehl, 2006)

(Figure 5-8). Thompson and Solomon (2002) also identified an associated strengthening of the westerlies over the

Southern Ocean, corresponding to a poleward shift of the

storm track, which is also simulated in response to stratospheric ozone depletion (Gillett and Thompson, 2003;

CLIMATE-OZONE CONNECTIONS

5.17

Linear trends in Z500 Linear trends in SAT/wind

Model Observations from Thompson and Solomon 2002)

Figure 5-8. Simulated (left column) and observed (right column) changes in (upper row) 500 hPa geopotential height (in m) and (lower row) near-surface temperature (in K) and winds. Observed changes (Thompson

and Solomon, 2002) are 22-year linear trends in 500 hPa geopotential height and 925 hPa winds (1979 to

2000), and 32-year linear trends in surface temperature (1969 to 2000) averaged over December to May.

Simulated changes are differences between the perturbed and control integrations in 500 hPa geopotential

height, 950 hPa winds, and land surface temperature averaged over December to February. The longest

wind vector corresponds to ~4 m/s. From Gillett and Thompson, 2003.
Shindell and Schmidt, 2004) (see also Figure 5-8).

Simulated and observed wind changes of ~1 m/s extend to

30°S (Gillett and Thompson, 2003). 

Stratospheric ozone depletion likely influences tropospheric climate through both radiative and dynamical

processes. Idealized model simulations indicate that a perturbation to the diabatic heating in the stratosphere gives

an annular-mode response in sea level pressure over

intraseasonal time scales (Polvani and Kushner, 2002;

Kushner and Polvani, 2004; see also Section 5.2.2.6).

Recent results suggest that the tropospheric response to

stratospheric perturbations may result directly from wave

driving and radiative forcing changes in the stratosphere

(Thompson et al., 2006; see Section 5.2.2.6), although this

direct forcing effect cannot explain the latitudinal structure

of the tropospheric response. Antarctic stratospheric ozone

depletion acts as a direct radiative cooling influence at the

surface. Indeed, an early study of the response to stratospheric ozone depletion with a radiative-convective model

found a surface cooling over Antarctica in response to stratospheric ozone depletion (Lal et al., 1987). 

Maximum stratospheric ozone depletion close to the

tropopause occurs in December to January, more than a

month after the maximum ozone depletion at 70 hPa, due

to the downward transport of ozone-depleted air (Solomon

et al., 2005). Since surface temperature is particularly sensitive to changes in ozone concentration close to the

tropopause (Forster and Shine, 1997), this suggests that

surface cooling is radiatively induced, and that the apparent

lag between stratospheric and tropospheric responses is

due to the downward transport of ozone-depleted air toward

the tropopause, rather than any dynamical effect. 

Most studies suggest that Antarctic ozone depletion

is likely to peak sometime in the current decade, and that

a recovery is likely to follow over the next 50 years

(Chapter 6). Therefore, over the coming decades,

increases in stratospheric ozone should drive a decrease in

the SAM index toward values seen before ozone depletion. However, increasing GHGs will likely have the

opposite effect, contributing an increase in the SAM index

(Shindell and Schmidt, 2004; Arblaster and Meehl, 2006).

The magnitudes of the future ozone and GHG effects on

the SAM are therefore uncertain.

5.2.7.2 NORTHERN HEMISPHERE

Stratospheric ozone depletion has also occurred in

the Northern Hemisphere, where it has led to a smaller

cooling of the Arctic polar vortex, maximizing in the spring,

although model studies indicate that only part of the cooling

observed in the Northern Hemisphere stratosphere can be

explained by ozone depletion (Section 3.4.3.1 in WMO,

2003). Some studies have suggested that ozone depletion

has contributed to the observed trend toward the positive

phase of the NAO or NAM (Graf et al., 1998; Volodin and

Galin, 1999; Hartmann et al., 2000). Ozone was found to

contribute a small trend toward the positive phase of the

NAM in model simulations, albeit one rather smaller than

that due to GHGs, or indeed that observed (Graf et al., 1998;

Shindell et al., 2001). However, while the largest response

to stratospheric ozone depletion is simulated in the spring

(Graf et al., 1998), the largest trends have been observed in

the winter months (Thompson et al., 2000), when no significant trends are simulated in response to ozone depletion

(Gillett et al., 2003; Gillett, 2005). Overall, therefore, based

on theoretical considerations and some modeling studies, it

can be concluded that Arctic ozone depletion has likely contributed to the weak positive trend in the NAM in the spring,

but that it cannot explain the observed winter trends.

5.3 EFFECTS OF ANTHROPOGENIC CLIMATE

CHANGE AND OF EMISSIONS ON STRATOSPHERIC OZONE

5.3.1 Stratospheric Temperature Changes

Due to Shifts in Radiative Forcing

Long-term changes in radiative forcing over the

next few decades will continue to impact global mean temperatures in both the troposphere and stratosphere. Over

the past three decades, increases in well-mixed greenhouse

gas (WMGHG) concentrations and declines in stratospheric ozone have been the primary forcing mechanisms

affecting stratospheric climate.

Global concentrations of WMGHGs will continue

to rise in the next half century, although significant uncertainties remain as to the exact rate of increase. These

changes will act, on average, to cool the stratosphere (see

Section 5.2.1), but there can still be a seasonal warming,

particularly at high latitudes, due to changes in planetary

wave activity. Therefore, the assessment of the future evolution of polar temperatures is uncertain. Future changes

in stratospheric water vapor are also difficult to predict, in

part because the changes observed over the last four

decades are still not fully understood (e.g., Randel et al.,

2004b, 2006; see also Section 5.2.5). While declines in

stratospheric ozone also act to cool the stratosphere, within

this decade global ozone levels will likely begin to rise

(Chapter 6). Higher ozone levels will increase stratospheric ozone heating, which will at least partially offset

the cooling due to increases in WMGHG concentrations.

Because ozone concentrations are so sensitive to the background temperature field, understanding the complex

CLIMATE-OZONE CONNECTIONS

5.18
interaction between changing constituent concentrations

and temperature requires an evaluation of the coupling

between chemistry, radiation, and atmospheric dynamics. 

The primary method used for quantifying the interaction between changing chemistry and dynamics is the

Chemistry-Climate Model (CCM). With advances in

CCMs, the nonlinear coupled evolution of chemistry and

dynamics can be studied. One of the continuing challenges

for climate models (AOGCMs, coupled ocean-atmosphere

general circulation models) is predicting future changes in

tropospheric and stratospheric wave dynamics, which in

turn affect the structure and evolution of lower and upper

atmospheric temperature (see Section 5.2.2). Improvements in CCMs (Eyring et al., 2006; see also Section 6.6)

are providing valuable insight into how temperatures will

evolve in the future.

Model simulations for doubled CO2 conditions

with and without interactive ozone chemistry show a substantial temperature decrease throughout most of the

middle atmosphere, with a maximum cooling of 10-12 K

near the stratopause (Jonsson et al., 2004; Sigmond et

al., 2004; Fomichev et al., 2006; see Figure 5-13a). In

these calculations, stratospheric ozone increases are associated with lower temperatures, which partially modulate the CO2-induced cooling through enhanced ozone

heating. While simulations using CCMs demonstrate

how the inclusion of interactive chemistry alters model

meteorology (Austin and Butchart, 2003; Manzini et al.,

2003; Tian and Chipperfield, 2005), more fundamental

dynamical processes such as the parameterization of

gravity waves and propagation characteristics of planetary waves remain significant modeling challenges (see

also Section 5.2.2.4). 

AOGCMs (i.e., AGCM coupled with an ocean model)

whose output will be used in the Intergovernmental Panel

on Climate Change (IPCC) Fourth Assessment Report (AR4;

IPCC, in preparation for 2007) represent the most advanced

and comprehensive set of climate simulations so far produced. To better represent the many physical processes

responsible for climate variability and change, enhanced

horizontal and vertical resolution has been used in these

models compared with earlier IPCC Assessments. While

the AOGCMs used for the IPCC were not specifically

designed for stratospheric simulations, it is becoming

increasingly apparent that accurate simulations of the stratosphere are important in determining the evolution of the surface climate, as well as other aspects of climate change. It is

for this reason that the assumptions that have been chosen in

the IPCC AOGCM simulations for the 21st century (i.e.,

future greenhouse gas (GHG) concentrations are only specified by different scenarios from the Special Report on

Emissions Scenarios (IPCC, 2000) that differ primarily in

the emissions of WMGHGs) constitute a restriction for an

assessment of climate change within the next few decades.

In the high (A2) scenario, concentrations of CO2 increase

from today’s value (~380 ppmv) to approximately 850 ppmv

by 2100, while the low (B1) scenario reaches 550 ppmv by

2100 (IPCC, 2001). In these AOGCM simulations, ozone

concentrations are prescribed and not interactively computed

during the 21st century, which is an obvious restriction and

has an impact on the assessment of the future development

of lower stratospheric temperature (see below).

Figure 5-9 shows a time series of global-mean temperature anomalies between 2000 and 2100 at 50 hPa from

fifteen AOGCMs adopting the A2 emission scenario used

for the IPCC AR4. See Appendix 5A for information concerning the specific AOGCMs. While all simulations

show 50 hPa global temperatures declining over the 21st

century, the range of predictions for temperature change

varies from –0.5 to –3.5 K by 2100. Differences in model

ozone concentrations over the 21st century are likely

responsible for some of the range in model predictions,

while variations in model dynamics, the limited vertical

extent of many of the AOGCMs, and differences in radiation schemes are also likely contributors. 

The relative uncertainty in model predictions of

future stratospheric temperature is further illustrated in

Figure 5-10, which shows the globally averaged temperature trend computed during the 21st century using the

models submitted to the IPCC AR4 for low (B1) and high

(A2) emission scenarios. In the troposphere, temperatures rise during the next century, ranging between 0.2

K/decade (low) to 0.5 K/decade (high) for the different

emission scenarios. In the stratosphere, the rate of cooling

is also strongly dependent on the emission scenario,

ranging from 0.07 ± 0.20 K/decade (low) to 0.18 ± 0.20

K/decade (high) at 50 hPa and 0.38 ± 0.09 K/decade (low)

to 0.72 ± 0.47 K/decade (high) at 10 hPa. However, it

should be reemphasized that in the present set of IPCC

AOGCM simulations, the ozone forcing in the 21st century varies from constant ozone to a slow recovery by

2050, and thus contributes to the large model-to-model

variability (Hare et al., 2004). 

A similar analysis of global-mean temperature

trends from CCMs (Eyring et al., 2006) that consider

interactive ozone feedback (see Box 5-1) generally shows

less model-to-model variability compared with the

AOGCM results submitted to the IPCC AR4. See

Chapter 6, Table 6-4 for information concerning the specific CCMs. Figure 5-11 (bottom panel) shows that

global temperature trends at 50 hPa from CCMs are in

reasonable agreement with ERA-40 (European Centre

for Medium-Range Weather Forecasts (ECMWF) 40-

year Reanalysis) and radiosonde data (Radiosonde

CLIMATE-OZONE CONNECTIONS

5.19
CLIMATE-OZONE CONNECTIONS

5.20

2000 2020 2040 2060 2080 2100

Year

-4

-2

0

2

4

Temperature (K)

miroc3.2-m: -0.037 ipsl-cm4: -0.086 cgcm3.1-t47: -0.068 inm-cm3.0: -0.11 echam5-mpi-om: -0.12 pcm: -0.12 gfdl-cm2.0: -0.16 ukmo-hadgem1: -0.18 bccr-bcm2.0: -0.13 mri-cgcm2.3.2a: -0.19 csiro-mk3.0: -0.18 gfdl-cm2.1: -0.19 giss-er: -0.22 ccsm3: -0.30 cnrm-cm3: -0.46 

Figure 5-9. Evolution of global-mean temperature anomaly at 50 hPa between 2000 and 2100 as estimated

from AOGCMs for the A2 emission scenario. The temperature trend in K/decade is given next to the name of

each participating model. Anomalies are computed with respect to temperatures between 2010 and 2020. For

clarity, models are labeled in order from weakest to strongest temperature variation. Information about the

AOGCMs is given in Appendix 5A.

-1.0 -0.5 0.0 0.5 1.0

Trend (K/decade)

1000

100

10

Pressure (hPa)

A2

B1

Figure 5-10. Global and annual

mean 21st century temperature trends

from AOGCMs using A2 (high) and B1

(low) emission scenarios. The boxes

indicate the average trend computed

for all models, while the thin horizontal

lines indicate the range of modelcalculated trends. Information about

the AOGCMs is given in Appendix 5A.
CLIMATE-OZONE CONNECTIONS

5.21

 

-10

-5

0

5

10

15

Anomaly (K)

AMTRAC: -0.095

CCSRNIES 60°N-90°N (Feb-Apr) : -0.99

CMAM: -0.29

E39C: -0.30

G OSCCM: -2.1

MDZrepro: -3.0 

MAECHAM4CHEM: 1.5

MRI: -3.5

SOCOL: 1.0

ULAQ: -0.13

UMETRAC: -1.

UMSLIMCAT: -1.0

WACCM: -0.12

ra-40: - 2.1 

 

-10

-5

0

5

10

15

Anomaly (K)

AMTRAC: -3.6

60°S-90°S (Sep-Nov) CCSRNIES: -2.0

CMAM: -0.97

E39C: -1.9 

GEOSCCM: -1.6

MDZrepro: -3.9 

MAECHAM4CHEM: -4.1 

MRI: -2.8 

SOCOL: -2.2

ULAQ: -3.2 

UMETRAC: -3.5

UMSLIMCAT: -3.1

WACCM: -2.8

era-40: - 2.4

1960 1970 1980 1990 2000

Year

-2

0

2

4

Anomaly (K)

AMTRAC: -0.88 

CCSRNIES: -0.70 Global (Annual) CMAM: -0.22

E39C: -0.76

GEOSCCM: -0.4

MDZrepro: -0.40 

MAECHAM4CHEM: -0.87

MRI: -0.64

SOCOL: -0.49

ULAQ: -0.4

UMETRAC: -0.99

UMSLIMCAT: -0.29 

WACCM: -0.43 

era-40: - 0.77

RATPAC: Global: 50hPa-1.

Figure 5-11. Time series of temperature anomalies for Arctic spring (top), Antarctic spring (middle), and the

annual global mean (bottom) at 50 hPa derived from the REF1 (20th century climate) CCM simulations and

from observations (i.e., ECMWF 40-year reanalysis, ERA-40, and Radiosonde Atmospheric Temperature

Products for Assessing Climate, RATPAC). The temperature anomalies are calculated with respect to a mean

reference period between 1980 and 1989. A linear temperature trend in K/decade is calculated for each model

using data between 1980 and 1999. AMTRAC, MAECHAM4CHEM, MRI, UMETRAC, SOCOL, ULAQ, and

E39C are shown with dashed lines, all other CCMs with solid lines. The temperature trend is given next to the

name of each participating model (from Eyring et al., 2006). See Table 6-4 and Appendix 6A of Chapter 6 for

information about the CCMs and the model runs.
Atmospheric Temperature Products for Assessing

Climate, RATPAC; Free et al., 2005) and display an

obvious cooling from 1980 to present day. The model

temperature trends range from –0.22 K/decade to –0.99

K/decade (average = –0.64 K/decade), with 7 out of 13

CCMs showing statistically significant trends, compared

with the ERA-40 (–0.77 K/decade; not significant) and

RATPAC (–1.0 K/decade; significant) data. Corresponding AOGCM results show model temperature trends

ranging from –0.03 to –0.97 (average = –0.38 K/decade).

Perturbations by volcanic eruptions are well captured in

many of the CCMs, but generally the temperature

response is over-predicted (see Section 5.2.4; Figure 5-

3). A similar over-prediction of the warming effect at 50

hPa from volcanic aerosols was also found with the

AOGCMs. The evolution of lower stratospheric (50 hPa)

springtime temperatures in the polar regions (Figure 5-

11, top and middle panel) indicates a much greater interannual variability, which is caused by the specific impact

of wave dynamics (Section 5.2.2.1; see also Chapter 4,

Section 4.1.1.1). This is another reason for larger uncertainties in the assessment of the future evolution of stratospheric springtime temperatures in these geographical

regions. 

Figure 5-12 shows that there is good agreement

between the CCM results showing a global cooling trend

at 50 hPa through the first half of the 21st century. The

CLIMATE-OZONE CONNECTIONS

2000 2020 2040 2060 2080 2100

Year

-4

-2

0

2

4

Anomaly (K)

AMTRAC: -0.15 

CCSRNIES: -0.29 

CMAM: -0.30 

E39C: -0.23 

GEOSCCM: -0.14 

MAECHAM4CHEM: -0.26 

MRI: -0.16 

SOCOL: -0.26 

ULAQ: -0.29 

UMSLIMCAT: -0.17 

WACCM: -0.26 

Figure 5-12. Time series of global-mean annual temperature anomalies at 50 hPa from the REF2 (21st century climate, A1B scenario; IPCC, 2001) CCM simulations (see Chapter 6, Appendix 6A). The temperature

anomalies are calculated with respect to a mean reference period between 2000 and 2010. AMTRAC, E39C,

MAECHAM4CHEM, MRI, SOCOL, and ULAQ are shown with dashed lines, all others CCMs with solid lines. A

linear temperature trend in K/decade is calculated for each model using data after 2000 and is shown next to

the name of each participating model. See Chapter 6, Table 6-4 and Appendix 6A for information about the

CCMs and the model runs.

Table 5-1. Mean linear temperature trend in

K/decade calculated from different types of models

and for different emission scenarios using model

data after year 2000. See text for detailed description.

Emission Scenario (Model Employed)

Pressure B1 A1B A2

Level (AOGCM) (CCM) (AOGCM)

–0.07 –0.23 –0.18

±0.20 ±0.09 ±0.20

–0.38 –0.63 –0.72

±0.09 ±0.23 ±0.47

5.22

50 hPa

10 hPa
CCMs adopted a medium (A1B) emission scenario used

for the IPCC AR4, which reaches 720 ppmv by 2100.

Considering the CCM results up to the year 2050, the

cooling trend ranges from 0.14 to 0.30 K/decade. In the

lower stratosphere, the temperature trends derived from

the CCM simulations using the medium (A1B) emission

scenario are larger compared with the AOGCM results

which are related to the high (A2) emission scenario (cp.

Table 5-1). At 10 hPa, where ozone changes are expected

to be weak, AOGCM and CCM simulations are in line,

i.e., the calculated trends derived from the AOGCM A2

simulations are stronger than those derived from the CCM

A1B simulations (see Table 5-1). The fact that the different CCM results are more tightly constrained may be

due to the smaller range in the prescribed forcing (see

Chapter 6, Appendix 6A). Both results from AOGCMs

and CCMs point to the changing nature of the stratosphere

and troposphere and the important role that emission of

WMGHGs has in the prediction of future temperature in

the stratosphere. 

5.3.2 Stratospheric Temperature Changes

Due to Additional Effects

Future changes in radiative forcing will have both a

direct in-situ effect on stratospheric temperatures (Section

5.3.1) and an indirect effect resulting from changes in the

wave forcing from the troposphere. For the stratosphere,

to a first approximation, the effects of in-situ radiative

changes and those of changes in wave forcing from the

troposphere are additive and therefore can be assessed separately (Sigmond et al., 2004; Fomichev et al., 2006).

Figure 5-13 illustrates the response of the middle atmosphere (10-80 km) to a doubling of CO2 in the troposphere

and middle atmosphere separately, as well as together

(Sigmond et al., 2004). Anthropogenic climate change is

likely to have an impact on sea surface temperatures

(SSTs), land-sea temperature differences, tropospheric

jets, and synoptic scale activity, which in turn are likely to

have an impact on the planetary wave forcing originating

in the troposphere, as well as their regional effects in the

lower stratosphere. Climate change can also affect the

propagation of planetary waves into the stratosphere (Rind

et al., 2005a; 2005b; Scott and Polvani, 2004; Scott et al.,

2004). Changes in planetary wave forcing and propagation in the polar winter are a major source of uncertainty

for predicting future levels of Arctic ozone loss (Austin et

al., 2003). CCM (see Box 5-1) results used for WMO

(2003) did not agree on the sign of the trend in the planetary wave flux from the troposphere. Recently, Fomichev

et al. (2006) found that the response in the polar winter

stratosphere to a doubling of CO2 in the first 15 years of a

30-year simulation differed from the response in the last

15 years, which also supports earlier findings (e.g.,

CLIMATE-OZONE CONNECTIONS

a) b) c )

Figure 5-13. Changes in the middle atmosphere resulting from different doubling CO2 model runs for the months

of December-January-February: (a) The change in the zonally averaged temperature due to a doubling of CO2 in

the middle atmosphere, ∆TM ; (b) the change due to a doubling of CO2 in the troposphere, ∆TT ; and (c) the degree

of non-additivity, i.e., the sum of the temperature changes due to a doubling of CO2 separately in middle atmosphere and troposphere, minus the temperature change resulting from a doubling of CO2 in all the atmosphere

(∆TM + ∆TT –∆TA). Light (dark) shading denotes significance at 95% (99%) level of the changes in (a) and (b) and

of the non-additivity in (c). The contour interval is 1 K. The thick dashed line denotes the position of the tropopause.

From Sigmond et al., 2004.

5.23
Butchart et al., 2000) that on a decadal time scale, the

dynamics of the polar winter stratosphere are unpredictable due to internal variability.

Since WMO (2003), a number of transient CCM

simulations have been run into the 21st century using the

same climate forcing and halogen loadings (see Chapter

6, Table 6-4 and Appendix 6A). Eyring et al. (2006)

assessed the CCMs used for these simulations and found

that temperature biases in the northern winter in the polar

lower stratosphere were quite small, and most of the

models exhibited the correct temperature sensitivity to

variations in the wave forcing from the troposphere. These

are both notable improvements relative to the CCMs used

by Austin et al. (2003). In contrast, there was little

improvement in the Southern Hemisphere, with most of

the CCMs still predicting Antarctic cold biases in spring,

and the polar vortex breaking down much later than

observed. Again, these latest CCM simulations do not

provide any consensus as to what will happen to the planetary wave forcing from the troposphere in a future climate (see Figure 5-14), despite each model using the same

amounts of greenhouse gases and halogens. 

The reasons why the models continue to fail to provide a consensus are unclear. In part it could be due to

model deficiencies, especially in the SH where many of

the models are unable to reproduce the correct polar temperature sensitivity to the wave forcing from the troposphere. On the other hand, as argued in WMO (2003,

Chapter 3), it could reflect true atmospheric behavior: in a

truly chaotic system, planetary wave forcing from the troposphere may not be inherently predictable, and so models

are giving a range of results reflecting that unpredictability.

Clarification is strongly required because the assessment

of the evolution of the ozone layer depends on the ability

of CCMs to predict wave changes. 

The most important SST factor affecting the troposphere and the stratosphere is tropical SST gradients, as

found during El Niño-Southern Oscillation (ENSO)

events. Anomalous tropical SST gradients are also known

to have an impact on planetary wave forcing. The stratoCLIMATE-OZONE CONNECTIONS

5.24

Northern Hemisphere

1960 1980 2000 2020 2040

Year

10

15

20

Heat flux

Souther Hemisphere

1960 1980 2000 2020 2040

Year

−12

−8

−4

Heat flux

AMTRAC

CCSRNIES

CMAM

E39C

GEOSCCM

MAECHAM4chem

MRI

SOCOL −5

ULAQ

UMSLIMCAT

WACCM

Figure 5-14. Evolution of the heat flux (K m/s) at 100 hPa as derived from CCM calculations. The analysis is

based on model data derived from the REF2 and SCN2 simulations (see Chapter 6, Appendix 6A). Left panel:

Heat flux averaged over 40°N to 80°N in January and February. Right panel: Heat flux averaged over 40°S to

80°S in July and August. Results are shown for each year of the simulations but have been smoothed with a 5-

year running mean. The straight lines represent the least squares linear fit for the individual model runs. The

heat flux at 100 hPa provides a measure of the wave flux propagating from the troposphere to the stratosphere. A negative heat flux in the SH corresponds to an upward planetary wave flux.
spheric polar vortex tends to be weaker than average

during warm ENSO years, when SSTs are anomalously

warm in the eastern tropical Pacific (van Loon and

Labitzke, 1987), while cold ENSO years are associated

with suppressed incidences of stratospheric sudden warmings (Limpasuvan et al., 2005). Similar relationships have

been noted in several numerical simulations (Hamilton,

1993; Sassi et al., 2004; Manzini et al., 2006). Moreover,

there are indications of increased incidences of sudden

stratospheric warmings in CCM simulations run with high

SST anomalies in the eastern tropical Pacific (Taguchi and

Hartmann, 2006; see Figure 5-15). Hoerling et al. (2001;

2004) and Hurrell et al. (2004) argued that changes in

Indian Ocean SSTs have a demonstrable impact on the

Northern Hemisphere Annular Mode/North Atlantic

Oscillation (NAO/NAM) on decadal time scales, but the

extent to which these changes extend to stratospheric

levels remains unclear.

Outside the polar regions in winter, changes in the

tropospheric forcing are also expected to have an impact

on lower stratospheric temperatures throughout the year.

There are indications from model assessments that anthropogenic climate change will increase the upwelling across

the tropical tropopause and downwelling in the extratropics (see Sections 5.3.3 and 5.3.7). The associated adiabatic heating will then lead to a cooling of the tropical

lower stratosphere and warmer extratropics in all months

(Figure 5-16), but it is more pronounced during the boreal

winter and spring, which is consistent with the asymmetry

in the SST change between the NH and SH in the adopted

model (Fomichev et al., 2006). Other climate change simulations (Butchart et al., 2006) also indicate that, on

average, the increase in planetary wave driving is greater

during the northern winter than during the southern winter.

All CCMs and AOGCMs used to predict stratospheric temperature changes (see Section 5.3.1) parameterize the effects of small-scale gravity waves. The more

sophisticated of these schemes parameterize the vertical

propagation and breaking of the waves in terms of the

large-scale flow, and therefore are able to respond to climate perturbations in the stratosphere resulting from radiative or planetary wave forcing. Moreover, research by

Shaw and Shepherd (2007) suggests that, provided the

parameterizations are implemented to conserve momentum, the climate response will be robust to differences in

the gravity wave source spectrum, background flow,

gravity wave-breaking criterion, and model upper

boundary. McLandress and Scinocca (2005) also showed

that the response is not very dependent on the choice of

gravity wave drag parameterization, at least among those

schemes that have been developed to represent the propagation and breaking. It is not clear what proportion of the

change in gravity wave forcing will result from the indirect effect of climate change modifying the propagation

and breaking of the waves, and how much will result

directly from a change in the source. The greatest uncertainty is in knowing how the source spectrum will evolve

in a changing climate. Unfortunately, the spectrum is

rather poorly constrained by observations for the present

climate, and it will be difficult to quantify the effects of

CLIMATE-OZONE CONNECTIONS

5.25

Figure 5-15. Climatological differences between

WACCM (Whole Atmosphere Community Climate

Model) simulations forced with high and low SSTs in

the eastern tropical Pacific: (a) zonal mean temperature [T] in K and (b) zonal mean zonal wind [U] in

m/s. Both panels display results for the Northern

Hemisphere, with the equator on the left and the

North Pole on the right. From Taguchi and

Hartmann, 2006. 

P r e

s

s

u r e ( h

P

a ) P r e

s

s

u r e ( h

P

a )

0 30 60 90

Latitude
changes in gravity wave forcing on the stratospheric temperatures without further progress toward specifying

sources in terms of other flow parameters to allow for climate feedbacks.

5.3.3 Troposphere-Stratosphere Exchange

and the Brewer-Dobson Circulation

The net mass exchange between the troposphere

and stratosphere is associated with the large-scale BrewerDobson circulation (Holton et al., 1995; Shepherd, 2002),

with a net upward flux in the tropics balanced by a net

downward flux in the extratropics. However, near the

tropopause itself, the picture is more complex, with twoway mixing across the extratropical tropopause at and

below synoptic scales, and vertical mixing in the tropicaltropopause layer (TTL) resulting from convective

processes (see Section 5.3.4). Nonetheless, above the

lowermost extratropical stratosphere and at the top of the

TTL, the exchange is more one-way with, in particular,

air slowly rising into the stratosphere above the TTL.

Model studies indicate that climate change will impact

the mass exchange across the tropopause. Rind et al.

(2001) estimated a 30% increase in the mass flux due to a

doubling of atmospheric CO2 amounts, and Butchart and

Scaife (2001) estimated that the net upward mass flux

above the TTL would increase by about 3% per decade

due to climate change. In both of these studies, the

changes in the mass flux resulted from more vigorous

wave propagation from the troposphere into the stratosphere. Modeling studies of tropospheric ozone (Collins

et al., 2003; Zeng and Pyle, 2003; Sudo et al., 2003) also

found that climate change caused a comparable percentage increase in the extratropical stratosphere-totroposphere ozone flux. 

For a doubled CO2 concentration, all 14 climatechange model simulations in Butchart et al. (2006)

resulted in an increase in the annual mean troposphere-tostratosphere mass exchange rate (Figure 5-17), with a

mean trend of 11 Gg s-1 year-1, or about 2% per decade.

The predicted increase occurred throughout the year but

was, on average, larger during the boreal winter than

during the austral winter. The Butchart et al. study was

unable to conclude whether stratospheric ozone changes

or ozone feedbacks had a significant impact on the underlying trend in the mass exchange rate. Other simulations

CLIMATE-OZONE CONNECTIONS

5.26

 Tropical and Extratropical Temperature Cycle 

Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec

Month

-10

-5

0

5

10

Temperature (K)

Figure 5-16. Annual cycle of the tropical (30°S-30°N; negative curves) and extratropical (positive curves)

mean temperature, with global mean subtracted, at 50 hPa. The dashed curve results from a model simulation in which the atmospheric CO2 was doubled, but the SSTs were not adjusted for the doubled CO2 conditions. The solid curve is from a simulation with both the CO2 doubling and adjusted SSTs, and therefore

includes an additional effect on the temperatures due to changes in wave forcing from the troposphere. From

Fomichev et al., 2006.
(e.g., Austin et al., 2006) suggest that the trend in tropical

upwelling is not constant. Periods of enhanced upwelling

coincide with periods of significant ozone depletion. 

The calculation of tropical upwelling (or of mean

age of air1; see below) from measurements or assimilated

datasets is challenging, and it is not possible to calculate

reliable trends. Calculations of age of air using a transport model and meteorological fields (Schoeberl et al.,

2003; Meijer et al., 2004; Scheele et al., 2005) gave values

that were lower than those determined using observations

of long-lived tracers (e.g., Boering et al., 1996; Andrews

et al., 2001; Schoeberl et al., 2005). Similar results have

been obtained recently from a suite of CCMs (Eyring et

al., 2006; see Chapter 6). 

Tropical upwelling is inversely related to model age

of air (Austin and Li, 2006), so that the age of air changes

as the stratospheric climate changes. As noted for the tropical upwelling, the age of air does not change steadily

(Figures 5-18 and 5-19). In both CCMs AMTRAC (e.g.,

Austin et al., 2006) and WACCM (e.g., Beres et al., 2005),

age of air decreased significantly from about 1975 to 2000,

consistent with an increase in tropical upwelling, which in

AMTRAC was identified as due in part to ozone depletion (Austin et al., 2006). As shown in Figure 5-19,

WACCM also indicates that age of air remains constant

for conditions of fixed GHG concentrations and sea surface temperatures (SSTs) (see also Section 5.3.7), indicating that changes in GHG concentrations and SSTs are

major influences for age of air changes in the future. The

overall decrease of age of air and increase in tropical

upwelling on climate time scales implies a more rapid

removal of the long-lived CFCs from the entire atmosphere (Butchart and Scaife, 2001), as well as source

species such as CH4 and N2O. Once enhanced CH4 concentrations reached the stratosphere, enhanced CH4 oxidation would occur, leading to a faster increase in water

CLIMATE-OZONE CONNECTIONS

5.27

1 Age of air: The length of time that a stratospheric air mass has been out of contact with the well-mixed troposphere. The content of

a unit element of air at a particular location and particular time of year in the stratosphere can be thought of as a mixture of different

air parcels that have taken different routes from the tropopause to arrive at that location. The mean age of air at a specific location is

defined as the average transit times of the elements since their last contact with the tropopause.

Annual

 

1.0 kt s−1 ear−1

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2.1 FUB−CMAM

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8.2 MAECHAM4chem

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 17.4 UM49L

7.8 M64L

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4.7 WACCM

a)

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14.9 kt s−1 ear−1

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4.6 M64L

9.6 M64Lchem

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5.6 WACCM

b)

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9.5 kt s−1 ear−1

6. MAM

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6.5 ISS

5.9 ISSchem

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8.4 MAECHAM4chem

4.1 MRI

2.9 M49L

4.3 M64L

7.9 M64Lchem

 11.4 M64LS

5.6 WACCM

c)

Figure 5-17. Schematic representation of the trends in troposphere-to-stratosphere mass exchange rate computed in a set of 14 CCMs and climate models that consider the complete stratosphere (from Butchart et al.,

2006). The trends (Gg s-1 year-1) are given by the slope of the lines. For the transient simulations, trends were

calculated from a least squares fit to the mass flux data. For the time-slice simulations (see Box 5-1) the difference in flux was converted to a mean mass flux trend, assuming a doubling time of 70 years for CO2. The

dashed line is the multi-model mean; the standard error is given by the gray shading.
CLIMATE-OZONE CONNECTIONS

5.28

Age of Air 20°N-20°S, 1.3 hPa

1950 2000 2050 2100

Date (year)

5.0

4.5

4.0

3.5

3.0

Age (years)

Figure 5-18. Mean age

of air in the tropical

upper stratosphere for

the period 1960-2100

computed from the CCM

AMTRAC. The three different colored curves

indicate individual simulations. A piecewise

linear curve (black

broken line) is included to

illustrate changes in

trends. Modified from

Austin and Li, 2006.

Age of Air 22°N-22°S, 1.24 hPa

1960 1980 2000 2020 2040 2060

Year

3.0

3.5

4.0

4.5

Mean Age (years)

REF1 (3)

REF2 (3)

NCC (3)

Figure 5-19. Mean age

of air for the CCM

WACCM (Whole Atmosphere Community Climate Model) in the tropical upper stratosphere.

The individual simulations are described in

Chapter 6 (Appendix 6A)

and cover the recent

past (REF1), the future

(REF2), and an ensemble of simulations

with fixed GHG concentrations (“no climate

change” scenario (NCC);

see also Section 5.3.7).
vapor amounts than would be anticipated on the basis of

the tropospheric CH4 concentrations alone. 

5.3.4 Changes in the Tropical and

Extratropical Tropopause Layer 

5.3.4.1 THE TROPICAL TROPOPAUSE LAYER (TTL)

The TTL (Sherwood and Dessler, 2001) is typically

defined as the body of air extending from the level of the

temperature lapse rate minimum at 11-13 km (Gettelman

and Forster, 2002) to the level of highest convective overshoot, slightly above the cold point tropopause (CPT) at

16-17 km (see Chapter 2, Section 2.4.1 for other definitions). This encompasses the level of zero net radiative

heating (z0) that marks the transition from radiative cooling

to radiative heating, divides the TTL into the lower and

upper TTL (Section 2.4.1), and depends on the presence

of clouds (Corti et al., 2005). Below z0, the cooling air

sinks back into the troposphere, whereas above z0 the

warming air rises and eventually enters the stratosphere.

The chemical state of this air depends on its residence time

in the TTL (Folkins et al., 1999; Thuburn and Craig, 2002;

Bonazzola and Haynes, 2004; Fueglistaler et al., 2004).

In addition, the minimum temperature experienced by this

air is crucial for dehydration along the transport pathway

in the TTL, and thus for stratospheric humidity (see

Sections 5.2.5 and 5.3.5). 

The TTL has undergone changes within the last few

decades that are not well known or understood, and our

predictive capabilities remain extremely limited. There is

no merged reference dataset of long-term global temperature observations for this height region. 

The TTL is sandwiched between a warming troposphere and a cooling stratosphere, which makes it difficult

to produce a theoretical estimate of the response of the

CPT and stratospheric humidity. A simple conceptual picture is shown in Figure 5-20. If one assumes a convectively controlled troposphere with a constant lapse rate,

then tropospheric warming raises and warms the

tropopause (cold point temperature increasing from T1 to

T2). A cooling of the stratosphere further raises the

tropopause but leads to a cooling (from T2 to T3). It seems

likely that the tropospheric warming should be the dominant effect, because the enhanced infrared cooling by

enhanced greenhouse gas concentrations is weak due to

the very low temperatures close to the tropopause (e.g.,

Clough and Iacono, 1995). The observed temperature

trends just above the tropical tropopause correspond to a

cooling of less than 0.4 K/decade but are not statistically

significant (Chapter 4 in WMO, 2003). The general concept behind Figure 5-20 has been corroborated by AGCM

time-slice (Shepherd, 2002) and transient simulations

(Santer et al., 2003) (see Box 5-1). 

The evolution of TTL temperatures is further complicated by the fact that there is a tropospheric amplification of surface warming (Santer et al., 2005). Conversely,

a strengthening of the Brewer-Dobson circulation, as suggested by model climate studies with enhanced CO2

(Butchart and Scaife, 2001; Rind et al., 2002a, b; Sigmond

et al., 2004; Eichelberger and Hartmann, 2005), would

imply a lowering of TTL temperatures. 

Seidel et al. (2001) obtained an increase in CPT

height of about 40 meters and a decrease in pressure of

about 1 hPa during 1978-1997. Both Seidel et al. (2001)

and Zhou et al. (2001) have also noted a decrease of tropical tropopause temperatures by about 1 K during this

period, resulting in a decrease in the saturation volume

mixing ratio of water of about 0.5 ppmv during 1978-1997.

These temperatures are therefore at odds with tropospheric

warming dominating the response of the CPT. Rather, they

CLIMATE-OZONE CONNECTIONS

5.29

T3

T2

T

T T3 T2

Z

T

Figure 5-20. Simplified sketch of the sensitivity of the

tropopause temperature and height with regard to

changes of temperature in the stratosphere and troposphere. Solid lines: reference profile with cold point T1.

Long-dashed line with solid line: perturbed profile

reflecting tropospheric warming with cold point T2.

Long-dashed line with short-dashed line: perturbed profile reflecting tropospheric warming and stratospheric

cooling with cold point T3. After Shepherd, 2002.
suggest that the CPT is being largely controlled by

increases in the Brewer-Dobson circulation and by

increased convection as suggested by Zhou et al. (2001).

Zhou et al. (2004) have also shown how the QBO and El

Niño-Southern Oscillation effects produce extremely high

and low tropical CPT temperatures.

Figure 5-21a indicates a very small negative trend

of tropical temperatures on the 100 hPa level over the past

decade, which merges into an enhanced negative evolution during 2001-03. The reasons for the most recent

development are not clear at present. Figure 5-21b shows

stratospheric total water content. Large anomalies are centered in the tropics and are in phase with the 100 hPa tropical temperatures. The low 2001-03 water vapor anomaly

covers nearly the entire globe. The observed strong seasonal and interannual T-H2O correlations suggest stratospheric total water content to be strongly controlled by

temperatures between the 100 hPa level and the CPT,

which is in agreement with recent trajectory calculations

(Hatsushika and Yamazaki, 2003; Bonazzola and Haynes,

2004; Fueglistaler et al., 2005; Fueglistaler and Haynes,

2005). 

Changes in the TTL may also affect the abundance

of many other species in the stratosphere. This may concern short-lived chemical species, such as biogenic

bromine compounds that may be carried to the stratosphere

via deep convection followed by transport through the TTL

(see Chapter 2, Section 2.4.1). Changes in deep convection may further affect the transport of longer-lived species

produced by biomass burning, such as methyl bromide

(Andreae and Merlet, 2001). Finally, species may be transported in particulate form across the tropical tropopause,

e.g., organic sulfur-containing species (Notholt et al.,

2005). Little is known about these processes, and even

less is known about climate-induced changes.

In summary, given the uncertainties in our understanding of mechanisms in the TTL and of their previous

changes, predictions of future changes in TTL morphology, processes, or transport are rendered difficult. A

future atmosphere with increasing greenhouse gas loadCLIMATE-OZONE CONNECTIONS

5.30

b

92 04 94 96 98 00 02

60°N

30°N

0

30°S

60°S

Latitude

a 2

-2

-1

0

-3

Temperature (K)

Figure 5-21. (a) Time series of deseasonalized 100 hPa temperature anomalies over

10°N-10°S, showing results from six different datasets (red = radiosondes, thick

black = GPS, and various model results).

(b) Latitude-time cross sections at 82 hPa

of deseasonalized anomalies in H2O + 2 ✕

CH4 from HALOE measurements. Contours

are ± 0.1, 0.3, etc., ppmv. Dashed vertical

lines show the onsets of various cold

phases ascribed to the quasi-biennial oscillation and El Niño-Southern Oscillation.

From Randel et al., 2004a.
ings is expected to develop a warmer troposphere with

enhanced deep convection. But for the reasons mentioned

above, it remains speculative that this will be reflected in

a warmer tropopause, higher water mixing ratios, and as a

consequence, a moister stratosphere with less rapid

recovery of ozone. The past decades suggest the dominance of other processes, possibly related to changes in

the Brewer-Dobson circulation, which may give rise to

surprises.

5.3.4.2 THE EXTRATROPICAL TROPOPAUSE LAYER

(EXTL) 

As described in Section 2.4.2 of Chapter 2, the

ExTL is a layer of air adjacent to the local extratropical

(thermal) tropopause, which has been interpreted as the

result of irreversible mixing of tropospheric air into the

lowermost stratosphere (Hoor et al., 2004; Hegglin et al.,

2005) or as the result of two-way stratosphere-troposphere

exchanges (Pan et al., 2004; Bischoff et al., 2006). 

The origin of ozone in the ExTL changes markedly

with season, with photochemical production dominating

in summer and transport from the stratosphere dominating

in winter and spring. A general upward trend of the extratropical tropopause height has been identified and has been

related to ozone column changes (Steinbrecht et al., 1998;

Varotsos et al., 2004). This long-term change provides a

sensitive indicator of human effects on climate (Santer et

al., 2004) and may be related to changes such as are

sketched in Figure 5-20. However, it is presently not clear

how these changes have affected, or will affect, the dominant transport pathways of air between the troposphere

and the stratosphere on the synoptic scale or mesoscale. 

Possible future changes in midlatitude circulation—

increased strength of Brewer-Dobson circulation (e.g.,

Butchart and Scaife, 2001), changes in synoptic eddies

(e.g., Schneider, 2004), changes in midlatitude convection

(e.g., IPCC, 2001)—are all likely to change ozone concentration in the upper troposphere and lower stratosphere.

However, limitations of our knowledge regarding the

ExTL prevent any firm predictions on the future of the

ExTL. In particular, the relative contribution of isentropic

(quasi-horizontal) and convective (vertical) transport of

tropospheric air into the lowermost stratosphere is not well

established, though there is evidence that vertical mixing

(both convectively and radiatively driven) might have a

significant influence on the large-scale trace gas distributions in the lowermost stratosphere (Hegglin et al., 2005).

If the frequency or intensity of midlatitudinal deep convection were to change in a future warmer climate, this

could affect the lowermost stratosphere and, thus, the midlatitudinal ozone layer. Simulations of the 20th century in

the IPCC climate models (AOGCMs) generally show a

decrease in the total number of extratropical storms in both

hemispheres, but an increase in the number of the most

intense storms (Lambert and Fyfe, 2006). Predictions of

implications for ExTL dynamics and chemistry will

require further investigations.

5.3.5 Impact of Future Water Vapor

Changes on Ozone Chemistry 

It is difficult to assess the magnitude of recent

changes to stratospheric water vapor concentrations

(Section 5.2.5), and future changes are uncertain. Water

vapor concentrations may remain similar to current values,

or possibly increase. For example, most of the CCM future

simulations (i.e., REF2; see Chapter 6, Appendix 6A) show

increasing stratospheric water vapor concentrations in the

stratosphere. In this section, the implications for ozone of

possible future increases in water vapor concentrations

are discussed. 

A positive trend in stratospheric water vapor concentrations would affect stratospheric ozone production

and loss chemistry. Increases in water vapor would cause

increases in hydrogen oxide (HOx) radicals, affecting

ozone loss processes. HOx chemistry is the primary ozone

loss process in the lower stratosphere (Wennberg et al.,

1994; 1998), except over the polar caps. In addition, balloon constituent data, in conjunction with a photochemical box model, have highlighted the important role that

HOx plays in the stratosphere and lower mesosphere

(Osterman et al., 1997; Jucks et al., 1998). 

Evans et al. (1998) examined the effect of increased

water vapor concentrations on ozone; they found that

increased humidity slightly increased ozone in the middle

stratosphere, and decreased ozone in the upper stratosphere. The decrease in upper stratospheric ozone was

primarily from increased loss by the HOx catalytic cycles.

Dvortsov and Solomon (2001) modeled past and future

changes in greenhouse gas concentrations and atmospheric

chlorine and bromine loading. The model was forced to

have an annual stratospheric H2O increase of 1% per year.

The water vapor trend in the model intensified the

Northern Hemisphere midlatitude ozone trends, primarily

due to increased ozone loss in the lower stratosphere from

enhanced HOx, and subsequent HOx-amplified ClOx ozone

loss. The study suggests that a 1% per year stratospheric

humidity trend during this century would enhance ozone

loss via the HOx catalytic cycles and delay the ozone

recovery by approximately 10 years. Shindell (2001),

using a CCM, also showed that including a similar future

trend in water vapor may delay the ozone recovery time

by approximately 15 years. 

CLIMATE-OZONE CONNECTIONS

5.31
The CCM study of Stenke and Grewe (2005) found

that an increase of 1 ppmv in stratospheric water vapor

would result in a 5-10% hydroxyl radical (OH) increase in

the tropical lower stratosphere between 100 and 30 hPa

(Figure 5-22, top panel). In the model, the OH increase

caused an increase in the HOx ozone destruction cycle of

about 6%. Ozone in the lower stratosphere decreased by

1-3% (Figure 5-22, bottom panel), reducing the column

by less than 1% in nonpolar latitudes (Figure 5-23). 

Increases in water vapor concentration can also

affect polar ozone depletion on polar stratospheric clouds

(PSCs). Enhanced water vapor concentrations would

increase the critical temperature below which heterogeneous reaction on liquid aerosols become important (KirkDavidoff et al., 1999). Tabazadeh et al. (2000) estimated

that the enhancement of PSC formation for an addition of

1 ppmv water vapor is approximately the same as the PSC

enhancement from a 1 K radiative cooling. Stenke and

Grewe (2005) concluded that increased humidity will

enhance heterogeneous ozone depletion in the Antarctic

spring due to a longer PSC- existence period (Figure 5-23). 

5.3.6 Impact of Temperature Changes on

Ozone Chemistry

The assessment of ozone chemistry’s sensitivity to

changes in temperature in the atmosphere is complicated

by the concomitant changes that occur in the dynamics,

transport, and radiation as the temperature changes. In

addition, the chemical state of the atmosphere will change

as the concentration of trace species change and ozonedepleting substances (ODSs) decrease over the coming

decades. This in turn will alter the sensitivity of the stratospheric chemical system to temperature changes.

5.3.6.1 UPPER STRATOSPHERE

There is a relatively solid understanding of the sensitivity of ozone chemistry in the upper stratosphere to

changes in temperature. In this region the chemical system

is generally under photochemical control and is constrained by gas-phase reaction cycles that are well understood. The largest stratospheric cooling associated with

CLIMATE-OZONE CONNECTIONS

5.32

Figure 5-22. Impact of 1 ppmv

water vapor perturbation above a

baseline simulation in the CCM

E39C. The top panel shows the

percentage change in zonally and

monthly averaged local OH. The

bottom panel shows the percentage change in zonally and

monthly averaged local O3. Light

(dark) shading indicates regions

in which differences are statistically significant at the 95% (99%)

level (t-test). From Stenke and

Grewe, 2005.
increased greenhouse gas concentrations (see Sections

5.2.6 and 5.3.1) has been observed in the upper stratosphere and mesosphere. The dominant ozone loss cycles

in the upper stratosphere (NOx, ClOx, and HOx) are

expected to slow with decreasing temperatures (e.g., Haigh

and Pyle, 1982). CCM simulations for doubled-CO2 conditions presented by Jonsson et al. (2004) indicated an

increase of 15-20% in the ozone mixing ratios in the upper

stratosphere associated with a 10-12 K temperature

decrease (see Section 5.3.1). In the lower mesosphere, the

ozone increase is primarily due to the negative temperature dependence of the O + O2 + M → O3 + M reaction.

The situation is more complex in the upper stratosphere,

with different loss cycles having greater influence on

ozone concentrations at different altitude ranges (50-60

km: HOx; 45-50 km: all cycles; below 45 km: NOx). The

slower loss rates are controlled by the temperature dependence of the reaction rate constants and by the reduction in

the abundance of atomic oxygen (change in Ox partitioning). The rate-limiting reactions for all the loss cycles

are proportional to the atomic oxygen number density. The

atomic oxygen number density, in turn, is strongly determined by the reaction O + O2 + M → O3 + M (Jonsson et

al., 2004). 

5.3.6.2 POLAR LOWER STRATOSPHERE

In the springtime polar lower stratosphere, the

gas-phase loss cycles described above play a similar role

in determining the ozone concentration. Any decrease

in temperature is expected to slow the rate of ozone loss

(Zeng and Pyle, 2003). Counteracting this effect,

chlorine- and bromine-containing reservoir species are

activated via heterogeneous processes on cloud and cold

aerosol particles, leading to markedly increased ClOx

and BrOx concentrations. This in turn leads to significant ozone loss via ClOx and BrOx catalytic cycles in the

presence of sunlight. The rate of chlorine and bromine

activation on the surface of cloud and aerosol particles

is strongly temperature dependent, increasing sharply

below approximately 195 K. 

For the current polar lower stratosphere with elevated ODS concentrations, halogen activation, and consequent ozone losses at lower temperatures, offset any effect

of ozone increases through reduction in NOx and HOx gasphase ozone loss. This is clear from comparisons of the

Antarctic and Arctic spring ozone concentrations (see

Chapter 4). The Antarctic stratosphere routinely experiences temperatures below the threshold for heterogeneous

halogen activation during the winter and early spring, with

the consequent significant loss of ozone. The Arctic lower

stratosphere, on the other hand, is highly variable and lies

close to the halogen activation threshold. This means that

a significant change in Arctic stratospheric temperatures

would strongly influence springtime ozone concentrations

in this region.

Arctic Spring

Chapter 4 discusses the compact linear relationship

between chemical ozone loss in the Arctic winter and the

vortex-averaged volume of air that is below the PSC

threshold temperature (VPSC) introduced by Rex et al.

(2004). They deduce a linear relationship of approximately 15 DU of additional chemical ozone loss per Kelvin

of cooling of the lower stratosphere, on the basis of

ozonesonde soundings. This result is supported by a similar study by Tilmes et al. (2004). The VPSC, as defined by

Rex et al. (2004), is essentially a temperature metric and

thus the relationship they derive indicates the temperature

CLIMATE-OZONE CONNECTIONS

5.33

0 5

0 5

0 5 0 5

-0.75 0 75

1

2 3

-90

-60

-30

0

30

60

90

Latitude

Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec 

4

Figure 5-23. Impact of

1 ppmv water vapor perturbation above baseline simulation in the

CCM E39C. This figure

shows the percentage

change in zonally averaged total column

ozone. Light (dark)

shading indicates regions in which differences are statistically

significant at the 95%

(99%) level (t-test).

From Stenke and

Grewe, 2005.
sensitivity of chemical Arctic ozone loss, but other factors

(e.g., vortex isolation) are also coupled to vortex temperature and influence the chemical ozone loss over the Arctic

vortex period.

The ozone loss vs. VPSC linear relationship determined by Rex et al. (2004) is compact and robust for recent

winters. This is surprising given the highly nonlinear

nature of some processes within the system being

described. Their relationship was derived by averaging

over the Arctic vortex for the period mid-December to the

end of March, which concatenates a number of complex

nonlinear processes (microphysical details of PSC formation, denitrification and dehydration, and chemical activation on PSC surfaces). 

It should be noted that the ozone loss vs. VPSC relationship is valid for current conditions of elevated ODSs.

As concentrations of ODSs in the stratosphere decrease

through the early part of this century, it is likely that the

ozone loss vs. VPSC relationship will change. Also, vortex

conditions (e.g., size) may change, so it is unclear whether

the ozone loss vs. VPSC relationship has any use as a predictive tool.

Antarctic Spring

In contrast to the Arctic polar vortex, the extent of

ozone loss within the Antarctic polar vortex during winter

and spring is primarily dictated by the amount of stratospheric halogen present. 

A major predictor for ozone mass deficit over the

Antarctic vortex period for the years 1979 to 2003 is the

lagged equivalent effective stratospheric chlorine (EESC);

it explains 82% of the variance in the ozone mass deficit

(Huck et al., 2005). The 100 hPa polar temperatures modulate the interannual variability in the ozone mass deficit

anomaly. The Antarctic ozone hole size is primarily sensitive to EESC and secondarily to temperatures near the

edge of the vortex (collar temperature) (Newman et al.,

2004). Changes in sulfate aerosol particle surface area

and surface reactivity were used to explain the correlation

between the residual ozone hole size and collar temperature. Using the residuals from a quadratic fit of the ozone

hole size to the EESC amounts, they estimated that a 1 K

decrease in collar temperature leads to an increase in size

of the late September ozone hole of 1.1 million km2. 

Both Huck et al. (2005) and Newman et al. (2004)

used regression analysis based on data from recent (1979-

2003) years to derive empirical models of Antarctic ozone

loss. The models can only be expected to remain valid for

descriptions of near-future ozone losses when stratospheric conditions remain broadly similar to the 1979-

2003 period used to construct the models. Both models’

ozone hole indices will become less meaningful as measures of Antarctic ozone levels as the ozone hole begins to

close from the middle of this century.

5.3.7 Impact of Climate Change on Ozone

Recovery 

As discussed in the previous sections, the long-term

evolution of stratospheric ozone concentrations depends

not only on changes of many stratospheric constituents

(including ozone-depleting substances (ODSs), greenhouse gases (GHGs), water vapor, and aerosols), but also

on changes in the climate of the troposphere and stratosphere caused by natural variability and anthropogenic

forcing. While it is expected that the reduction of ODSs

in the next years to decades will lead to an increase in

ozone, this increase could be affected by changes in temperature (Section 5.3.6) and in chemical composition and

transport (Sections 5.3.3 through 5.3.5). This section discusses results of sensitivity studies that use CCMs to elucidate how climate change could affect formation and

destruction of future stratospheric ozone.

Recent Assessment reports (Section 12.2.1.3 in

WMO, 1999; Section 4.8 in WMO, 2003) presented results

derived from an ensemble of two-dimensional (2-D)

models (see Box 5-1) that calculated the evolution of total

ozone through the year 2050. For some models, nearglobal mean (60°N-60°S) column ozone remained up to

1% below 1980 levels even in 2050, while for others the

ozone column amounts in 2050 were 3.5% higher than in

1980. Several of the 2-D models did not include the impact

of greenhouse gas-induced stratospheric cooling on ozone.

Temperatures affect ozone by changing the reaction rates

that determine Ox abundance, primarily in the middle and

upper stratosphere. Moreover, in the polar lower stratosphere, in addition to gas-phase photochemistry, chlorine

and bromine reservoir species are activated via heterogeneous processes on cloud and cold aerosol particles (see

Section 5.3.6). For example, in the 2-D models of

Rosenfield et al. (2002) and Chipperfield and Feng (2003),

ozone recovers to 1980 values 10-20 years earlier in many

latitudes and seasons, due to the inclusion of the stratospheric cooling effect from enhanced GHG concentrations

(e.g., Figure 6-11 of Chapter 6). 

The aforementioned 2-D model studies illustrate that

the future evolution of ozone is sensitive to changes in both

chemical constituents and climate. Although a few 2-D

models include some of the effects of changes in dynamics,

they do not consider any variability in tropospheric wave

forcing due to climate change and the impact this may have

on the stratosphere (see Sections 5.2.2 and 5.3.2). Another

weak point of 2-D models is that they are simply not useful

for investigations of polar regions (WMO, 2003).

Nowadays CCMs offer a more complete possibility to furCLIMATE-OZONE CONNECTIONS

5.34
ther investigate the impact of climate change on ozone,

since these models do consider interactions and feedbacks

of radiative, dynamical, and chemical processes. Recently

a number of CCMs (see Chapter 6, Table 6-4) have been

used to simulate the past and future evolution of ozone (see

Chapter 6, Appendix 6A). The following discussion highlights some results of additional CCM sensitivity studies

(using the “no climate change” scenario, NCC) that enable

a more detailed investigation of the impact of fixed greenhouse gas concentrations on the recovery of the stratospheric ozone layer. The results derived from these model

sensitivity studies provide useful and supplementary information to better understand the expected evolution of the

ozone layer in the 21st century (Chapter 6). 

Three CCMs (E39C, ULAQ, WACCM) were

employed to carry out long-term sensitivity simulations

(through 2050), wherein concentrations of well-mixed

greenhouse gases (CO2, CH4, and N2O) are held constant

after a given date (E39C: 1980; ULAQ: 1970; WACCM:

2000), whereas the evolution of halocarbon is based on

the “Ab” scenario from WMO (2003). Additionally, in

E39C and ULAQ, the sea surface temperatures (SSTs)

were prescribed according to observations of the years

1970 to 1979. This SST dataset was used again for every

succeeding decade in the model simulation (E39C: 1980

to 2020; ULAQ: 1960 to 2050). The methods used to prescribe SSTs in the NCC simulations are a bit simplified,

but reasonable for this kind of sensitivity study. In reality,

the thermal inertia of the oceans must be considered. A

mixed-layer ocean model needs approximately 20 to 30

years to reach a new equilibrium, and a deep-ocean model

even several hundred years (e.g., Hansen et al., 2005). For

this kind of sensitivity study, the neglect of this effect is

expected to be of small significance. Nonetheless, this

effect was considered in the WACCM NCC simulations

using SSTs from a coupled ocean-atmosphere model

(AOGCM) that was run in support of the most recent IPCC

Assessment (IPCC CCSM/CAM3). The CCSM/CAM3

simulation was run with IPCC emissions for 2000 held

constant from 2000 to 2100, i.e., the same GHG scenario

as adopted for the NCC simulations. 

In all three CCMs, ozone and water vapor are prognostic variables, i.e., the radiative feedback of ozone and

water vapor changes is considered. A direct comparison

of results from the NCC simulations with results from reference simulations (prescribing past and expected future

changes, i.e., REF1 and REF2/SCN2; see Chapter 6,

Appendix 6A) allows a qualitative estimation of the future

impact of climate change on ozone. Figure 5-24 (top) displays the evolution of the near-global (60°N-60°S) mean

temperature deviations of the lower stratosphere at 50

hPa. The results derived from the reference simulations

show an obvious cooling trend because of increasing

GHG concentrations (see Section 5.3.1). Nearly zero

trends are calculated by all three CCMs when fixed GHG

concentrations are assumed. It takes about 20 years before

a difference is apparent between curves derived from the

reference and the sensitivity simulations. This statement

is not valid for mean temperatures (at 50 hPa) over the

Northern Hemisphere polar region (i.e., 60°-90°N, Figure

5-24 middle panel), which does not show statistically significant differences between the reference and the sensitivity simulations. The picture is unclear for the Southern

Hemisphere polar region (i.e., 60°-90°S, Figure 5-24

bottom panel). E39C shows no obvious differences

between the two simulations, whereas WACCM results

indicate slightly higher and ULAQ much higher temperatures in the NCC simulations. 

Investigations of the CCM results do not indicate a

clear impact of climate (temperature) change on metrics

like minimum ozone column values in the polar stratosphere or “maximum ozone hole area,” either in the

Northern Hemisphere or the Southern Hemisphere (not

shown). These metrics are evidently insensitive to temperature changes of less than about 1 K. But the results

shown in Figure 5-25 demonstrate that the recovery of

near-global mean (60°N-60°S) column ozone is accelerated in a changing climate: since the global stratosphere is

under photochemical control (constrained by gas-phase

reaction cycles), the ozone loss cycles in the stratosphere

(primarily Ox) slow down with decreasing temperature.

When climate change is included, ULAQ and WACCM

show a return to the 1980 global ozone amount approximately 10 to 15 years earlier than without climate change.

This is in agreement with former estimates derived from

2-D model studies (see above). The results derived from

the CCM E39C are not as clear in this sense, but here it

must be considered that in contrast to ULAQ and

WACCM, E39C has a low uppermost layer which is centered at 10 hPa, and therefore neglects the impact of

dynamical and photochemical effects on ozone in the

upper part of the stratosphere. In any case, the E39C simulations were not long enough for the difference between

NCC and REF simulations to become apparent. 

Results derived from ULAQ and WACCM show

that by 2050, the middle to upper stratosphere is 5-10 K

colder in the reference runs than in corresponding NCC

simulations (not shown). As a result of this cooling, ozone

mixing ratios in the middle and upper stratosphere are

higher in the reference cases than in NCC (up to 15%

below the stratopause). The most important impact is

through the effect of temperature on the ozone loss rate.

Below 10 hPa, all three CCMs show a consistent behavior.

These results indicate a delay in the recovery of ozone in

CLIMATE-OZONE CONNECTIONS

5.35
the Antarctic lower stratosphere (not shown) due to “intensified” heterogeneous chemistry that is caused by reduced

temperatures in the lower polar stratosphere. The changes

are not statistically significant in the Arctic lower stratosphere. Interestingly, all models very clearly show less

ozone in the tropical lower stratosphere in the reference

simulations, which may point to an enhanced updraft in

the tropics (intensified Brewer-Dobson circulation) in the

reference simulations. 

Qualitatively, the evolution of column ozone and

lower stratospheric temperatures (50 hPa) in the polar

regions (60°-90°) is very similar in the reference and NCC

simulations (Figure 5-26). The results seem to indicate

that the net impact of changes of well-mixed greenhouse

gas concentrations and SSTs on total ozone is perhaps

slightly larger in the Arctic stratosphere than in the

Antarctic stratosphere. Differences in ozone column

between results derived from reference simulations and

NCC are not apparent in the polar SH. In the polar NH,

the WACCM model indicates that after the year 2020, the

ozone column is systematically larger in the reference simulations. So, lower temperatures in the ozone loss regions

of high latitudes have the opposite effect of lower temperatures in the ozone production region in the tropics.

Notwithstanding, it may be that in the NH the transport of

air with enhanced ozone mixing ratios to high latitudes by

an intensified Brewer-Dobson circulation dominates the

effect of enhanced chlorine/bromine catalysis of ozone

destruction in the polar region. The fact that it takes at

least until 2020 for a noticeable difference to emerge

between the reference simulations and NCC in the NH

may be consistent with this conjecture.

CLIMATE-OZONE CONNECTIONS

5.36

Figure 5-24. Time series of zonally averaged nearglobal (top panel, 60°N-60°S) temperature deviations

at 50 hPa between 1960 and 2050 relative to the year

1980 (in K). Data have been smoothed with a 13-

month running mean. Respective time series are

shown for the northern polar region (middle panel,

60°N-90°N) and the southern polar region (lower

panel, 60°S-90°S). Here, data have been considered

only for springtime months, i.e., 3-month mean

values of February, March, and April in the NH, and

August, September, and October in the SH. For each

model, i.e., E39C (red lines), ULAQ (black lines), and

WACCM (blue lines), the solid curves show results

derived from the reference simulations (REF). Since

WACCM and E39C have performed ensemble simulations, the envelope of results is shown for each

model. The dotted curves indicate the results derived

from the “no climate change runs” (NCC). 
CLIMATE-OZONE CONNECTIONS

Figure 5-25. Time series of zonally averaged near-global (60°N-60°S) total ozone

deviations between 1960 and 2050 with

regard to the year 1980 (in %). For each

model, i.e., E39C (red lines), ULAQ (black

lines), and WACCM (blue lines), the solid

curves show results derived from the reference simulations (REF). Since WACCM and

E39C have performed ensemble simulations, the envelope of results is shown for

each model. The dotted curves indicate the

results derived from the “no climate change

runs” (NCC). All data are smoothed with a

13-month running mean.

Figure 5-26. Time series of zonally averaged total

ozone deviations at southern polar latitudes (top

panel, 60°S-90°S) and northern polar latitudes

(lower panel, 60°N-90°N) between 1960 and 2050

with regard to the year 1980 (in %). For each model,

i.e., E39C (red lines), ULAQ (black lines), and

WACCM (blue lines), the solid curves show results

derived from the reference simulations (REF). Since

WACCM and E39C have performed ensemble simulations, the envelope of results is shown for each

model. The dotted curves indicate the results

derived from the “no climate change runs” (NCC).

Data have been considered only for springtime

months, i.e., 3-month mean values of February,

March, and April in the NH, and August, September,

and October in the SH.

5.37
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Appendix 5A

AOGCMs USED IN THIS CHAPTER

The coupled atmosphere-ocean general circulation models (AOGCMs) used in this chapter, and the country of the

sponsoring organization. The model simulations were completed to support the Fourth Assessment Report of the

Intergovernmental Panel on Climate Change (IPCC, in preparation for 2007). The models submitted 21st century simulations using the A2 (high) and B1 (low) emission scenarios. Also listed is the key reference for the atmospheric model, the

approximate altitude in kilometers of the model top, and the number of stratospheric levels at or above 16 km. The experiments include various emission scenarios of the 20th and 21st century and focus on simulating the surface and tropospheric climate. The stratosphere of most of these models tends to be poorly resolved. While all models include changes

in well-mixed greenhouse gases, other forcings such as changes in ozone, volcanic and other aerosols, and changes in

solar radiation are only included in a subset of the models. This set of experiments was produced by 15 modeling groups

from 9 countries and represents one of the most comprehensive international climate model efforts ever attempted. 

Stratospheric

Model Key Reference Model Top Levels

BCCR-BCM2.0, Norway Déqué et al., 1994 33 5

CCSM3, USA Collins et al,. 2004 40 7

CGCM3.1(T47), Canada Flato, 2005 49 11

CNRM-CM3, France Déqué et al., 1994 76 17

CSIRO-Mk3.0, Australia Gordon et al., 2002 38 3

ECHAM5/MPI-OM, Germany Roeckner et al., 2003 29 4

GFDL-CM2.0, USA GFDL GAMDT, 2004 35 3

GFDL-CM2.1, USA GFDL GAMDT, 2004 35 3

GISS-ER, USA Schmidt et al., 2006 67 9

INM-CM3.0, Russia Galin et al., 2003 32 6

IPSL-CM4, France Hourdin et al., 2006 32 7

MIROC3.2-M, Japan K-1 Developers, 2004 67 6

MRI-CGCM2.3.2, Japan Shibata et al., 1999 54 8

PCM, USA Kiehl et al., 1998 43 7

UKMO-HadGEM1, UK Martin et al., 2004 39 10
CHAPTER 6

The Ozone Layer in the 21st Century

Lead Authors:

G.E. Bodeker

D.W. Waugh

Coauthors:

H. Akiyoshi

P. Braesicke

V. Eyring

D.W. Fahey

E. Manzini

M.J. Newchurch

R.W. Portmann

A. Robock

K.P. Shine

W. Steinbrecht

E.C. Weatherhead

Contributors:

J. Austin

S. Bekki

C. Brühl

N. Butchart

M. Chipperfield

M. Dameris

T. Egorova

V. Fioletov

A. Gettelman

M.A. Giorgetta

D. Kinnison

E. Mancini

M. Marchand

P.A. Newman

S. Pawson

G. Pitari

D. Plummer

B. Rognerud

E. Rozanov

R.J. Salawitch

T.G. Shepherd

K. Shibata

B.-M. Sinnhuber

M. Sinnhuber

S. Smyshlyaev

R. Stolarski

H. Struthers

W. Tian

G. Velders

D. Weisenstein

E.-S. Yang

Final Release: February 2007

From Scientific Assessment of Ozone Depletion: 2006

CHAPTER 6

THE OZONE LAYER IN THE 21ST CENTURY

Contents

SCIENTIFIC SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.1

6.1 INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.5

6.2 A FRAMEWORK FOR EVALUATING CHANGES IN OZONE ABUNDANCES . . . . . . . . . . . . . . . . . . . . . 6.5

6.2.1 Ozone Changes in the Near and Long Term . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.5

6.2.2 Stages in the Evolution of the Ozone Layer . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.6

6.2.3 Milestones in the Evolution of the Ozone Layer . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.7

6.2.4 Using Data and Models to Evaluate Ozone Milestones and Milestone Parameters . . . . . . . . . . . . . . . 6.8

6.3 FACTORS AFFECTING THE DETECTION, ATTRIBUTION, AND TIMING OF MILESTONES . . . . . . . 6.8

6.3.1 Stratospheric Halogen Loading . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.8

6.3.2 Atmospheric Chemical Composition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.9

6.3.3 Stratospheric Temperatures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.9

6.3.4 Atmospheric Transport . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.10

6.3.5 The Solar Cycle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.10

6.3.6 Volcanic Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.11

6.4 STATISTICAL METHODS FOR DETECTION OF MILESTONES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.11

6.4.1 Change in Linear Trends . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.12

6.4.2 Cumulative Sum of Residuals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.12

6.5 ATTRIBUTION OF THE RECENT BEHAVIOR OF OZONE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.13

6.5.1 Upper Stratospheric Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.13

6.5.2 Lower Stratospheric and Total Column Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.15

6.5.3 Polar Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.18

6.6 PROJECTIONS OF THE FUTURE BEHAVIOR OF OZONE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.20

6.6.1 Model Descriptions and Scenarios . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.20

6.6.2 Model Evaluation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.22

6.6.3 Midlatitude and Tropical Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.27

6.6.4 Polar Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.30

6.6.4.1 The Antarctic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.31

6.6.4.2 The Arctic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.33

6.6.5 Uncertainties in Model Projections and Open Questions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.36

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.37

APPENDIX 6A: MODEL SCENARIOS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.43

21st CENTURY OZONE LAYER

6.1

SCIENTIFIC SUMMARY

Global ozone levels are now no longer declining as they were from the late 1970s until the mid-1990s, and some

increases in ozone have been observed. These improvements in the ozone layer have occurred during a period when stratospheric halogen abundances reached their peak and started to decline. These declining halogen abundances clearly

reflect the success of the Montreal Protocol and its Amendments and Adjustments in controlling the global production and

consumption of ozone-depleting substances (ODSs).

Stratospheric ozone abundances are affected by a number of natural and anthropogenic factors in addition to the

atmospheric abundance of ODSs, e.g., temperatures, transport, volcanoes, solar activity, and hydrogen and nitrogen

oxides (Chapter 3). Separating the effects of these factors is complex because of nonlinearities and feedbacks in the

atmospheric processes affecting ozone. For the purposes of this Assessment, we consider specifically the recovery of

ozone from the effects of ODSs, because the primary audience is the group of Parties to the Montreal Protocol, whose

purview is ozone-depleting compounds. In this Assessment, the metric used to gauge the overall burden of ozonedepleting halogens in the stratosphere from the ODSs is equivalent effective stratospheric chlorine (EESC).

The Process of Ozone Recovery from the Effects of ODSs

• In this Assessment, the recovery of ozone from depletion caused by increases in ODSs is discussed as a process

involving three stages:

(i) The slowing of ozone decline, identified as the occurrence of a statistically significant reduction in the rate of

decline in ozone due to changing EESC. 

(ii) The onset of ozone increases (turnaround), identified as the occurrence of statistically significant increases

in ozone above previous minimum values due to declining EESC.

(iii) The full recovery of ozone from ODSs, identified as when ozone is no longer significantly affected by ODSs.

In the absence of changes in the sensitivity of ozone to ODSs, this is likely to occur when EESC returns to pre1980 levels.

The first two stages of recovery either have already occurred or are expected to occur within the next two decades.

The third stage is expected to occur around the middle of the century. Because of changes in atmospheric composition and dynamics, this third stage may or may not be accompanied by the actual return of ozone to pre-1980 levels,

and it is possible that ozone could return to 1980 levels before the effects of ODSs disappear. 

• In reaching full recovery of ozone, the milestone of the return of ozone to pre-1980 levels is considered important

because ozone was not significantly affected by ODSs prior to 1980. As a consequence, this milestone is useful, for

example, to gauge when the adverse impacts of enhanced surface ultraviolet (UV) radiation on human health and

ecosystems caused by ozone depletion are likely to become negligible. However, as mentioned above, the return of

ozone to pre-1980 levels may not occur at the same time as the return of EESC to pre-1980 levels, and in fact may

never occur because of changes in the atmosphere since 1980 that are not caused by ODSs. Therefore, this milestone alone cannot be used to identify the recovery of ozone from the effects of ODSs.

The Role of ODSs in Recent Ozone Trends

• The slowing of the decline and leveling off of midlatitude upper stratospheric (35-45 km) ozone over the past

decade has very likely been dominated by changes in EESC. Gas-phase chemistry, modulated by changes in

temperature and other gases such as methane (CH4), directly controls ozone in this region, and observed ozone

changes are similar to those modeled from EESC decreases.

• Over the past decade, changes in EESC have likely contributed to the slowing of the midlatitude total column

ozone decline and the leveling off of ozone. However, evidence suggests that changes in transport have also

played an important role, particularly in the lowermost stratosphere, making attribution of specific ozone changes to

EESC more complicated. For northern midlatitudes, increases in ozone have been greater than expected from
21st CENTURY OZONE LAYER

6.2

EESC decreases alone, while over southern midlatitudes the observed ozone changes are broadly consistent with the

expectations from EESC decreases alone.

• Inside the Antarctic vortex, the interannual variations in ozone depletion observed from 2001 to 2005 have

not been caused by changes in EESC. At current EESC concentrations, nearly total loss of ozone occurs in the

lowermost stratosphere inside the ozone hole in September and October, and EESC concentrations often exceed

those necessary to cause total loss. The Antarctic ozone hole, therefore, has low sensitivity to moderate decreases in

EESC and the unusually small ozone holes in some recent years (e.g., 2002 and 2004) are strongly attributable to a

dynamically driven warmer Antarctic stratosphere.

• In the collar region of the Antarctic vortex (60°S-70°S), where ozone destruction is not complete, reductions

in EESC have likely contributed to the slowing of ozone decline observed over the past decade. However,

uncertainty in the estimation of EESC in the collar region and the ozone response to temperature changes confound

the attribution of observed ozone changes to reductions in EESC.

• The decline in EESC has not caused the large interannual variations observed in Arctic ozone depletion.

Indeed there has been no detection of any ozone recovery stages in the Arctic. The large interannual variations in

ozone are driven by changes in meteorology and are likely to delay the detection of the first stage of recovery.

Expected Future Changes in Ozone

Two-dimensional (2-D) models and three-dimensional (3-D) coupled Chemistry-Climate Models (CCMs), both of

which have achieved significant successes in simulating many or nearly all of the factors that affect ozone and their feedbacks, have been used to project the evolution of ozone throughout the 21st century. The evolution of tropical and midlatitude ozone was examined in all models. CCMs are generally believed to better represent key processes relating to threedimensional transport in the polar regions and, therefore, only CCMs were used for polar regions. The projected total

column ozone was examined for three periods:

(i) The beginning of the century (2000-2020), when EESC is expected to start decreasing or continue to decrease

(ii) Mid-century (2040-2050 in extrapolar regions, 2060-2070 in polar regions), when EESC is expected to reach

and fall below 1980 values

(iii) End of the century (2090-2100), when changes in factors other than ODSs are expected to control changes in

stratospheric ozone

Because modeled changes in column ozone to 2100 are not specifically attributable to changes in EESC, ozone column

amounts when EESC returns to pre-1980 levels and the timing of the return of ozone to pre-1980 levels are examined in

the model projections.

• The CCMs used to project future ozone abundances have been critically evaluated, and more emphasis has

been given to those models that best represent the processes known to strongly affect column ozone abundances. The CCMs vary in their skill in representing different processes and characteristics of the atmosphere.

However, there is sufficient agreement between the majority of the CCMs and the observations that some confidence can be placed in their projections.

BEGINNING OF THE CENTURY

• Averaged between 60°S and 60°N, total column ozone is projected to increase in all models between 2000 and

2020, with most of the increase of 1 to 2.5% occurring after 2010. The small interannual variability shown in 2-

D models allows more precise identification of key ozone change dates compared with CCMs that show large interannual variability, similar to that seen in observations. Nonetheless, both the 2-D models and CCMs suggest that

minimum total column ozone values have already occurred in this latitude region.
21st CENTURY OZONE LAYER

6.3

• Over the Antarctic, most CCMs predict column ozone increases in spring of around 5 to 10% between 2000

and 2020. Different diagnostics of ozone depletion show different sensitivities to EESC. The most rapid change

(decrease) occurs in the ozone mass deficit and the slowest change (increase) in ozone minimum values and October

ozone anomalies. Minimum ozone values remain roughly constant between 2000 and 2010 in many models. The

projected onset of decreases in the ozone mass deficit occurs between 2000 and 2005, whereas the projected onset

of increases in minimum Antarctic ozone does not occur until after 2010 in many models.

• Over the Arctic, most CCMs predict that springtime column ozone in 2020 will be 0 to 10% above 2000 levels

and that ozone turnaround in the Arctic will occur before 2020. Over the Arctic, the large interannual variability

in the CCM projections obscures the year when the ozone turnaround due to decreasing EESC occurs.

MID-CENTURY

• Averaged between 60°S and 60°N, total column ozone is projected to be close to or above 1980 values when

EESC in that region of the stratosphere declines to 1980 values (2040-2050). This occurs in nearly all models

that include coupling between well-mixed greenhouse gases (WMGHGs; carbon dioxide (CO2), methane (CH4),

and nitrous oxide (N2O)) and temperature (interactive 2-D models and the CCMs). Thus, outside polar regions,

ozone is projected to reach 1980 values at about the same time or before EESC returns to 1980 values.

• Two-dimensional models that include coupling between WMGHGs and temperature (interactive 2-D models)

predict that column ozone averaged over 60°S to 60°N will in general exceed 1980 values up to 15 years earlier than models that do not include these feedbacks (non-interactive 2-D models). This suggests that the earlier return to 1980 values is caused mainly by response to stratospheric cooling linked to increased WMGHGs. The

response results from the temperature dependence of the gas-phase photochemistry of ozone. In some non-interactive 2-D models, column ozone never increases to 1980 values throughout the 21st century in some regions.

• Over the Antarctic, most models predict that ozone amounts will increase to 1980 values close to the time

when Antarctic EESC decreases to 1980 values. That time is later than at midlatitudes due to the delay associated

with transport of stratospheric air to polar regions. A new empirical model, based on observations, indicates a return

of Antarctic EESC to 1980 values between 2060 and 2075. 

• Over the Arctic, CCMs show ozone values exceeding 1980 values before EESC decreases to 1980 values, with

ozone increasing to 1980 values between 2020 and 2040. The increases in ozone do not follow the decreases in

EESC as closely as in the Antarctic, and in the majority of CCMs Arctic ozone exceeds 1980 values before the

Antarctic. There is no indication of future severe reductions in Arctic column ozone in any of the model simulations. There is large uncertainty in projections of Arctic ozone because of the smaller ozone depletion and the larger

interannual variability in the Arctic stratosphere in comparison with the Antarctic.

END OF THE CENTURY

• Averaged between 60°S and 60°N, total column ozone is projected to be around 2 to 5% above 1980 values

between 2090 and 2100. This result is obtained in all 2-D models that include coupling between WMGHGs and

temperature and in one CCM that extends to 2100. This CCM predicts that from 2090 to 2100, Arctic ozone will be

substantially above 1980 values, while Antarctic ozone will be close to or just below 1980 values.

• Projected ozone amounts in 2100 are sensitive to future levels of WMGHGs. For example, expected future

increases in N2O will increase stratospheric nitrogen oxides (NOx), which may exacerbate ozone depletion. However,

the expected stratospheric cooling induced by increasing concentrations of greenhouse gases, primarily CO2, is

expected to slow gas-phase ozone depletion reactions and, thereby, increase ozone. The net effect on ozone amounts

will depend on future levels of the different WMGHGs. The importance of this temperature feedback is demonstrated by the non-interactive 2-D models, which predict that extrapolar column ozone will be less than or near 1980

values through the latter half of the century.

6.1 INTRODUCTION

It is now clear that the depletion of the ozone layer,

both globally and in the polar regions, is attributable to an

atmospheric halogen burden that is strongly enhanced

compared with natural levels by anthropogenic emissions

of ozone-depleting substances (ODSs). Over the past

decade, as a consequence of adherence to the Montreal

Protocol and its Amendments and Adjustments, equivalent effective stratospheric chlorine (EESC; see Chapter

8), a commonly used measure of the stratospheric halogen

burden, has peaked and begun to decline (Chapter 1).

However, owing to the long lifetime of the most important

halogen source gases, the removal of ODSs from the

atmosphere will take many decades even with continued

compliance with the Montreal Protocol (Chapter 8).

Ozone is expected to continue to respond to these changes

in ODSs but the timing and sensitivity of the response will

depend on other changes in the atmosphere, e.g., increases

in well-mixed greenhouse gases (WMGHGs) that have

occurred since the onset of significant ozone depletion in

1980.

Chapters 3 and 4 discuss ozone changes observed

to date and interpret the underlying causes of the changes,

while Chapter 5 examines the mechanisms connecting

ozone depletion with climate change. This chapter builds

on the material presented in these three chapters, with a

focus on two overarching themes:

(1) Analysis of recent ozone measurements for the first

signs of ozone recovery attributable to decreasing

ODS concentrations

(2) Projections of how the global ozone layer will evolve

during the 21st century and how this evolution will

depend on ODSs and other concomitant changes in

the atmosphere

The first of these two themes uses the observational

time series presented in Chapters 3 and 4 to assess whether

ozone recovery (defined in Section 6.2) has been detected.

The interpretation of observations presented in the two

earlier chapters is extended by applying different statistical tools (Section 6.4) to the detection of the first stage

of ozone recovery, i.e., a slowing of the ozone decline

attributable to decreasing ODSs. This chapter also examines when increases in ozone due to decreases in ODSs

are likely to occur, i.e., the second stage in the ozone

recovery process.

The second theme builds on the discussion of the

feedbacks between ozone depletion and climate change

presented in Chapter 5. A group of two-dimensional

(2-D) models and coupled Chemistry-Climate Models

(CCMs), incorporating climate feedback processes, is used

to predict the evolution of the ozone layer throughout the

21st century. While Chapter 5 discusses ozone depletion/-

climate change interactions, this chapter uses models to

investigate the effects of those interactions and addresses

how they may affect the evolution of ozone throughout

this century. The resulting ozone projections are then used

in Chapter 7 to show how surface clear-sky ultraviolet

(UV) radiation is expected to change in the future.

Section 6.2 describes the framework for assessing

ozone changes through the 21st century. Specifically it discusses the time scales of the expected changes, stages in the

recovery process and the milestones defining those stages.

Section 6.3 includes a discussion of how factors other than

ODSs affect the detection and timing of milestones in the

ozone recovery process and the attribution to changing

EESC. Section 6.4 presents how least squares regression

models (also discussed in Chapter 3) can be used to evaluate the role of ODSs and other factors in ozone recovery.

The interpretation of recent observations of ozone in the

context of ozone recovery is presented in Section 6.5. The

longer-term perspective of the evolution of global ozone

through the 21st century is provided in Section 6.6, where

the results from the 2-D models and CCMs are assessed.

6.2 A FRAMEWORK FOR EVALUATING

CHANGES IN OZONE ABUNDANCES

In considering the ozone layer in the 21st century,

two topics are prominent in the scientific and policy communities. The first is obtaining evidence as soon as possible that the ozone layer is responding to the decline in

EESC. As ODSs are removed from the atmosphere, the

destruction of ozone attributed to ODSs is also expected

to decline. The second topic is the projection of changes

that might occur in the ozone layer by mid-century and by

the end of the century. By approximately mid-century,

substantial or full recovery of the ozone layer is expected

globally. By the end of the century, other changes in

atmospheric chemistry and transport become a prominent consideration in projected ozone amounts. Here,

we have adopted a specific framework, described below,

to evaluate changes in ozone over the recent past and

until the end of the century. This framework is conceptual in design and is intended to encompass both ozone

changes that have already occurred as well as potential

future changes.

6.2.1 Ozone Changes in the Near and Long

Term

The discussion of ozone changes in the 21st century

is divided into those in the “near term,” which includes

21st CENTURY OZONE LAYER

6.5
the recent past, the present day, and the first decades of

the century, and those in the “long term,” which includes

the remaining decades of this century. In the near term,

when ozone observations and model simulations can be

compared, the first changes in the ozone layer attributable

to the ongoing removal of ODSs can be detected. In contrast, changes in the last decades of this century can only

be discussed as model projections and, hence, are associated with large uncertainties.

Within each time period, the analyses of ozone

changes will focus separately on changes that occur, for

example, in global average total column ozone, midlatitude and tropical column ozone, as well as springtime polar

ozone amounts. The role of ODSs in ozone changes is of

particular importance, but the role of other factors, such

as changes in abundances of other gases and of sulfate

aerosols, stratospheric temperatures, atmospheric transport, solar output, and volcanic emissions, will also be

carefully examined (see Section 6.3). Without a comprehensive examination of all factors that significantly influence ozone amounts, the contribution of ODSs cannot be

quantified with sufficient confidence.

6.2.2 Stages in the Evolution of the Ozone

Layer

Ozone changes in the 21st century are expected to

encompass the period of the “recovery” of ozone from the

influence of ODSs that have been released in anthropogenic

activities. ODSs have led to the decline in global ozone

amounts, with measurable changes beginning in the 1980s

and the largest changes found in the winter/spring polar

stratospheres (see Chapters 3 and 4). The total abundance

of ODSs is in decline in the troposphere and stratosphere

as a result of the effective actions undertaken as part of the

Montreal Protocol (see Chapter 1). The start of the decline

in EESC marks the conceptual start of the ozone recovery

process. The continued analysis of ozone measurements

21st CENTURY OZONE LAYER

6.6

Figure 6-1. A schematic diagram of the temporal evolution of global ozone amounts beginning with pre-1980

values, which represent amounts before significant depletion due to anthropogenic ODS emissions, and stopping at the end of the 21st century. Observed and expected ozone amounts (solid red line) show depletion from

pre-1980 values and the three stages of recovery from this depletion (see Section 6.2). The red-shaded region

represents the range of observations and model results for both near-term and long-term ozone changes. The

blue-shaded region represents the time period when declining global ODS concentrations are expected to reach

1980 values. The full recovery of ozone from ODSs may be delayed beyond the return of ODSs to 1980 levels

by factors (e.g., a volcanic eruption close to that time) that could change the sensitivity of ozone to ODSs.

Stage i: Slowing 

of ozone decline

1960 End of 21st 

century Time

Ra ge of p ojectio s

Stage ii On t o

ozone inc ea

Stage iii: F ull 

ecov r y of ozone 

om ODS

i c easi g o o e dec easi g o o e

Ozone change from pre 1980

1980

Expected 

etu rn of 

E ESC to 980

le el
during the 21st century is expected to reveal and confirm

three key stages in the recovery process:

(i) The slowing of ozone decline, identified as the occurrence of a statistically significant reduction in the rate

of decline in ozone due to changing EESC

(ii) The onset of ozone increases (turnaround), identified as the occurrence of statistically significant

increases in ozone above previous minimum values

due to declining EESC

(iii) The full recovery of ozone from ODSs, identified as

when ozone is no longer affected by ODSs. In the

absence of changes in the sensitivity of ozone to

ODSs, this is likely to occur when EESC returns to

pre-1980 levels

As illustrated in Figure 6-1, the first two stages of

recovery are expected to occur in the near term and the

last stage in the long-term future. These three stages apply

to both total column ozone and ozone at a specific altitude. However, the timing of the stages may be different

for these different measures of ozone. Documenting how

and when each of these stages of ozone layer recovery is

reached will be of interest to both the scientific and policy

communities.

The focus of this chapter is how, and to what extent,

observed and projected changes in ozone during the

recovery process will be attributable to changes in ODSs

and to other contributing factors as noted above.

The role of contributing factors in the recovery

process is important because the physical and chemical

environment of the atmosphere has changed significantly

since the onset of observable ozone depletion in the 1980s.

A key aspect of ozone recovery is whether or not ozone

abundances in stage (iii) will be greater than or less than

those present in 1980 before significant depletion by ODSs

occurred. As shown in Figure 6-1, projections of ozone

abundances show a range of values during the full recovery

stage because of uncertainties in the role of contributing

factors in controlling ozone in the coming decades. At

present there is no standard definition of stages in the

ozone recovery process and no standard framework that

can be used to analyze ozone observations for signs of

recovery. Authors of previous studies have used terms

such as “recovery” with a range of different meanings.

For example, not all studies have incorporated the concept

of attribution in the definition of ozone recovery. The

framework presented here with three defined stages of

ozone recovery will provide a more rigorous basis for

future analyses of ozone changes.

6.2.3 Milestones in the Evolution of the

Ozone Layer

The stages of recovery outlined above apply to the

overall response of the ozone layer. As an aid to monitoring and documenting each stage of recovery, milestones

can be defined within the recovery stages. A milestone is

a point in a recovery stage that a specific change in a specific ozone parameter can be said to have occurred.

Milestones and milestone parameters reflect that ozone

changes during the recovery process will vary with geographic region and altitude, following in large part the

known variations and differences in ozone depletion.

Primary geographic regions are the Antarctic and Arctic

regions, midlatitudes, and the tropics. Primary altitude

regions are the upper and lower stratosphere. Table 6-1

summarizes a number of possible milestone parameters,

some of which will be discussed in the following sections.

21st CENTURY OZONE LAYER

6.7

Table 6-1. Examples of milestone parameters in ozone recovery stages. 

Stratospheric Region Milestone Parameters

Antarctic during winter/early spring (60°S-90°S) Daily minimum column ozone values

Minimum values of the ozone partial column between 12 and 20

km above the South Pole

Area of the 220 DU contour of the ozone hole

Vortex average daily ozone mass deficit

Seasonal trend in average column ozone 

Arctic during winter/early spring (60°N-90°N) Daily minimum column ozone values

Seasonal trend in average column ozone

Midlatitudes (35°-60°) Trend in column ozone 

Trend in upper stratospheric column ozone (35-45 km)

Trend in lower stratospheric ozone (15-20 km)

Tropics (25°N-25°S) Trend in column ozone
Milestone parameters are all quantifiable from

observational datasets. A milestone derives directly

from a parameter by choosing a quantitative threshold

or limit value. Reaching or passing a milestone then has

a quantitative and statistical basis that can be objectively

evaluated. For example, milestones derive easily from

parameters associated with trends in Table 6-1: 

Milestone 1: the year(s) when the negative ozone trends

seen in the 1980s and 1990s begin to weaken (related

to Stage (i))

Milestone 2: the year(s) that there is a trend reversal

(related to Stage (ii))

Milestone 3: the year(s) that EESC returns to pre-1980

values 

Milestone 4: the year(s) that the parameter reaches pre1980 values (related to Stage (iii))

Clearly, many milestones can be defined using the

parameters listed and the expectations of the stages of

ozone recovery as outlined above and as temporally displayed in Figure 6-1. Of greatest interest and importance

are those milestones that primarily result from changes in

ODS amounts, that will occur soonest, and that will have

the largest statistical significance using available observational datasets. Also of interest are those milestones that

can be simulated in global models of ozone recovery using

known changes in ODSs.

6.2.4 Using Data and Models to Evaluate

Ozone Milestones and Milestone

Parameters

The principal tools to evaluate ozone milestone

parameters are statistical analyses of ozone observations,

the results of global models for past ozone amounts, and

the model projections of future amounts. Observational

time series are available from multiple sources, including

satellite-, ground-, aircraft-, and balloon-based instruments. Each has a role in documenting ozone parameters

over time in profile and column amounts, and each has

value in establishing accurate ozone trends. Statistical

tools are required to derive ozone trends because ozone

amounts are subject to significant natural variability

throughout the available time series and the quality of

ozone data is not uniform. The variability arises because

ozone amounts reflect transport as well as chemical production and loss processes that are affected by a wide range

of factors as noted in Sections 6.2.1 and 6.3, and Section

3.4 of Chapter 3.

Models and statistical methods are key tools in the

attribution of changes in milestone parameters to the various controlling factors. Statistical analyses of time series

are used to derive trends and their uncertainties (see

Chapter 3). Photochemical box models constrained by

observations of chemical composition are useful for quantifying the role of ODS changes in observed parameter

changes. More complex models are needed to address

global ozone changes caused by the full range of contributing factors; namely, changes in atmospheric composition, stratospheric temperatures, atmospheric transport,

solar output, and volcanic emissions. Those relied upon

in this Assessment are 2-D models and CCMs. Descriptions of these models are provided in Chapter 5 while the

role of each model in this chapter is detailed in the following sections. An important challenge for these models

is to represent the atmospheric processes sufficiently well

that observed changes can be understood and attribution

of changes can be discussed. A large source of uncertainty

in model projections of ozone is the scenarios that must

be adopted by the models to account for changes in atmospheric parameters related to climate change.

6.3 FACTORS AFFECTING THE DETECTION,

ATTRIBUTION, AND TIMING OF

MILESTONES

In this section we discuss to what extent different

factors contributing to the variability in stratospheric

ozone are likely to affect the detection, attribution, and

timing of milestones. A major issue is the variability in

ozone induced by these factors that masks or resembles

the expected ozone change due to halogen loading (see

Figure 3-1 of Chapter 3). For projections over the rest of

the century, the issue is how factors other than ODSs may

change the expected increases in ozone from decreasing

halogen loading.

6.3.1 Stratospheric Halogen Loading 

The evolution of stratospheric halogen loading is

an obvious factor impacting ozone. Correctly estimating

stratospheric halogen loading is important for attributing

observed changes in ozone to decreases in EESC. The

global mean EESC from WMO (2003) is generally used

in such studies. However, if the estimated EESC is incorrect or not appropriate for the region being considered,

then the attribution or inferred timing of a milestone may

be incorrect. Possible causes of incorrect EESC include

using an inappropriate mean age of air, neglecting mixing

processes in the atmosphere (i.e., neglect of age spectra),

or incorrectly accounting for the bromine contribution (see

Chapter 1). The effect of errors in EESC estimation is

most likely largest for the detection of milestones in polar

regions, where the EESC from WMO (2003) peaks too

21st CENTURY OZONE LAYER

6.8
early and decays too rapidly because of the older mean

age of air in polar regions and the neglect of the age spectra

(see Newman et al., 2006, and the figure in Box 8-1 of

Chapter 8).

Future stratospheric halogen concentrations will

depend on future emissions of ODSs and on transport into

and through the stratosphere. Model simulations suggest

that increases in WMGHGs (carbon dioxide (CO2), nitrous

oxide (N2O), and methane (CH4)) may lead to an increased

stratospheric circulation and to reduced transport time

scales (e.g., Butchart and Scaife, 2001; Butchart et al.,

2006; Section 5.3.3 of Chapter 5). A reduction in transport time scales will result in a reduction of EESC, which

could affect the timing of long-term milestones.

6.3.2 Atmospheric Chemical Composition

Apart from changing ODSs, changes in other gases

could affect the evolution of ozone and the timing of ozone

recovery by changing the background chemical composition of the atmosphere. In particular, increases in gases

producing radicals that catalytically destroy ozone (e.g.,

N2O, CH4, molecular hydrogen (H2), and water (H2O)) are

likely to change ozone. As discussed in Section 6.3.3, temperature changes due to WMGHGs are also important.

Section 1.4 of Chapter 1 discusses the possible long-term

changes in gases other than ODSs that may affect ozone

and Section 6.6 discusses model estimates of ozone evolution based on emission scenarios of these gases.

Catalytic ozone loss in the stratosphere occurs from

the reactive nitrogen (NOx), hydrogen (HOx), oxygen (Ox),

chlorine (ClOx), and bromine (BrOx) families. Ozone loss

through these families is strongly altitude and latitude

dependent (see Figure 1.11 of IPCC/TEAP, 2005), with

NOx dominating in the middle stratosphere (approximately

25-40 km), and HOx dominating in the lower and upper

stratosphere. Under conditions of high chlorine loading,

ClOx is important in the upper stratosphere (peak impact

near 40 km) and in regions where heterogeneous reaction

rates are large, such as in the polar regions during spring.

However, in many regions of the stratosphere, these

changes are strongly buffered by induced changes on the

other chemical families, which can reduce the primary

impact (Nevison et al., 1999). For example, NOx increases

in the lower stratosphere cause decreases in HOx and ClOx

catalyzed losses, along with increases in “tropospheric”

ozone production mechanisms. In the middle stratosphere,

NOx induced changes are reduced by interactions with

chlorine species. On the other hand, in some cases coupling between different chemical processes can amplify

the effects of source gas emissions, e.g., nitrogen dioxide

(NO2) concentrations over southern midlatitudes have

risen at approximately twice the rate of its source gas N2O

as a result of changes in ozone (Section 3.3.2 of Chapter

3; McLinden et al., 2001).

Another compositional change that could affect

ozone is a change in stratospheric water vapor. An increase

in water vapor would increase HOx and thus cause ozone

decreases in the upper and lower stratosphere (KirkDavidoff et al., 1999; Dvortsov and Solomon, 2001),

although these changes are reduced by complex buffering

interactions. In the polar regions, increases in water vapor

would cause an increase in heterogeneous reaction rates

(e.g., hydrogen chloride (HCl) plus chlorine nitrate

(ClONO2)) and an increase in the surface areas of polar

stratospheric cloud (PSC) particles. Both effects are likely

to lead to an increase in chlorine activation and ozone loss.

The effects of water vapor increases on ozone (via HOx),

induced by increases in methane, are partially offset by

the reaction of methane with atomic chlorine, which deactivates ClOx and reduces ClOx-driven ozone loss (this

could be important throughout the stratosphere). The coupling of water vapor and methane with ClOx induced ozone

loss will be eliminated by decreasing ODS levels during

the 21st century.

While the Assessment does not deal with tropospheric ozone per se, it should be noted that increases in

tropospheric ozone in some regions might have masked

stratospheric decreases (such as in the tropics, where stratospheric decreases are relatively small). In such regions,

total column ozone measurements cannot be used for

detection and attribution of recovery milestones, and vertical ozone profile data are required. Vertically resolved

observations in the stratosphere allow separation of

regions controlled by ODSs and those dominated by transport (see Section 6.5). 

6.3.3 Stratospheric Temperatures

As discussed in Chapter 5, the rates of chemical

reactions, and the formation of PSCs, depend on temperature. Thus, changes in temperature can have a large influence on ozone. Temperature changes need to be accounted

for when attributing observed ozone variations to changes

in halogen loading and when predicting future ozone

levels. This is especially important for attribution in polar

regions, where interannual variations in ozone are closely

coupled to variations in polar temperatures; in the Arctic,

ozone loss rates are related to the volume of PSCs (e.g.,

Rex et al., 2004), whereas in the Antarctic, the size of the

ozone hole depends on the temperature in the vortex

“collar region” (60°S-70°S) (e.g., Newman et al., 2004). 

Future stratospheric temperatures are a major

source of uncertainty when predicting future ozone levels.

21st CENTURY OZONE LAYER

6.9
They are a key parameter for the questions of whether and

when ozone will return to pre-1980 levels. As discussed

in Chapter 5, stratospheric temperatures depend on stratospheric dynamics, radiation, and composition. Future

changes in temperature, and hence ozone, are not likely to

be uniform throughout the stratosphere. Cooling due to

increased CO2 (and other WMGHGs), particularly in the

upper stratosphere, is expected to slow gas-phase ozone

loss reactions. When EESC decreases to pre-1980 levels,

and if there are no other changes, the cooling will lead to

an increase in ozone to values higher than in 1980.

However, as noted in the previous subsection, increases in

WMGHGs will also alter the chemical composition of the

stratosphere and possibly the Brewer-Dobson circulation.

These effects are also likely to affect ozone; see Section

6.6 for further discussion. The impact of stratospheric

cooling on ozone might be the opposite in polar regions.

Here, cooling could result in increases in PSCs, which,

given enough halogens, would increase ozone loss.

6.3.4 Atmospheric Transport

Atmospheric transport is a major factor contributing

to stratospheric ozone variability. Accounting for this variability is an important issue both for detecting and

attributing ozone recovery milestones. As discussed in

Chapters 3 and 4, changes in the stratospheric meridional

circulation, in the stratospheric polar vortices, and in tropospheric weather systems, all have a strong influence on

stratospheric ozone and can produce variability on a wide

range of time scales. Some of these changes can be linked

to waves propagating from the troposphere, but internal

stratospheric dynamics also play a role.

The detection of recovery milestones is essentially

a signal-to-noise ratio problem, and much of the “noise”

in the ozone signal results from variability in transport.

Geographical differences in the variability in ozone

induced by this factor will affect where and when certain

recovery milestones can best be detected. Alternatively, it

determines the number of years of measurements required

to detect a milestone. Reinsel et al. (2002) found that for

midlatitudes (30°-60°), a statistically significant change in

total column ozone trend can be detected with ~7-8 years

of data following the period of linear decline, whereas

detection of ozone turnaround (the second stage in the

ozone recovery process) can require ~15-20 years for

southern midlatitude zonal average column ozone and

~20-25 years for northern midlatitude zonal average

column ozone. This result is in agreement with earlier

findings (Weatherhead et al., 2000). Analyzing vertically

resolved atmospheric regions is an effective technique to

increase the signal-to-noise ratio because of the vertical

separation of the forcing functions (Yang et al., 2006;

Weatherhead and Andersen, 2006).

Statistical regression models are often used to

remove competing drivers of ozone variability in the attribution of ozone changes to ODSs. They are based on the

assumption that proxies describing the dynamical state of

the atmosphere and resultant effects on ozone can be provided as model basis functions. Commonly used dynamical proxies are equatorial zonal winds to capture the

effects of the quasi-biennial oscillation (QBO), the Arctic

Oscillation or Antarctic Oscillation (AO or AAO), the

North Atlantic Oscillation (NAO), the El Niño Southern

Oscillation (ENSO) index, and measures of wave activity

(e.g., latitudinally averaged tropopause Eliassen-Palm

fluxes). However, because such preselected proxies are

not necessarily independent (they are not orthogonal

within the regression model), the partitioning of the ozone

variance among the different proxies by the model requires

some interpretation. It is difficult to establish how much

of the transport-driven ozone variability is appropriately

accounted for. An alternative approach is to use idealized

modeling studies to quantify the impact of different circulation regimes on ozone. For example, the magnitude of

ozone changes related to a strong warm ENSO event was

investigated by Brönnimann et al. (2004) and Pyle et al.

(2005).

Changes in temperatures and transport not only

complicate the detection and attribution of recovery milestones, they also affect ozone projections over the rest of

this century (see Chapter 5).

6.3.5 The Solar Cycle

When attributing recent changes in stratospheric

ozone to changes in ODSs, it is important to consider

ozone variations related to the 11-year solar cycle because

the timing of the most recent maximum in solar activity,

between 1999 and 2003, was around the time when EESC

peaked in the stratosphere. As discussed in Chapter 3,

observations continue to indicate a statistically significant

solar variation of ozone, with ozone in phase with solar

activity. This suggests that an increase in solar activity

during the 1999-2003 solar maximum will have contributed to the slowing of the decline and increase of ozone

(e.g., Dameris et al., 2006). Proper attribution of the cause

of the ozone changes in recent years requires the separation of ozone increases due to changes in solar irradiance

from those due to changes in halogen levels. However,

this is difficult, as the magnitude of the solar influence on

ozone is somewhat uncertain.

The amplitude of ozone changes due to solar

activity varies with altitude and latitude. In the upper strat21st CENTURY OZONE LAYER

6.10
osphere, ozone during solar maximum is 2 to 5% higher

than in solar minimum, with an uncertainty around 2%

(McCormack and Hood, 1996; Steinbrecht et al., 2004a;

Figure 3-19 of Chapter 3). Sensitivity studies by Cunnold

et al. (2004) indicate that current estimates of the solar

cycle effect on ozone are probably sufficiently accurate to

allow the separation of halogen decrease-related ozone

increases from solar cycle effects in the upper stratosphere.

Recent data from 2004 and 2005, i.e., from the beginning

of the solar minimum, confirm this result (Steinbrecht et

al., 2006a).

The situation is less clear for total column ozone.

Depending on latitude and location, total column ozone

is between 2 and 10 Dobson units (DU) higher during

solar maximum, both in observations and model simulations, with uncertainty ranging from 2 to over 5 DU

(McCormack et al., 1997; Steinbrecht et al., 2006b;

Reinsel et al., 2005). One reason for this large uncertainty

in the magnitude of the solar cycle variation in total ozone

is the fact that the two solar maxima before 1999-2003

coincided with large volcanic eruptions. It is difficult to

separate the impacts of eruptions and solar cycle on

observed ozone (Solomon et al., 1996). There were no

major volcanic eruptions during the 1999-2003 solar maximum. However, in this case, the ozone response is contaminated by the ozone changes related to the turnaround

of ODSs, that we are trying to quantify. It is further

unlikely that the solar cycle signal in ozone exactly follows the simple proxies used in most analyses, such as the

Mg II core-to-wing ratio, 10.7 cm radio flux, or an 11-year

harmonic function (Steinbrecht et al., 2004a,b). The

response may inherently vary from one solar maximum to

the next (Ruzmaikin et al., 2003). All of the above factors

add uncertainty to estimates of ozone increases in all three

past solar maxima. In particular, they complicate the separation of recent or near-term increases in ozone due to

the EESC turnaround from increases due to the 1999 to

2003 solar maximum. It is likely that observations at least

to the end of the next solar minimum in 2008 will be

required to allow better separation of solar cycle effects

from possible ozone increases due to decreases in EESC.

6.3.6 Volcanic Eruptions

As discussed in the previous Ozone Assessment

(WMO, 2003) and in Chapter 5, volcanic eruptions can

have a large impact on stratospheric ozone by changing

heterogeneous chemistry, thermal structure, and circulation in the stratosphere. Because of this, it is necessary to

consider volcanic eruptions both when interpreting

observed changes and when making projections of future

changes of ozone.

There have been no large volcanic eruptions since

the 1991 Mt. Pinatubo eruption, and the stratospheric

aerosol loading in the four years since the previous Ozone

Assessment has remained at low, nonvolcanic levels.

However, the impact of the Mt. Pinatubo eruption still

needs to be considered when attributing changes in ozone

in the last decade of the 20th century to changes in ODSs

(or any other factor). The Mt. Pinatubo eruption contributed to a large decline in Northern Hemisphere ozone,

which was followed by an increase in ozone as stratospheric aerosols decayed back to low, nonvolcanic levels.

This decrease in aerosol levels occurred at around the same

time that the growth in EESC slowed and reached its peak

value. Both the changes in aerosols and EESC led to

changes in ozone levels, and it is difficult to separate their

impact on ozone.

If one or more large volcanoes erupt in the next 50

years, it is likely to impact the ozone recovery process.

The overall impact of volcanic eruptions varies with

halogen levels (see Chapter 5). Outside the polar regions,

the primary effect of an increased rate of heterogeneous

reactions is to cause a reduction of nitrogen oxides. In the

current high-chlorine conditions, this causes an increase

in reactive chlorine and increased ozone depletion, as

observed following the Mt. Pinatubo eruption (Brasseur

and Granier, 1992). However, in low-chlorine conditions,

a large volcanic eruption could cause a small ozone

increase due to the suppression of nitrogen oxides (e.g.,

Tie and Brasseur, 1995). Hence, a large, Pinatubo-like

eruption within the next 20 years, when there will still be

significant amounts of halogens in the stratosphere, may

lead to an increase in ozone destruction by ODSs and a

temporary delay in ozone recovery; whereas a similar

eruption in the more distant future, when EESC has

decreased to around or below 1980 values, may lead to an

increase in ozone levels. In both cases, the exact impact

will depend on the latitude and size of the volcanic eruption, and ozone may also be impacted by changes in the

stratospheric thermal structure and circulation caused by

the eruption.

6.4 STATISTICAL METHODS FOR DETECTION

OF MILESTONES

Detection of recovery milestones, and in particular

the detection of the onset of ozone recovery (i.e., a slowing

of the ozone decline), cannot be achieved simply by examining raw measurement time series. Statistical methods

are required to distill the subtle changes in trend from an

observed ozone time series. Typically, multiple linear

regression fits are used with various proxies accounting

for different contributions to the ozone variation, which

21st CENTURY OZONE LAYER

6.11
have been described in Section 6.3. Autocorrelation in the

residual time series also needs to be considered (Tiao et

al., 1990). Several different approaches have been used

on model output, total column data, and vertically resolved

data, to test the last few years of data for an increase in

values: the change in linear trend (“change-in-trend”)

method (Reinsel et al., 2002), the multivariate adaptive

regression splines method (Krzy ́scin et al., 2005), the flexible tendency method (Harris et al., 2001), or the cumulative sum of residuals (CUSUM) method (Reinsel, 2002;

Newchurch et al., 2003). Each has its merits, offers

slightly different information, and tests different hypotheses with the data. The change-in-trend and CUSUM

approaches have been used extensively to examine recent

data and test for early signs of changes in ozone trends.

They are discussed in more detail below. Both of these

methods assume linear trends in ozone between the start

of the analysis period (usually 1980) and a selected “turnaround date” (usually 1996). Over this period, when EESC

increased approximately linearly, the expectation is, therefore, that ozone loss is linearly proportional to EESC. This

assumption has been shown to be valid over midlatitudes

(Yang et al., 2006), over the Arctic (Douglass et al., 2006),

and over the Antarctic when ozone destruction is incomplete (Jiang et al., 1996).

6.4.1 Change in Linear Trends

In the change-in-trend method (Reinsel et al.,

2002), the standard least-squares regression model (see

Chapter 3) is extended by adding a second basis function,

set to zero before a selected turnaround date and proportional to time thereafter (Figure 6-2). The turnaround date

is preselected and its choice can affect the results. Such a

regression model gives an estimate of the magnitude of

the trend before the turnaround date (m1), the slope change

at the turnaround date (m2), and the magnitude of the trend

after the specified turnaround date (m1+m2). The statistical significance of m2 can be tested and a change-intrend milestone is reached when m2 is significantly different from zero at the 2σ level. This method assumes

that the ozone record exhibits a discrete change from a

downward trend to a less steep trend or an upward trend

across the turnaround date (Reinsel et al., 2002).

Calculations indicate that the actual shift from upward

trend to downward trend in EESC may take a few years,

but includes a near linear trend prior to the turnaround

and a near linear trend for a few years after this turnaround

(WMO, 2003). If the underlying trend in ozone follows

this EESC pattern, the piecewise linear and contiguous

approach may be appropriate. Each data point, in most

cases monthly and regionally averaged data, is weighted

by the square of the distance from the fit, resulting in large

deviations having a strong influence on the final trend

estimates. Anomalous years either near the turnaround or

near the beginning or end of the record have particularly

high leverage on the trend estimates.

6.4.2 Cumulative Sum of Residuals

The CUSUM approach (Reinsel, 2002; Newchurch

et al., 2003) assumes that the data, after accounting for

various non-ODS influences, follow a near-linear downward trend until some preselected specified date, analogous to the turnaround date of the change-in-trend method.

After that date, all later data are evaluated to determine if

they represent a likely deviation from that downward

trend. This evaluation is made by examining the cumulative sum of residual deviations from the extrapolated trend

after the specified point (see Figure 6-2). The choice of

the turnaround date is not very important provided it is not

too far from the physically appropriate point (e.g., the date

at which EESC maximizes). In contrast to the change-in21st CENTURY OZONE LAYER

6.12

time 

ozone deviation [%]

time

cumulative sum [%]

trend m

trend after m +m2

summed

residuals

CUSUM

c hange of trend m2

CUSUM of data

CUSUM of trendline

( m2 2/ 2)

t urnaround date

(a)

(b)

~~

Figure 6-2. A schematic representation of (a) the

change-in-trend method, and (b) the CUSUM

method. See Section 6.4 for discussion.
trend approach, CUSUM makes no assumptions about the

temporal path of the deviation. Data shifts, gradual reductions in trend, and slow changes to an extrapolated trend

are all detected. However, little information is provided

on the temporal shape of a change. Because the cumulative sum of deviations is considered, as opposed to the

minimum of the sum of the squares as in the change-intrend method, highly unusual points are somewhat less

influential. The explanatory variables, such as the QBO,

solar cycle, and an underlying trend, are fit for the entire

time series and removed from the data to create the

monthly ozone residuals. These residuals are used to form

the trend estimate calculated over the period up to the turnaround date and extrapolated thereafter for the CUSUM

calculation.

The CUSUM approach detects arbitrary deviations

from the extrapolated trend, whereas the change-in-trend

method looks for a specific temporal path. Comparison

between the change-in-trend and CUSUM approaches is

possible using the fact that if there is a change of trend by

m2 per unit time, the CUSUM value after the turnaround

increases with time approximately as m2 t 2 /2 (Yang et al.,

2006). For example, a representative change of trend m2

for total ozone at northern midlatitudes is 2 DU per year,

or 0.166 DU per month (see Figure 6-4, panel b). When

accumulated over 84 months from January 1997 to

December 2004, this change of trend corresponds to a

CUSUM value of 595 cumulative DU by the end of 2004.

This compares very well with the CUSUMs given in

panels (d) and (f) of Figure 6-3 for partial ozone columns.

By the end of 2004, these CUSUMs amount to 250 and

350 DU, or 600 DU combined for the total column.

6.5 ATTRIBUTION OF THE RECENT BEHAVIOR

OF OZONE

To attribute the recent changes in ozone to changes

in ODSs, we require more than the observation of a

slowing of the ozone decline or an increase in ozone. We

require that the observed changes in ozone occur at approximately the right time to be associated with ODS concentrations, that the magnitude of change in ozone trend be of

the appropriate magnitude expected due to changes in ODS

concentrations, and that the latitudinal, altitudinal, and

seasonal changes are in agreement with the changes

expected due to a turnaround in ODS concentrations. Even

if all of the above are observed, scientific judgment is

involved in assessing whether the recent changes in ozone

are appropriately linked with the changes in ODS concentrations. Scientific judgment in this assessment process

takes into account statistical studies, a variety of models

involving a range of appropriate assumptions, an understanding of the time scales of relevance in the atmosphere,

possible uncertainties in mechanisms governing ozone

concentrations, as well as uncertainties in available ozone

measurements.

Attribution is fundamentally difficult because, as

discussed in Section 6.3, many factors other than ODSs

can affect ozone on time scales of a few years. The separation of these other factors from any underlying ODS

signal in ozone is further complicated by the fact that the

quantification of the many factors affecting ozone is uncertain, can be nonlinear, and involves feedbacks through a

variety of mechanisms. Furthermore, attribution requires

analyzing data with high natural variability and using

models that cannot represent atmospheric processes at

arbitrary degrees of spatial and temporal resolution.

Several approaches have been used to assess whether the

changes in ozone can be attributed to changes in ODSs.

These include using statistical methods to separate effects

of transport and temperature from chemical effects of

ODSs, and modeling efforts to partition the past changes

into those due to ODS changes and those due to known

transport and temperature effects.

6.5.1 Upper Stratospheric Ozone

As described in Section 6.2, the first stage of the

ozone recovery process is defined as a statistically significant deviation above the previous linear decline in ozone

that must be attributable to changes in EESC. We expect

that this stage of ozone recovery will be passed first in

those regions of the atmosphere where ozone changes are

most closely controlled by changes in EESC. One such

region is the upper stratosphere, where the largest ozone

decline due to gas-phase chemical reactions has been

recorded (WMO, 2003) and where few factors other than

gas-phase chemistry directly control ozone. Reinsel

(2002) and Newchurch et al. (2003) showed that a statistically significant deviation in the ozone decline can be

found in observations of ozone in this altitude region. This

deviation is illustrated in Figure 6-3(a), which shows

ozone residuals for northern midlatitudes and the altitude

range from 35 to 45 km. Residuals were defined by subtracting the mean annual cycle and estimated variations

due to the solar cycle and QBO. Up to about the end of

1996, ozone residuals closely followed a linear decline of

approximately −7% per decade, and since then, this steep

decline has not continued and ozone levels have been

essentially constant or may have even increased. This

behavior is quantified in Figure 6-3(b), where the

CUSUMs from the extrapolated 1979 to 1996 trend are

shown. The CUSUM increases with time since 1996,

demonstrating that the residuals lie significantly above the

21st CENTURY OZONE LAYER

6.13
21st CENTURY OZONE LAYER

6.14

80 85 90 95 00 05

-10

0

10

20

Residual ozone (%)

SAGE ozone at 35-45 km, 30-60N

L(ClOx/BrOx)/L(total) fit

EESC fit

979-1 96 trend = -7.6 +/- 1.1 %/decade

(a)

80 85 90 95 00 05

-200

-100

0

00

00

300

400

CUSUM (%)

SAGE

(ClOx/BrOx)/L(total) f

EESC f

(b)

80 85 90 95 00 05

-10

-5

0

5

10

15

Residual ozone (DU)

SAGE/HALOE ozone at 18-25 km, 30-60N

(ClOx/BrOx)/L(total) fit

ESC fit

979-1 96 trend = -4.4 +/- 1.2 DU/decade

(c)

80 85 90 95 00 05

-200

-100

0

00

00

300

400

CUSUM (DU)

SAGE/HALOE

(ClOx/BrOx)/L(total) f

ESC f

(d)

80 85 90 95 00 05

Year

-20

-10

0

10

20

Residual ozone (DU)

ozonesonde ozone at TP-18 km, 30-60N

EESC fit

979-1 96 trend = -2.1 +/- 1.6 DU/decade

(e)

80 85 90 95 00 05

Year

-200

-100

0

00

00

300

400

CUSUM (DU)

ozonesonde

EESC f

(f)

Figure 6-3. Time series of monthly average ozone residuals plus linear trend (left panels) and cumulative sum

(CUSUM) of residuals (right panels) in percent or DU. Panels on the left show ozone column residuals between

30°N and 60°N, at 35-45 km from SAGE (top), at 18-25 km from SAGE and HALOE (middle), and from the

tropopause (TP) to 18 km from ozonesondes (bottom). Ozone residuals, calculated over the period 1979-

2005, have annual cycle, QBO, and solar cycle effects removed. In all left panels, EESC fits to the residuals

are shown in red. The solid black lines on the left panels indicate the ozone trend calculated from observations for 1979-1996 and forecasted linearly afterward (dotted). The blue lines in the top and middle panels

show the ozone evolution expected from photochemical model calculations. Cumulative sums of residuals of

traces in the left hand panels are shown in the panels on the right together with the 95% confidence envelopes

of departure from natural variability and model uncertainty as dotted lines. Updated from Newchurch et al.

(2003) and Yang et al. (2006).

extrapolated trend line, and since about 1999 the CUSUM

has exceeded the 2-sigma error envelope (95% confidence

level) of a continuing decline. Similar behavior is found

for equatorial latitudes and Southern Hemisphere midlatitudes (Newchurch et al., 2003).

The significant positive deviation of upperstratospheric ozone levels from the previous linear

decline shown in Figure 6-3(a,b) is corroborated by

Petropavlovskikh et al. (2005) using Umkehr and

SBUV(/2) data, and by Steinbrecht et al. (2006a) using
ground-based and satellite records for four stations

between 45°S and 44°N and the change-in-trend method.

The latter showed an increase in ozone residuals over the

last few years for some of the stations (see also Figure 3-6

of Chapter 3). The lessening of the trends in the residuals

depends to some degree on the estimates of the solar-cycle

and QBO amplitude in ozone (Steinbrecht et al., 2004a;

Cunnold et al., 2004), as discussed in Section 6.3. Also,

the reduction in trend in upper stratospheric ozone was not

significant at the most northerly station (48°N) considered

by Steinbrecht et al. (2006a), see Figure 3-6. At northern

midlatitudes, other unquantified factors seem to mask the

reduction of the ozone decline expected from the beginning of the decline of EESC. A general problem is the lack

of long-term upper stratospheric ozone records at latitudes

north of 60°N. Nevertheless there is now conclusive evidence for a reduction in the negative trend in upper-stratospheric ozone over a wide region between southern and

northern midlatitudes.

The observations of a slowing of the ozone decline

discussed above do not alone constitute the first stage of

recovery as these changes need to be attributed to changes

in ODS amounts. While the ozone changes are in general

consistent with those expected from changes in EESC, we

recognize that we do not have a full quantitative understanding of ozone loss in the upper stratosphere (Chapter

3). The smooth curves in Figure 6-3(a,b) show fits to the

ozone residuals using the EESC (red curve) and ozone loss

fraction due to chlorine and bromine (blue curves). The

latter was calculated using a photochemical model constrained by observations of total reactive nitrogen (NOy),

inorganic chlorine (Cly), CH4, and H2O (Yang et al., 2006).

As quantified by the CUSUMs, these fits show a significant deviation from the previous decline up to 1997. The

good fit to the observed ozone residuals by 1) EESC and

2) the calculated ozone loss fraction due to EESC, as well

as the fact that most other sources of ozone variability have

been removed, provides strong evidence that the slowing

of the ozone decline in the midlatitudes in the upper stratosphere can be attributed to changes in ODSs.

6.5.2 Lower Stratospheric and Total

Column Ozone

A similar slowing of the decline since 1997 has also

been observed in total column ozone, and in some regions

of the globe, column ozone has even increased since 1997

(Chapter 3). This deviation from the previous decline is,

in general, statistically significant (e.g., Reinsel et al.,

2005; Krzy ́scin et al., 2005; Dhomse et al., 2006; Miller et

al., 2006; Krzy ́scin, 2006; Yang et al., 2006).

The change in total column ozone evolution is illustrated in Figure 6-4, which shows calculations of the

overall trend and the change in trend (see Section 6.4) of

column ozone as a function of latitude (Reinsel et al.,

2005). The change-in-trend term is statistically significant at the 95% confidence level between 40°N and 60°N,

and between 50°S and 60°S. At lower latitudes the change

in trend is not statistically significant. The trends and

changes in trends shown in Figure 6-4 are representative

of the range of results obtained in Chapter 3 for various

total column ozone datasets. Furthermore, several other

studies (Krzy ́scin et al., 2005; Dhomse et al., 2006;

Krzy ́scin, 2006; Yang et al., 2006) came to a similar conclusion based, in part, on additional data and using somewhat different approaches. A statistically significant

slowing of the decline is also found for ozone in different

altitude regions. For example, Figures 6-3(c) and (e) show

residuals for ozone between 18 and 25 km and between

the tropopause and 18 km, respectively. The corresponding CUSUMs (Figures 6.3(d) and (f)) indicate a similarly

significant deviation from the previous decline.

The above studies have shown that there is now substantial evidence that a slowing of ozone decline has

occurred not only for upper stratospheric ozone (Section

6.5.1), but also for lower-stratospheric and total column

ozone over substantial parts of the globe. However, it is

still necessary to attribute these ozone deviations to

changes in EESC in order to label this as stage (i) of ozone

recovery. This attribution can be done by estimating

changes in chemical ozone destruction (Yang et al., 2006),

by eliminating as many non-chemical ozone variations as

possible (e.g., Hadjinicolaou et al., 2005; Reinsel et al.,

2005; Miller et al., 2006; Dhomse et al., 2006), or by comparisons with multidimensional models (e.g., Anderson et

al., 2006; Weatherhead and Andersen, 2006).

Yang et al. (2006) compared the changes in ozone

residuals for column ozone, ozone between 18 to 25 km,

and ozone between the tropopause and 18 km with changes

in EESC and those in the chemical ozone loss fraction due

to chlorine and bromine (using the same model as

described in Section 6.5.1). As shown in Figures 6-3(c)

and (d), the ozone residuals in the 18 to 25 km altitude

range closely follow both the EESC and the modelcalculated ozone loss fraction due to chlorine and bromine,

and CUSUM calculations indicate statistical significance

above the 95% confidence level. This result is robust for

both northern and southern midlatitudes. It is a strong

indication that the beginning of the reduction in EESC is

mirrored in photochemical ozone loss rates and that this is

the major contributor to the recent leveling off of ozone

residuals in the 18 to 25 km range (Yang et al., 2006). The

situation is however different for the lower altitude range

21st CENTURY OZONE LAYER

6.15
from the tropopause to 18 km (Figure 6-3(e) and (f)).

Although ozone residuals between the tropopause and 18

km have increased, the low cumulative sum of residuals

for the EESC fit (red curve in Figure 6-3(f)) suggests that

these changes have not been driven by declining EESC.

Because ozone is significantly controlled by transport in

this region, the low CUSUM for EESC indicates that

dynamical and transport changes likely account for the

major part of the ozone increases between the tropopause

and 18 km.

The altitude partitioning of total column ozone

changes before and after 1997, according to Yang et al.

(2006), is summarized in Table 6-2. About 80% of the

total ozone decline between 1979 and 1996 occurred above

18 km, and, as discussed above, this is largely due to

increasing EESC. The remaining 20% of the total ozone

decline occurred below 18 km, where transport changes

played a major role. For the recent total ozone increases,

however, only about half comes from altitudes above 18

km, and can be attributed to changes in EESC. The other

half comes from altitudes between the tropopause and 18

km, and must largely be due to dynamical and transport

changes. The error bars on this partitioning are substantial, of the order of ±30%; nonetheless, the CUSUM statistics are well above the 2σ level.

Numerous other studies, using a variety of

approaches, have also concluded that changes in transport make a major contribution to the recent increase of

total column ozone. Multilinear regression studies by

Dhomse et al. (2006), Reinsel et al. (2005), or Krzyscin  ́

(2006) show that dynamical factors account for a major

fraction, but not all, of the recent increases in total ozone.

Figure 6-4, for example, indicates that for northern lati21st CENTURY OZONE LAYER

6.16

Change in trend (DU/year) with 1σ mits

O veral l (turnaro und) trend (DU/year) with 1σ mits

a )

m1

b) 

m2 c) 

m1+m2

Lat itude 

Latitude 

Latitude 

2.5

-0.5

0.0

0.5

.0

.5

2.0

0

2

3

4

Pr e-turnaro und trend (DU/year) with 1σ mi ts

2.5

-2.0

.5

.0

-0.5

0.0

Figure 6-4. Trends in zonal mean (5°) total column

ozone from the merged Total Ozone Mapping

Spectrometer/Solar Backscatter Ultraviolet spectrometer (TOMS/SBUV) satellite dataset (see Chapter 3).

(a) Pre-turnaround trends calculated over the period

1979-1996; m1 as shown in Figure 6-2. (b) The

change in trend at the turnaround date, calculated

using data over the period 1979 to 2002; m2 as shown

in Figure 6-2. (c) the overall trend; m1 + m2 as shown

in Figure 6-2. All trends are shown with 1σ error bars.

Trends from a regression model incorporating basis

functions to account for changes in ozone driven by

the Arctic Oscillation, Antarctic Oscillation, and wave

activity are shown using filled circles, while trends calculated using a traditional regression model incorporating trend, QBO, and solar cycle terms are shown

using open circles. From Reinsel et al. (2005).
tudes, trends before and after 1996 are comparable in

magnitude. Since EESC is declining at a rate three times

slower than the previous increase, this is a strong indication for a major contribution from transport to the postturnaround trend, at least in the Northern Hemisphere.

Chemical transport model simulations using the European

Centre for Medium-Range Weather Forecasts ERA 40

meteorological reanalyses with chlorine fixed at 1980

levels can reproduce the total ozone increase from 1996

to 2003 (Hadjinicolaou et al., 2005; see also Chapter 3).

Analysis of Canadian ozonesondes indicates significant

increases in ozone below 60 hPa (20 km), mostly related

to dynamical changes in the occurrence of ozone laminae

(Tarasick et al., 2005). Furthermore, studies suggest that

changes in the mean meridional Brewer-Dobson circulation, e.g., inferred by 100 hPa Eliassen Palm flux, and in

lower tropospheric wave forcing, e.g., described by potential vorticity, can account for a substantial fraction of the

pre-1997 total ozone decline outside the polar regions and

also for the recent increases in northern midlatitude spring

(e.g., Salby and Callaghan, 2004; Malanca et al., 2005;

Hood and Soukharev, 2005; see also Section 3.4.2 of

Chapter 3).

Andersen et al. (2006) compared the recent changes

in total column ozone with fourteen 2-D and 3-D model

estimates of recovery rates. The comparison showed that

for most areas, the changes in long-term trends were in

agreement with the range of model expectations for the

long-term trends. However, the change in trend was larger

than any of the models predicted for the same short time

period (1996-2003). In particular, they showed that the

large changes observed in the high northern latitudes were

larger than any of the models predicted, indicating that

much of the change observed north of 50°N may be natural variability and such trends will not continue into the

future at the same rate. Weatherhead and Andersen (2006)

extended this analysis and showed that the latitudinal, altitudinal, and seasonal signatures of change in ozone from

1996 through 2005 are in rough agreement with what

models are predicting for long-term ozone change

(between 1996 and 2050 for the 2-D models and for

periods of time-slice runs for 3-D models); see Figure 6-5.

The models and measurements agree well in the Southern

Hemisphere, and in the Northern Hemisphere show rough

agreement for past trends but significant disagreement for

the magnitude of trends since 1996. Some of this disagreement could be due to the different time periods considered

for observations and models.

All of the above approaches suffer from methodological and statistical uncertainties. Multilinear regression

methods are limited by the simplicity of the underlying

assumptions, and are only an attempt to describe very complex processes in the atmosphere with a simple set of

indices. Not all effects, e.g., solar cycle effects, may be correctly removed on the basis of these simple indices. The

reconstruction of chemical ozone loss may require substantial data extrapolation to include periods before measurements of key species (e.g., methane) were available (Yang

et al., 2006). Also, bromine may be more important in the

lower stratosphere, and lead to higher ozone loss than previously thought (Salawitch et al., 2005). Estimations on the

basis of chemical transport models and reanalysis datasets

(National Centers for Environmental Prediction (NCEP) or

ERA-40), e.g., Hadjinicolaou et al. (2005), suffer from

uncertainty regarding the long-term consistency of the

reanalysis datasets and their ability to correctly describe the

slow mean meridional circulation (see Chapter 3). Also,

the meteorological data may inherently contain effects from

21st CENTURY OZONE LAYER

6.17

Table 6-2. Total column ozone trends and trend changes, and contributions from different altitude

ranges. The column ozone is obtained from the Dobson/Brewer (total), Stratospheric Aerosol and Gas

Experiment (SAGE) satellite (above 25 km), SAGE (18 to 25 km), and ozonesonde (tropopause (TP) to 18 km)

measurements. The fractions with respect to the ozone column above the tropopause are listed in percent.

Results are for the 30° to 60°N latitude band. Uncertainties given are 2σ. Adapted from Yang et al. (2006).

Trend Fraction of Total Change of Trend Fraction of Change

Altitude Range 1979-1996 Trend 1979-1996 1997-2005 of Trend 1997-2005

(DU/decade) (%) (DU/decade) (%)

Total column −8.7 ± 2.3 17.4 ± 8.2

Above 25 km −4.3 ± 1.0 42 ± 10 4.9 ± 2.9 30 ± 18

18 to 25 km −3.9 ± 0.9 38 ± 9 3.4 ± 2.8 21 ± 17

TP to 18 km −2.1 ± 1.6 20 ± 16 8.0 ± 5.3 49 ± 3
the true ozone changes, making it difficult to separate chemical and transport processes.

Even though all approaches have uncertainties, the

complementary approaches come to similar conclusions:

that changes in transport have contributed a major fraction to the leveling off and increase in midlatitude total

ozone since 1997, particularly in the lower stratosphere

over northern midlatitudes. However the changes in

trends are in agreement, on a number of criteria, with

changes in EESC. It is likely that we have passed the

“slowing of ozone decline” stage for total ozone over large

parts of the globe and in the middle and upper stratosphere. However, at this time, due to the described statistical and methodological uncertainties, an unambiguous

attribution to changes in EESC, with a high level of confidence, is not possible.

6.5.3 Polar Ozone

The sustained increase in the severity of Antarctic

ozone depletion from the early 1980s to the later 1990s

has not continued. Recent changes in metrics of the

severity of Antarctic ozone depletion have ranged from

little or no change over the past 10 years (ozone hole area),

to small (minimum ozone levels) and moderate (ozone

mass deficit) signs of ozone increases (Bodeker et al.,

2005; Section 4.3.2 and Figure 4-8 of Chapter 4). The

cessation of ozone hole growth can be ascribed to the fact

that at current ODS levels, all or most of the ozone is

destroyed between 12 and 24 km (e.g., WMO, 2003).

Recent reductions in Antarctic ozone loss can be

explained by anomalously high stratospheric temperatures and reduced frequency of PSCs in recent years

(Solomon et al., 2005; Hoppel et al., 2005). Therefore,

although recent stabilization and reduction of Antarctic

ozone depletion has occurred at approximately the same

time as the growth in EESC has slowed and passed its

peak, this ozone change cannot be attributed solely to

changes in EESC. It is therefore not possible to conclude

that either the first or second stage of Antarctic ozone hole

recovery has occurred. The near total destruction of ozone

inside the ozone hole means that there is low sensitivity

to moderate reductions in EESC. ODS amounts in the

Antarctic vortex are expected to decrease only slowly

over the next decade (polar EESC is decreasing at around

0.6%/year, see figure in Box 8-1 of Chapter 8). Therefore,

in the near future, only small changes in Antarctic ozone

area are expected as EESC declines, and these changes

will be masked by interannual variability due to temperature and transport variations (Newman et al., 2004;

Solomon et al., 2005).

Larger sensitivity to changes in EESC is expected

at the upper altitudes of the ozone hole (20-22 km) where

ozone depletion is not complete, and this has been identified as a possible region to detect Antarctic ozone recovery

(Hofmann et al., 1997). Hoppel et al. (2005) analyzed

measurements of ozone mixing ratios in the 20-22 km altitude range for October, and showed that the values for

2001-2004 were higher than for 1994-1996 and 1998-

21st CENTURY OZONE LAYER

6.18

atitude

Trend (Dobson Units / year)

-60 -30 EQ 30 60

-4

2

0

2

4

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o o

o o o o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o o o o o o o o o o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

o

Figure 6-5. Measured and modeled

column ozone trends by latitude for

1979-1995 (in red) and 1996-2005 (in

blue). Increases in ozone during the

past eight years coincide with the

regions in which past depletion has

occurred. Analyses of the measurements are represented by the solid

bars; analyses of 2-D models are

shown with horizontal dashes, and

results of 3-D models are shown with

circles. The model estimates for 1979-

1995 are shown in red; projections for

long-term recovery are shown in blue.

Trends before 1995 are statistically

significant (2σ) for middle and high latitudes. Data are from version 8 of the

merged TOMS/SBUV2 dataset. Eq.,

equator; negative latitudes are

degrees south. From Weatherhead

and Andersen (2006). 
2000. However, the higher ozone mixing ratios were

accompanied by higher temperatures and reduced frequency of PSCs. Although ozone is more sensitive to ODS

changes in the 20-22 km region, there is more dynamical

activity in this region, and changes in ozone due to ODS

changes cannot be clearly distinguished from those due to

changes in meteorological conditions. It is therefore not

possible to conclude that the first (or second) stage of

ozone recovery has occurred at 20-22 km in the Antarctic.

Another region where detection of the first stages

of ozone recovery may be possible is the polar vortex

collar region (60°S to 70°S). Yang et al. (2005) examined

the evolution of ozone in this region using several different

datasets. They accounted for dynamical variability using

the observed correlation between ozone and 100 hPa temperatures, and showed that temperature-adjusted ozone

anomalies were roughly constant since 1997 (see Figure

6-6). The change in the trend of these anomalies is significant at greater than 95% confidence limits. Yang et al.

also showed that complete ozone loss is infrequent in the

collar region and was not responsible for the slowing of

the ozone decline. As shown in Figure 6-6, the variation

in the ozone anomalies is similar to that of midlatitude

EESC, and fits to EESC can explain most of the long-term

variations in the ozone anomalies. This high correlation

and the limited role of complete ozone loss provide evidence that the first stage of ozone recovery may have

occurred in the stratospheric collar region surrounding the

Antarctica vortex. However, there is some uncertainty in

the EESC in the collar region, and the resulting fit with

temperature-adjusted ozone anomalies. For example, the

polar EESC shown in Box 8-1 continues to increase until

21st CENTURY OZONE LAYER

6.19

80 85 90 95 00 05

-20

0

20

40

Ozone anomaly (%)

79-96 trend = -21.4 +/ 3.0 %/decade

-52.8 +/- 7.5 DU/decade)

EESC it to ozone

(a)

80 85 90 95 00 05

-100

0

100

200

300

CUSUM (%)

Dobson

total column

(b)

80 85 90 95 00 05

Year

-40

0

40

80

Ozone anomaly (%)

79-96 trend = -44.6 +/- 11. %/decade

-50.2 +/- 12.4 DU/decade)

EESC it to ozone

(c)

80 85 90 95 00 05

Year

-200

-100

0

100

200

300

400

CUSUM (%)

Ozonesonde

column a 12-25 km

(d)

Figure 6-6. September and October temperature-adjusted ozone monthly anomalies (left panels) and CUSUM

of ozone residuals (right panels) in percent for Dobson spectrophotometer total ozone columns at Vernadsky

(65°S) and Syowa (69°S) (top) and ozonesonde ozone columns from 12-25 km at Syowa (bottom). The ozone

anomalies are calculated from a regression model fit to the data over the full period. The temperature-adjusted

ozone anomalies are obtained by subtracting the ozone equivalent temperatures from the ozone anomalies

(see Yang et al., 2005, for details). The black line indicates the ozone trend calculated from observations for

1979-1996 and forecasted linearly afterward. Linear trends and 95% confidence intervals for 1979-1996 are

listed in %/decade and DU/decade. The red line shows the EESC fits to the temperature-adjusted ozone. The

blue lines indicate the 95% confidence envelopes of departure from natural variability and trend model uncertainty. Adapted from Yang et al. (2005).
2001, in contrast to the ozone anomalies in the collar that

peak around 1997. It is unclear whether the midlatitude

or polar EESC is more representative for the collar region

which is outside the ozone hole. These uncertainties cast

some doubt on using EESC to quantify the role of ODSs

in the stabilization of ozone in the collar region.

The expected slow improvement of Antarctic

ozone over the next decade (e.g., Newman et al., 2004,

2006) means that variability will continue to complicate

detection of the first and second stages of ozone hole

recovery, even after accounting for temperature variations. Solomon et al. (2005) have suggested that the return

of (1) the relationship between temperature and ozone

and (2) the variance in ozone abundances to historical

values may provide early signals of the beginning of

recovery inside the ozone hole. However, after a stage of

recovery has occurred, it is unclear how long it will take to

achieve the detection of the stage using either diagnostic.

The issue of identifying recovery in the Arctic

springtime is even more difficult than in the Antarctic.

Arctic ozone depletion occurs in most years, but there is

substantial interannual variability (see Chapter 4). The

larger meteorological variability in the Arctic compared

with the Antarctic (e.g., Langematz and Kunze, 2006) and

smaller ozone depletion means that it will likely take many

years to detect any changes in ozone due to decreases in

EESC. Furthermore, it is likely that the first stage of

recovery (slowing of decline) cannot be detected for the

Arctic. As with the Antarctic, the first signals of recovery

might be found by looking at the relationship between

ozone and temperature. As discussed in Chapter 4, one

relationship that reduces the variability is the compact relationship between vortex average ozone loss and VPSC, the

volume of air potentially containing PSCs (Rex et al.,

2004). It may be possible to detect the recovery of polar

ozone in the Northern Hemisphere from a statistically significant and persistent deviation from this relationship.

However, no such deviation has been detected, and there

has been no detection of any ozone recovery stages in the

Arctic.

6.6 PROJECTIONS OF THE FUTURE

BEHAVIOR OF OZONE

As discussed above (and in Chapter 5), the process

of recovery of the ozone layer will depend not only on the

decline of ODSs but also on many other factors. Although

some of these factors can be accounted for empirically

when projecting future ozone (e.g., Knudsen et al., 2004),

coupling between the different chemical, dynamical, and

radiative processes involved requires the use of models

that include these interdependencies to make well-founded

projections. In particular, models used for prognostic

studies should incorporate the effects of changes in temperature and transport that are likely to occur as the concentrations of WMGHGs rise. This section addresses how

changes in ODSs couple with other atmospheric changes

to influence the long-term evolution of ozone.

The framework introduced in Section 6.2 is used to

examine the ozone recovery process in the model simulations. In particular, projections of total column ozone are

examined for three periods: 

(i) The beginning of the century (2000-2020), when

EESC is expected to start to decrease or continue to

decrease

(ii) Mid-century (2040-2050 in extrapolar regions, 2060-

2070 in polar regions), when EESC is expected to fall

below 1980 values

(iii) End of the century (2090-2100), when factors other

than ODSs are expected to control stratospheric ozone

Confidence in projections near the beginning of the

century is higher than near the middle or end of the century because the former can be supported by observations,

empirical studies, and extrapolations while the latter are

more influenced by uncertainties in the emissions scenarios and other boundary conditions. In general, a separation of the different factors contributing to the ozone

variability in the model output has not been performed.

Therefore the modeled ozone time series presented below

cannot be used for attribution of ozone changes to changes

in ODSs.

6.6.1 Model Descriptions and Scenarios

In this Assessment both two-dimensional chemical

transport models (2-D models) and three-dimensional coupled Chemistry-Climate Models (CCMs) are used to make

projections of the ozone layer in the 21st century. By using

both classes of models, with their respective advantages

and disadvantages, the conclusions drawn from the model

projections are likely to be more robust.

2-D models have been used extensively in previous

Assessments. Their relative computational efficiency

allows long integrations and a large number of sensitivity

studies. This capability makes these models a valuable

tool for understanding and quantifying the long-term

changes in ozone. However, due to the inherent zonal

averaging within these models, they are not well suited

for modeling polar processes. The 2-D models used in this

Assessment (see Table 6-3 for a summary of their characteristics) vary in the extent to which they incorporate interactions between model components. Most of the 2-D

models use prescribed temperatures and transport, and

therefore do not include the well-known temperature feed21st CENTURY OZONE LAYER

6.20
back from changes in WMGHGs, or the ozone radiativedynamical feedback, in which changes in ozone also affect

the radiative balance of the stratosphere and hence temperatures and transport. Three of the models (NOCAR,

Leeds-Bremen, and GSFC-INT) calculate temperatures

from the modeled atmospheric composition, and capture

some of the feedback on the circulation and transport.

However, these so-called “interactive” models do not

include changes in the wave forcing.

CCMs have a much more limited history in the

Ozone Assessments. WMO (2003) was the first time

CCMs were fully integrated into an Assessment. These

simulations focused on the polar regions and, due to computational limitations, were largely restricted to the recent

past and near future (roughly 1980 to 2020). Transient

simulations (see Box 5-1 in Chapter 5) are preferred for

predicting future ozone because in these simulations,

ozone responds interactively to the gradual secular trends

in WMGHGs, ODSs, and other boundary conditions. For

this Assessment we consider only transient simulations

from CCMs, and we examine global as well as polar ozone.

Since WMO (2003), several new CCMs have been devel21st CENTURY OZONE LAYER

6.21

Table 6-3. A summary of the 2-D models used in this chapter.

Model Institution(s) Investigators Temperature PSC Solar Reference

Scheme Cycle

AER AER, Inc., U.S. D. Weisenstein Specified Based on No Rinsland et al.

from NCEP model T (2003)

analyses

GSFC- NASA/Goddard E. Fleming, Interactive Based on No Rosenfield et

INT Space Flight Center C. Jackman observed T al. (1997)

GSFC NASA/Goddard E. Fleming, Specified from Distribution No Fleming et al.

Space Flight Center C. Jackman UKMO analyses based on NCEP (1999)

Leeds- University of M. Chipperfield, Interactive Based on No Chipperfield

Bremen Leeds, U.K.; M. Sinnhuber, model T and Feng 

University of B.-M. Sinnhuber (2003)

Bremen, Germany

MNP 1 Netherlands G. Velders Climatology No solid No Velders

Environmental PSCs (1995)

Assessment Agency

MPIC MPI for Chemistry, C. Brühl Specified from Based on model No Grooß et al. 

Mainz, Germany CIRA data and observed T (1998)

deviation

NOCAR NOAA/NCAR R. Portmann Interactive Based on NCEP No Portmann et al.

T, Liquid aerosols (1999)

OSLO University of B. Rognerud Specified Based on No Stordal et al. 

Oslo, Norway climatology (1985);

Isaksen et al.

(1990)

SUNY SUNY, U.S.; S. Smyshlyaev Specified from Based on Yes Smyshlyaev

St. Petersburg, NCEP analyses NCEP T et al. (1998)

Russia

AER, Atmospheric and Environmental Research Inc.; MPIC, Max-Planck Institute for Chemistry; NOAA, National Oceanic and

Atmospheric Administration; NCAR, National Center for Atmospheric Research; SUNY, State University of New York; NCEP,

National Centers for Environmental Prediction; UKMO, United Kingdom Meteorological Office; CIRA, COSPAR (Committee

on Space Research) International Reference Atmosphere. 1 This was referred to as the RIVM model in the 2002 Ozone Assessment (WMO, 2003).
oped, significantly deepening the pool of model simulations of future ozone. 

An additional improvement over the approach

used in WMO (2003) is the use of two standard simulations, defined as part of the CCM Validation Activity

(CCMVal; Eyring et al., 2005). One “past” simulation

(REF1) is designed to reproduce ozone changes from

1980 to the present when global ozone observations are

available. It allows a detailed investigation of the role of

natural variability and other atmospheric changes important for ozone trends. All forcing fields in this simulation are based on observations (Appendix 6A). The

“future” simulation (REF2) is a self-consistent simulation from the past into the future. In this simulation the

surface time series of halocarbons is based on the “Ab”

scenario from WMO (2003); see Appendix 6A. The new

halogen scenario A1 from this Assessment (Chapter 8)

has not been applied because the computational requirements of the CCMs meant that these simulations had to

be started well before the scenarios in Chapter 8 were

finalized. The WMGHG concentrations for the future

simulations are taken from the IPCC (2000) “A1B” scenario, while sea surface temperatures (SSTs) are taken

from different coupled ocean model simulations, either

from simulations with the ocean coupled to the underlying general circulation models, or from the U.K.

Meteorological Office HadGEM1 simulations using

IPCC (2000) scenario “A1B.” Some CCMs ran multiple

future simulations with the same boundary conditions

but different initial conditions (see Table 6-4). In general, the variability between ensembles from a single

model is much smaller than the inter-model differences

(see also Austin and Wilson, 2006; Dameris et al., 2006).

For consistency, the 2-D model simulations used

the same halogen and WMGHG scenarios as the CCMs.

Additional sensitivity runs with different scenarios were

also performed by some 2-D models (see discussion below

and Table 6-3). In this chapter, the focus is on the “future”

simulations out to 2100, whereas the “past” simulations

(up to 2004) are discussed in Chapter 3. Of the 13 CCMs

available (see Table 6-4 for a summary of their characteristics), all but UMETRAC and LMDZrepro provided simulations of future ozone changes, and their results form

the basis for the projections discussed in Sections 6.6.3

and 6.6.4.

6.6.2 Model Evaluation

Most of the 2-D models listed in Table 6-3 have

been used in earlier Assessments. Their use in the past has

often been accompanied by model intercomparisons and

comparisons with measurements to characterize their

capabilities and deficiencies. For example, Park et al.

(1999) assessed the chemistry and transport in a large

number of 2-D and 3-D chemical transport models (CTMs)

(including the AER, GSFC-INT, GSFC, and NOCAR 2-D

models considered here). This study showed that there

was a large variation in the transport in the models, and

most models produced air that was too young in the stratosphere and did not correctly simulate the tropical “tape

recorder” (Mote et al., 1996). Also, there were significant

differences in the NOy and Cly in the lower stratosphere.

Here, 2-D models are used to predict ozone changes in the

midlatitudes and tropics. More emphasis has been given

to interactive 2-D models because non-interactive 2-D

models neglect the important effects of temperature feedback from changes in WMGHGs and the ozone radiativedynamical feedback.

An evaluation of the CCMs used here is reported in

Eyring et al. (2006). They compared simulations of the

recent past (1960 to 2004) with meteorological analyses

and trace gas observations. This CCM evaluation provides guidance on the level of confidence that can be

placed on each model simulation.

The comparisons in Eyring et al. (2006) showed

that the models reproduce the global, annual mean temperature fairly well, but most CCMs still have a cold bias

in winter-spring in the Antarctic resulting in later polar

vortex breakup (most severe in LMDZrepro and E39C).

Most models display the correct stratospheric response to

wave forcing in the Northern Hemisphere, but in the

Southern Hemisphere several of the models (e.g.,

CCSRNIES, E39C, MAECHAM4CHEM, MRI, ULAQ,

and LMDZrepro) have temperatures that are rather unresponsive to changes in the heat fluxes, with the ULAQ

model having the incorrect sign. These biases indicate

problems in the simulation of the dynamical response of

the Southern Hemisphere stratosphere in winter in these

models.

Eyring et al. (2006) also evaluated the transport

in the CCMs by comparing simulated methane, mean

age of air, and the water vapor “tape recorder” with

observations. They found a wide spread in the model

results for each diagnostic, indicating large differences

in the transport. However, the majority of the spread is

due to only a small subset of models (E39C,

MAECHAM4CHEM, MRI, SOCOL, and ULAQ) where

large deviations from observations are apparent in several, if not all, of the transport diagnostics. This is illustrated for methane and mean age of air at 50 hPa in

Figure 6-7(a) and (b), where the above models are shown

as dashed curves. The cause of significant biases in the

tracer fields in these models is generally not known, but

in some cases it can be attributed to specific model

21st CENTURY OZONE LAYER

6.22
21st CENTURY OZONE LAYER

C H 50 hPa October Mean Age of Air 50 hPa

CH4 (ppmv)

HCl 50 hPa November

HCl (ppbv)

Mean Age of Air (years)

Latitude Latitude

Latitude

Figure 6-7. Comparison of climatological zonal-mean (a) methane in October, (b) annual mean of the mean

age of air, and (c) HCl in November, at 50 hPa from CCMs and observations. Observations in (a) and (c) are

from HALOE: Black dots are values averaged in latitudes (zonal means) and the gray area shows plus and

minus 1 standard deviation about the climatological zonal mean HALOE. Observations in (b) are based on

ER-2 aircraft measurements of CO2 from many different years and months (Andrews et al., 2001). Adapted

from Eyring et al. (2006).

6.23
21st CENTURY OZONE LAYER

6.24

Table 6-4. A summary of the CCMs used in this Assessment. See Appendix 6A for descriptions of the

model runs. All CCMs have a comprehensive range of stratospheric chemical reactions, except E39C and

MAECHAM4CHEM, which do not include bromine chemistry.

Model Institution(s) Investigators Horizontal No. Levels/ Runs Reference

Resolution Upper

Boundary

AMTRAC GFDL, U.S. J. Austin 2° × 2.5° 48 / 3×REF1a Austin et al.

R. Wilson 0.0017 hPa 3×SCN2b (2006); Austin

and Wilson

(2006)

CCSRNIES NIES, Tsukuba, H. Akiyoshi 2.8° × 2.8° 34 / REF1 Akiyoshi et al.

Japan T. Nagashima (T42) 0.01 hPa REF2 (2004); Kurokawa

M. Yoshiki et al. (2005)

CMAM MSC, Univ. of J. McConnell 3.75° × 3.75° 71 / 3×REF2e Beagley et al. 

Toronto and N. McFarlane (T32) 0.0006 hPa (1997);

York Univ., D. Plummer de Grandpré

Canada J. Scinocca et al. (2000)

T. Shepherd

E39C DLR Oberpfaf- M. Dameris 3.75° × 3.75° 39 / 3×REF1 Dameris et al.

fenhofen, V. Eyring (T30) 10 hPa 4×SCN2 (2005, 2006)

Germany V. Grewe 2×NCC

M. Ponater

GEOSCCM NASA/GSFC, A. Douglass 2° × 2.5° 55 / REF1c Bloom et al.

U.S. P. Newman 0.01 hPa REF2 (2005);

S. Pawson Stolarski

R. Stolarski et al. (2006)

LMDZrepro IPSL, France S. Bekki 2.5° × 3.75° 50 / REF1 Chemistry

F. Lott 0.07 hPa REF2 part: Lefèvre 

F. Lefèvre et al. (1994)

M. Marchand

MAECHAM4 MPI Mainz, C. Brühl 3.75° × 3.75° 39 / REF1 Manzini et al.

CHEM Hamburg, M. Giorgetta (T30) 0.01 hPa REF2 (2003); Steil

Germany E. Manzini SCN2 et al. (2003)

B. Steil

MRI MRI, Tsukuba, K. Shibata 2.8° × 2.8° 68 / REF1 Shibata and

Japan M. Deushi (T42) 0.01 hPa REF2 Deushi (2005);

Shibata et al.

(2005)

SOCOL PMOB/WRC E. Rozanov 3.75° × 3.75° 39 / REF1 Egorova et al.

and ETHZ, M. Schraner (T30) 0.01 hPa REF2 (2005); Rozanov

Switzerland et al. (2005)

ULAQ Univ. of E. Mancini 10° × 22.5° 26 / REF1d Pitari et al.

L’Aquila, G. Pitari 0.04 hPa REF2 (2002)

Italy NCC 

UMETRAC Met Office, N. Butchart 2.5° × 3.75° 64 / REF1 Austin (2002);

U.K.; NIWA, H. Struthers 0.01 hPa Struthers et al. 

New Zealand (2004)
features, e.g., very low horizontal resolution (ULAQ) or

a low upper boundary (E39C). For the remaining CCMs

there is, in general, reasonable agreement with observations. In the upper stratosphere, the model agreement

with age of air observations is poorer, with most models

underestimating the age (not shown). However, in general the mean age of air and the tape recorder are in better

agreement than reported in the Hall et al. (1999) assessment of 2-D and 3-D CTMs.

The ability of CCMs to reproduce past stratospheric chlorine and bromine concentrations clearly

affects the confidence that we can place in their projections of future ozone changes, particularly in the

Antarctic. Eyring et al. (2006) examined modeled HCl

and inorganic Cly fields, and again found a large model

spread, with some large deviations from observations, see

Figure 6-7(c) and Figure 6-8(a). The differences are most

pronounced in the polar lower stratosphere, where peak

Cly varies from around 1 ppb to over 3.5 ppb. Measurements of Cly in the Antarctic lower stratosphere (symbols in Figure 6-8(a)) clearly show that peak values of Cly

close to or less than 2.5 ppb, as simulated by several

CCMs, are unrealistic. Transport deficiencies are a major

contributor to deficiencies in the simulated HCl and Cly,

and the models discussed above that did not perform well

for transport diagnostics also showed differences from

observed HCl. In the MRI model, the age of air decreases

significantly with time and, unlike in all other models and

observations, Cly does not peak around 2000 but continues

to increase until after 2015. This unrealistic continued

increase of Cly lowers the confidence we can place in this

simulation. Transport deficiencies do not, however,

explain all of the differences. The initial decrease in Cly

in UMSLIMCAT is due to wrong initial conditions used

in this simulation. AMTRAC and UMETRAC have

higher HCl and Cly than other models but similar CH4 and

mean age of air, and presumably similar transport (Figure

6-7 and 6-8). The additional Cly results from photolysis

rates of organic chlorine species being artificially

increased by 25% in AMTRAC and UMETRAC so that

the Cly in the upper stratosphere is in close agreement

with observed Cly. Although this adjustment of the photolysis rates improves the agreement with observed Cly in

the upper stratosphere and polar lower stratosphere

(Figure 6-8(a)), HCl (and also Cly) in the extrapolar lower

stratosphere is unrealistically large (Figure 6-7(c)). It is

not clear whether the 25% increase in the photolysis rates

is responsible for the enhanced extrapolar Cly.

The CCMs are generally able to reproduce the

observed amplitude and phase of the mean annual cycle in

total column ozone, except over southern high latitudes

(see right panels in Figure 3-26, Chapter 3). However,

most models exhibit large offset biases in the mean annual

cycle, with the majority overestimating total column

ozone. Although it is not possible to trace all differences

in the simulated ozone to deficiencies in the simulated

temperature and tracers, in some cases a link can be made.

For example, the models with low tropical and midlatitude CH4 have high ozone (MAECHAM4CHEM and

MRI), those with low Cly in polar regions have smaller

21st CENTURY OZONE LAYER

6.25

Table 6-4, continued.

UMSLIMCAT University of M. Chipperfield 2.5° × 3.75° 64 / REF1 Tian and 

Leeds, U.K. W. Tian 0.01 hPa REF2 Chipperfield 

(2005)

WACCM (v.3) NCAR, U.S. B. Bovillei 4° × 5° 66 / 3×REF1a Park et al.

R. Garcia 4.5×10-6 hPa 3×REF2 (2004); Richter

A. Gettelman 2×NCC and Garcia 

D. Kinnison (2006)

D. Marsh

F. Sassi

GFDL, Geophysical Fluid Dynamics Laboratory; NIES, National Institute for Environmental Studies; MSC, Meteorological Service of

Canada; DLR, Deutschen Zentrum für Luft- und Raumfahrt; NASA, National Aeronautics and Space Administration; GSFC, Goddard

Space Flight Center; IPSL, Institut Pierre Simon Laplace; MPI, Max-Planck-Institut; MRI, Meteorological Research Institute;

PMOB/WRC, Physical-Meteorological Observatory/World Radiation Center; ETHZ, Eidgenössische Technische Hochschule Zurich;

NIWA, National Institute of Water and Atmospheric Research; NCAR, National Center for Atmospheric Research. a REF1 but with SSTs from the Smith/Reynolds dataset (J. Hurrell, private communication) and without QBO b SCN2 without QBO c REF1 without QBO, solar cycle, and volcanic eruptions d REF1 without QBO and solar cycle e Different SSTs were used in each simulation
ozone reductions there (SOCOL and E39C), and the model

with the largest cold bias in the Antarctic lower stratosphere in spring (LMDZrepro) simulates very low ozone.

CCMs show a large range of ozone trends over the

past 25 years (see left panels in Figure 3-26 of Chapter 3)

and large differences from observations. Some of these

differences may in part be related to differences in the simulated Cly, e.g., E39C and SOCOL show a trend smaller

than observed, whereas AMTRAC and UMETRAC show

a trend larger than observed in extrapolar area weighted

mean column ozone. However, other factors also contribute, e.g., biases in tropospheric ozone (Austin and

Wilson, 2006).

The CCM evaluation discussed above and in Eyring

et al. (2006) has guided the level of confidence we place

on each model simulation. The CCMs vary in their skill

in representing different processes and characteristics of

the atmosphere. Because the focus here is on ozone

recovery due to declining ODSs, we place importance on

the models’ ability to correctly simulate stratospheric Cly

as well as the representation of transport characteristics

and polar temperatures. Therefore, more credence is given

to those models that realistically simulate these processes.

Figure 6-7 shows a subset of the diagnostics used to evaluate these processes and CCMs shown with solid curves

in Figures 6-7, 6-8, 6-10 and 6-12 to 6-14 are those that

are in good agreement with the observations in Figure

6-7. However, these line styles should not be overinterpreted as both the ability of the CCMs to represent

these processes as well as the relative importance of Cly,

temperature, and transport vary between different regions

and altitudes. Also, analyses of model dynamics in the

Arctic, and differences in the chlorine budget/partitioning

in these models, when available, might change this evaluation for some regions and altitudes.

21st CENTURY OZONE LAYER

6.26

Figure 6-8. October zonal mean values of total inorganic chlorine (Cly in ppb) at 50 hPa and 80°S from CCMs.

Panel (a) shows Cly and panel (b) difference in Cly from that in 1980. The symbols in (a) show estimates of Cly

in the Antarctic lower stratosphere in spring from measurements from the UARS satellite in 1992 and the Aura

satellite in 2005, yielding values around 3 ppb (Douglass et al., 1995; Santee et al., 1996) and around 3.3 ppb

(see Figure 4-8), respectively.

50 hPa 80°S October 50 hPa 80°S October

Cl

y Cl

y (1980) (ppbv)

Cl

y (ppbv)

 Year Year
6.6.3 Midlatitude and Tropical Ozone

Figures 6-9 and 6-10 show the time evolution

through the 21st century of annual means of monthly total

column ozone anomalies from the 2-D models and CCMs,

respectively, for four different regions: extrapolar (60°S

to 60°N), tropics (25°S to 25°N), northern midlatitudes

(35° to 60°N), and southern midlatitudes (35° to 60°S).

For the 2-D models, ozone from the P1 and F1 simulations

is shown, while for the CCMs the ozone is from REF1 and

REF2 or SCN2 (depending on the CCM) simulations (see

Table 6-4 for description of simulations). Multiple curves

are shown for CCMs that ran multiple simulations with

the same forcing. The method used to calculate the ozone

anomalies is detailed in Eyring et al. (2006).

The evolution of the ozone anomalies is qualitatively similar for both the 2-D models and the CCMs: the

lowest ozone occurs around 2000 within a broad minimum

after which ozone increases, as expected from decreasing

EESC beyond 2000. A clear difference between the 2-D

model and CCM projections is the year-to-year fluctuations in the ozone anomalies. The 2-D models show a

smooth ozone evolution, whereas the CCMs show large

interannual variability. The variability in the CCMs arises

from internal dynamics and variations in prescribed sea

surface temperatures with time. Variability associated with

the QBO and 11-year solar cycle is also included in some

CCM simulations (see Table 6-4) and contributes to the

interannual and longer time variations. Since the focus is

on decadal changes, the ozone anomalies plotted in Figure

6-10 have been smoothed using a 1:2:1 filter applied 5

times iteratively. There is a large spread in projected ozone

anomalies during the beginning of the 21st century (2000

to 2020). Two CCMs (AMTRAC and MRI) show much

larger ozone anomalies than the other models. These two

models also have much larger HCl and Cly during this

period (AMTRAC because of increased CFC photolysis

rates, and MRI because of slower circulation; see discussion in Section 6.6.2), which may explain part of the larger

ozone anomalies.

Averaged between 60°N and 60°S, total column

ozone is projected to increase between 2000 and 2020 in

all models except MRI (see above discussion), with most

of the increase of 1% to 2.5% occurring after 2010. Over

midlatitudes, the majority of the models predict an increase

of 1.5% to 3.5%, while over the tropics smaller ozone

increases of less than 2% are projected. The evaluation of

the timing of ozone minima is biased by the large response

to Mt. Pinatubo in the 2-D models, and obscured by interannual variability in the CCMs. If smoothed CCM anomalies are used (as shown in Figure 6-10) and, to avoid local

ozone minima caused by Mt. Pinatubo, only minima after

1996 are considered, both classes of model predict that

minimum ozone values have already occurred between

60°S and 60°N. In midlatitudes, 2-D models predict ozone

minima between 1997 and 2001 while the CCMs (again

with the exception of MRI) predict ozone minima between

1997 and 2007 with no clear hemispheric differences.

Because the exact timing of minimum ozone in some of

the CCM simulations might be affected by the solar cycle

(e.g., Dameris et al., 2006) or other factors, the ozone

minima detected in the CCMs are not synonymous with

ozone turnaround and should not be interpreted as stage

(ii) of ozone recovery.

It is expected that EESC in the midlatitude lower

stratosphere will decrease to its 1980 value between 2040

and 2050 (see box 8-1). If no other factors played a role,

it would be expected that ozone would increase to its 1980

values at the same time. However, if there are changes in

temperature, transport, or the abundance of other gases,

ozone may return to 1980 values earlier or later than EESC.

Figure 6-9 shows that total column ozone in the interactive 2-D models that include temperature feedbacks

from changes in WMGHGs (GSFC-INT, LeedsBremen,

NOCAR) is above 1980 values between 2040 and 2050.

Ozone averaged between 60°S and 60°N in these models

returns to 1980 values between 2025 and 2035, so that by

the time EESC returns to 1980 values, ozone is 0.5% to

3.0% above 1980 values. In the tropics and northern midlatitudes, ozone returns to 1980 values between 2020 and

2030, while over southern midlatitudes this happens somewhat later, between 2025 and 2040. The increase in ozone

to levels higher than would be expected from ODS concentrations alone results from stratospheric cooling due to

increased concentrations of WMGHGs. In contrast to the

interactive 2-D models, the non-interactive 2-D models

predict that column ozone will be less than or around 1980

values in 2040, and averaged between 60°S and 60°N,

column ozone returns to 1980 values about 5 years (but

up to 15 years) later than in the interactive models. This

is consistent with temperature changes playing a major

role in the increases in ozone.

The CCM simulations that extend to 2050 show a

similar recovery process over midlatitudes as the interactive 2-D models. Most CCMs (all but one/two in

northern/southern midlatitudes) predict midlatitude ozone

to be on average higher than 1980 values between 2040

and 2050, with increases to 1980 values generally occurring between 2005 and 2035 over northern midlatitudes

and over southern midlatitudes somewhat later between

2025 and 2040, in agreement with the interactive 2-D

models (Figure 6-10). The earlier return of northern midlatitude ozone to pre-1980 values is echoed in the hemispheric differences in ozone anomalies between 2040 and

21st CENTURY OZONE LAYER

6.27
21st CENTURY OZONE LAYER

6.28

960 1980 2000 20 20 2040 2060 2080 2100

Year

- 5

0

5

5

0

Annual mean anomaly (% deviation from detrended 1980 1989 mean annual cycle) 5

1 0

- 5

0

5

1 0

- 5

0

5

60 S-60 N

6 0 S-35 S

2 5 S-25 N

3 5 N-60 N

AER

NOCAR SU NY MPIC

Le eds Br emen MNP

OSLO GSFC-INT GSFC

Figure 6-9. Annual mean zonal mean total column ozone anomalies from the P1 (1980-2004) and F1 (2005-

2100) runs (Appendix 6A) from all of the 2-D models (colored lines; solid for interactive models and dashed for

non-interactive models) and from four observational datasets (thick black line and gray shaded area show the

mean and range of observed anomalies; see Chapter 3). Area-weighted zonal mean time series are shown for

the global region (60°S to 60°N), the equatorial region (25°S to 25°N), the northern midlatitude region (35°N to

60°N), and the southern midlatitude region (60°S to 35°S) to match the analysis regions used in Chapter 3.

The observational time series have been smoothed by applying a 1:2:1 filter iteratively five times. The filter

width is reduced to 1 at the ends of the time series, effectively reproducing the original data and avoiding

anomalous edge effects. The light gray shading between 2040 and 2050 shows the period when EESC is

expected to return to 1980 values. Monthly anomalies were calculated by subtracting a detrended mean

annual cycle, calculated over the period 1980-1989, from each time series. The annual cycle was detrended

by fitting a regression model with seasonally dependent trends to the data and then reconstructing the “1980”

mean annual cycle using the stationary components of the regression model. For further details on the method

for calculating the anomalies, see Eyring et al. (2006).
21st CENTURY OZONE LAYER

6.29

Figure 6-10. Annual mean zonal mean total column ozone anomalies from CCMs (colored lines) and from four

observational datasets (thick black line and gray shaded area show the mean and range of observed anomalies; see Chapter 3). The time series are formed using the REF1 and REF2 or SCN2 simulations of each

model (see Table 6-4). The time series have been smoothed as in Figure 6-9. The light gray shading between

2040 and 2050 shows the period when EESC is expected to return to 1980 values. As for the 2-D model

results, the anomalies were calculated by subtracting the detrended 1980-1989 mean annual cycle. The mean

annual cycles subtracted from the raw monthly means to calculate these anomalies are shown in Figure 3-26.

For further details on the method for calculating the anomalies, see Eyring et al. (2006).

960 1980 2000 20 20 20 40 20 60 20 80 2100

Y ea r

-15

-10

- 5

0

5

- 5

0

Annual mean anomaly (% deviation from detrended 1980 1989 mean annual cycle) -10

- 5

0

5

- 5

0

5

60 S-60 N

60 S-35 S

25 S-25 N

35 N-60 N

AMTRAC

E39C

CCSRNIES

ULAQ

SOCOL

MAECHAM4CHEM

UMSLIMCAT

MRI WACCM

CMAM

GEOSCCM
2050 (NH value of 0.5 to 5%; SH value of −0.5 to 3%).

The two exceptions to the above are AMTRAC and MRI,

which show a slower return of ozone to 1980 values such

that ozone is below 1980 values in 2040. The later return

of ozone to pre-1980 values in AMTRAC and MRI is consistent with the later decrease of Cly to 1980 values in these

models (although other factors, such as biases in the troposphere (Austin and Wilson, 2006), may also contribute).

The recovery process of tropical ozone differs

between the CCMs and interactive 2-D models. As discussed above, the interactive 2-D models predict tropical

ozone to be above pre-1980 levels by 2040. In contrast,

the majority of the CCMs predict ozone around or less

than 1980 values in 2040. The cause of this difference is

not known. Austin and Wilson (2006) noted that during

northern spring, tropical ozone remained below 1980

levels out to 2100 in a narrow band around the equator.

Rosenfield and Schoeberl (2005), using the GSFC-INT

interactive 2-D model, showed that ozone at 51 hPa and

5°N remains low out to 2050 in their model as a result of

increasing CO2, which cools the upper stratosphere, slows

ozone gas-phase destruction reactions, and increases

ozone there, which in turn reduces the penetration of UV

to the lower stratosphere, which reduces ozone production

there (a reverse of the “self-healing” effect). However,

given that the GSFC-INT model also shows ozone

exceeding pre-1980 levels between 25°S and 25°N (Figure

6-9), these suppressed ozone concentrations in the tropical lower stratosphere must not dominate column ozone

changes.

After 2050, the ozone changes primarily reflect

changes in WMGHGs. In the non-interactive 2-D models,

the ozone is roughly constant or declines. In contrast, in

the interactive 2-D models and in the one CCM that ran

beyond 2050 (Austin and Wilson, 2006), ozone continues

to slowly increase (except tropical ozone in AMTRAC and

in the GSFC-INT model). By the end of the 21st century,

these models predict that ozone averaged between 60°S

and 60°N will be 2% to 5% above 1980 values.

In the simulations presented above, only a single

scenario for WMGHG emissions has been considered.

However, the sensitivity to temperatures and/or composition caused by WMGHG increases has been examined in

other simulations. For example, Rosenfield et al. (2002)

used an interactive 2-D model to examine the impact of

CO2 increase-induced cooling on ozone recovery and

found that ozone returns to 1980 values 10 to 20 years earlier in many latitudes and seasons. Randeniya et al. (2002)

used a non-interactive 2-D model to examine the impact

of WMGHG-induced atmospheric composition changes

on ozone through the 21st century and showed that in

northern midlatitudes, the ozone recovery depended

greatly on the emissions scenarios used to drive the model.

After 2050, NOx increases due to N2O increases caused

stratospheric ozone levels to start to fall, and this loss was

exacerbated if future methane levels remained approximately constant, rather than growing, as assumed in many

scenarios.

The combined impact of CO2-induced cooling

and changes in N2O was subsequently examined by

Chipperfield and Feng (2003), using the Leeds-Bremen

interactive 2-D model. They performed two simulations,

one with and the other without temperature changes due

to increasing CO2, for three different WMGHG scenarios

(one with high emissions (A1FI), one with low emissions

(B1), and a third with A1FI emissions except with reduced

CH4 emissions). As shown in Figure 6-11, these calculations indicate that the inclusion of the CO2-induced cooling

overcomes the N2O chemical effect and leads to ozone

reaching 1980 levels or above between 2030 and 2050,

depending on the WMGHG scenario used. Similar results

were obtained in NOCAR simulations performed for this

Assessment (not shown). Consistent with Randeniya et

al. (2002), Figure 6-11 also shows that for the cases where

the CO2-induced cooling is excluded, ozone does not

return to pre-1980 values in the 21st century. Therefore,

the CO2-induced cooling and the N2O chemical effect

oppose each other, with the net effect dependent on the

WMGHG scenario chosen.

The sensitivity of ozone recovery to WMGHG scenarios has also been examined using CCMs that include a

more complete representation of climate-ozone feedbacks

(e.g., the variability in tropospheric wave forcing due to

climate change). As discussed in Chapter 5, simulations

were also repeated for several CCMs with fixed

WMGHGs, rather than increasing WMGHGs as in the

“A1B” scenario. These simulations show higher stratospheric temperatures and slower increases in midlatitude

ozone than in the simulations with increasing WMGHGs

(see Figure 5-25 of Chapter 5).

6.6.4 Polar Ozone

As discussed in Chapter 5, the processes affecting

ozone recovery in the polar regions are different from those

influencing extrapolar recovery. For example, WMGHGinduced cooling of the stratosphere reduces ozone destruction in extrapolar regions, but enhances it in the polar lower

stratosphere where heterogeneous chemistry dominates.

To investigate the evolution of polar ozone through the

21st century, we focus on the CCMs as 2-D models do not

include the three-dimensional processes that play a key

role in polar ozone depletion and recovery. Several dif21st CENTURY OZONE LAYER

6.30
ferent diagnostics of polar ozone from the CCMs are considered, including:

• springtime total column ozone anomalies for the

Arctic (March, 60°N-90°N) and Antarctic (October,

60°S-90°S), Figure 6-12;

• the minimum total column ozone poleward of 60°N

for March-April (Arctic) and 60°S for SeptemberOctober (Antarctic), Figure 6-13;

• the ozone mass deficit (Bodeker et al., 2005) from

September to October, Figure 6-14; and

• the maximum Antarctic ozone hole area between

September and October, Figure 6-14.

A discussion of the Antarctic ozone hole indices is provided in Chapter 4.

We examine these diagnostics for the same time

periods as above: the beginning of the century (2000-2020)

when EESC declines substantially, mid-century (2060-

2070) when EESC is expected to decrease to 1980 values,

and the end of the century (2090-2100) when ODSs are

not expected to play a role in controlling polar ozone.

6.6.4.1 THE ANTARCTIC

In the Antarctic, the general characteristics of the

ozone recovery are similar in all models and similar to the

CCM projections shown in Austin et al. (2003) and WMO

(2003), namely, the peak depletion occurs around 2000

within a broad minimum, followed by a slow increase in

ozone values. All CCMs show Antarctic ozone increasing

between 2000 and 2020, although the evolution varies

between models and between diagnostics. In general,

column ozone increases by around 5% to 10% during this

period. In nearly all models, the year of lowest October

column ozone anomalies occurs between 1997 and 2010

but, as shown in Figures 6-12, 6-13, and 6-14, the response

of the ozone diagnostic to changes in ODSs varies between

the diagnostics. The increase in 60°-90°S October ozone

anomalies and minimum Antarctic ozone is relatively

slow, with values remaining constant between 2000 and

2010 in many models, whereas there is a relatively fast

decrease in the ozone mass deficit. This behavior matches

the differences in response between different Antarctic

ozone hole indices seen in observations (Bodeker et al.,

2005). The fact that some of the CCMs suggest that minimum October ozone anomalies have already occurred in

the Antarctic may appear to contradict the conclusion made

in Section 6.5.3 that neither the first, nor second stage of

Antarctic ozone recovery has been detected in observations. It should be noted that the availability of model

results beyond 2005, as opposed to measurements that terminate in 2005, facilitates the detection of the minimum.

Ozone turnaround in measurements may have already

occurred but lack of data beyond 2005 precludes its statistical detection. Nevertheless, because the turnaround in

ozone as simulated in the CCMs has not been attributed

here to changes in EESC, this behavior cannot be interpreted as stage (ii) of ozone recovery.

There is a wide range in the simulated peak ozone

depletion in the CCMs. For example, the smoothed wintertime minimum Antarctic ozone values in the CCMs vary

21st CENTURY OZONE LAYER

6.31

35oN -60oN Figure 6-11. Variation in

Northern Hemisphere midlatitude total column ozone (%

change since 1980) from six

runs of a 2-D model. The simulations used three different

WMGHG scenarios (A1FI, B1,

and A1FI with B1 CH4 emissions), with results shown with

(solid lines) and without

(dashed lines) stratospheric

cooling due to CO2 increases.

Also shown are observed past

changes from satellite data.

From Chipperfield and Feng

(2003).
21st CENTURY OZONE LAYER

6.32

960 1980 2000 20 20 20 40 20 60 20 80 2100

Y ea r

-50

-40

-30

-20

-10

0

1 0

2 0

-20

-10

0

0

2 0

March/October anomaly (% deviation from detrended 1980 1989 mean annual cycle)

60 N-90 N

60 S-90 S

AMTRAC

E39C

CCSRNIES

ULAQ

SOCOL

MAECHAM4CHEM

UMSLIMCAT

MRI WACCM

CMAM

GEOSCCM

Figure 6-12. Upper panel: March Arctic (60°N to 90°N) total column ozone anomalies from CCMs (colored

lines) and from 4 observational datasets (thick black line and gray shaded area show the mean and range of

observed anomalies; see Chapter 3). Lower panel: as for the upper panel but October Antarctic (60°S to 90°S)

total column ozone anomalies. Model simulations are as in Figure 6-10. The time series have been smoothed

as in Figure 6-10. The light gray shading between 2060 and 2070 shows the period when EESC is expected to

return to 1980 values.
between around 60 DU and 120 DU compared with

observed values around 80 DU (Figure 6-13(b)), while the

peak smoothed ozone mass deficit varies between 5 and

33 million tons compared with observed values around 31

million tons (Figure 6-14(a)). These variations highlight

the deficiencies in some of the CCM simulations of the

present-day ozone hole. Note, however, that because both

the ozone hole area and the ozone mass deficit are based

on 220 DU thresholds, a general high bias in the globalmean total ozone fields will generate a bias in these two

diagnostics (e.g., most pronounced in MAECHAM4CHEM).

There is a wide spread in the projected Antarctic

ozone anomalies in 2050 and in the projected date when

Antarctic ozone increases to or around 1980 values. The

anomalies in 60°-90°S ozone in 2050 (see Figure 6-12)

vary from around −24% to 3%, while the date for Antarctic

ozone to increase to 1980 values varies between 2035 and

2095. However, this wide spread is mainly due to a few

models (AMTRAC, MRI; see further discussion); in the

majority of the CCMs, ozone is around 1980 values

between 2040 and 2050.

There is no simple relationship in the CCMs

between the date that Antarctic ozone increases to 1980

values and either the date of the minimum in Antarctic

ozone or the ozone anomaly at this date. Hence, comparison of the simulated ozone hole with observations (as

shown in Figures 6-13 and 6-14) does not provide a good

indicator of differences in the CCM projections of the

future ozone hole, and the wide spread shown in Figures

6-12 to 6-14.

However, some insight into the model differences

can be obtained from comparisons of inorganic chlorine

(Cly) in the models. A relationship is expected between

the decrease in Cly and increase in ozone, and, in general,

the increase in Antarctic ozone follows the decrease in Cly

in the CCMs. As shown in Figure 6-8, there is a large

spread in the simulated Antarctic Cly, including in the peak

value and the date at which the Cly decreases to 1980

values. Models with a smaller peak have an earlier return

of Cly to pre-1980 levels (e.g., SOCOL and E39C) and

those with a larger peak have a later return (e.g.,

AMTRAC). Comparison with observations showed that

peak Cly in several models is unrealistically low (see

Section 6.6.2), and the above analysis indicates that the

return to pre-1980 values of both Cly and ozone will be

too early in these models. We thus put more weight on

results from CCMs with higher, and more realistic, Cly.

These models predict Cly and ozone back to pre-1980

values around or later than 2050. The MRI model simulates a later return of ozone to pre-1980 values due to an

unrealistic too slow decrease of Cly after 2000 and we put

less weight on this model simulation. The Cly in

AMTRAC matches the observations best, and predicts the

latest return to pre-1980 values (Figure 6-8). The better

agreement of polar Cly with observations gives support

for this late return, but the justification for altering CFC

photolysis rates has to be questioned (see Section 6.6.2).

The late return to pre-1980 values in the AMTRAC

model is, nevertheless, consistent with the parametric

modeling study of Newman et al. (2006). Using a parametric model of spring Antarctic ozone amounts that

includes EESC levels over Antarctica (see Box 8.1 of

Chapter 8) and analyzed Antarctic stratospheric temperatures, they predicted that the ozone hole area will remain

constant until around 2010, that a decrease of the ozone

hole area would not be statistically detectable until around

2024, and that return to pre-1980 levels in the Antarctic

would not occur until around 2068; see Figure 6-14(b).

AMTRAC is the only CCM to run past 2050. This

model predicts that springtime Antarctic ozone will be

close to or just below (−7% to 3%) 1980 values by the end

of the century (2090-2100). However, uncertainties in the

projection are high because they are based on a single

model and single WMGHG concentration scenario.

6.6.4.2 THE ARCTIC

As in the other regions, Arctic ozone is projected to

increase from 2000 to 2020 (see Figures 6-12 and 6-13).

The increases range between 0% and 10% and there is

large interannual variability. Using the smoothed minimum Arctic ozone or 60°-90°N ozone anomalies, where

the smoothing reduces the large variability, the date of

minimum ozone occurs between 1997 and 2015 in nearly

all of the models. This time range is similar to that for the

Antarctic. As in the Antarctic, although the timing of the

minimum anomaly is similar between models, there is a

substantial variation in the magnitude of the ozone when

this occurs, with the smoothed minimum Arctic ozone

varying from around 200 DU to over 300 DU, compared

with around 220 DU in the observations. This again highlights some substantial biases in many of the models.

Most CCMs that have been run to 2050 show Arctic

ozone values larger than 1980 values in 2050. In other

words, all CCMs show Arctic ozone increasing to pre1980 values before 2050 (and before EESC returns to 1980

concentrations; 2060-2070). There is a wide range of dates

for Arctic ozone to increase to 1980 values, with the date

that smoothed minimum Arctic ozone increases to 1980

values varying between around 2010 and 2040. Although

there are large differences in the CCM simulations of

Arctic ozone, all, except MRI, show an increase to pre21st CENTURY OZONE LAYER

6.33
21st CENTURY OZONE LAYER

6.34

Figure 6-13. (a) Minimum Arctic total column ozone for March to April and (b) minimum Antarctic total column

ozone for September to October for various transient CCM simulations. Model simulations are as in Figure 6-

10, except a single ensemble simulation is shown for each model. Model results are compared with observations calculated using the National Institute of Water and Atmospheric Research (NIWA) combined total column

ozone database (Bodeker et al., 2005). Solid and dashed curves show smoothed data derived by applying a

1:2:1 filter iteratively 30 times. The filter width is reduced to 1 at the ends of the time series, effectively reproducing the original data and avoiding anomalous edge effects. The light gray shading between 2060 and 2070

shows the period when EESC is expected to return to 1980 values.
21st CENTURY OZONE LAYER

6.35

Figure 6-14. (a) The Antarctic September to October average daily ozone mass deficit for each year from

each CCM and (b) the maximum Antarctic ozone hole area between September and October. Model simulations are as in Figure 6-10, except a single ensemble simulation is shown for each model. Model results are

compared with observations calculated using the National Institute of Water and Atmospheric Research (NIWA)

combined total column ozone database (Bodeker et al., 2005). The curves show the data smoothed as in

Figure 6-13. Gray circles show the projection from Newman et al. (2006). The light gray shading between

2060 and 2070 shows the period when EESC is expected to return to 1980 values.
1980 ozone before or around 2050 with no indication of

severe Arctic ozone loss over this time.

In all CCMs that have been run long enough, Arctic

ozone increases to 1980 values before Antarctic ozone

does. In some models the difference is only a few years,

but in others the difference is over 25 years. The projected

evolution of Arctic ozone does not follow the evolution of

halogens as closely as in the Antarctic, i.e., the spread of

the projected Arctic ozone is not strongly related to the

spread in simulated Cly (the simulated Arctic Cly is similar

to that shown in Figure 6-8). This indicates that changes

in other factors (e.g., temperature and transport) play a

significant role in determining when Arctic ozone returns

to pre-1980 values. Austin and Wilson (2006) reported

that the earlier return to pre-1980 values in the Arctic in

AMTRAC results from both an increased Brewer-Dobson

circulation and reduced gas-phase ozone depletion in a

cooler stratosphere. However, it is unclear whether this is

occurring in all models. For example, in the E39C, ULAQ,

and WACCM simulations with fixed WMGHGs discussed

in Chapter 5, the return to pre-1980 values is very similar

to that in simulations with increasing WMGHGs presented

here (see Figure 5-26 of Chapter 5). This suggests that in

these models, there is little impact of increased WMGHGs

on Arctic ozone recovery before 2050.

Projections from AMTRAC, the only CCM to run

past 2050, suggest that Arctic ozone in 2090 to 2100 will

be substantially above (8% to 19%) 1980 values. This is

in contrast to Antarctic ozone in this model, which is close

to or just below 1980 values by the end of the century.

The increase in the Arctic column ozone minima in

the above CCM simulations is much faster than the empirical estimation of Knudsen et al. (2004), where the Arctic

ozone losses would increase until 2010-2015 and decrease

only slightly afterwards. However there are large uncertainties in the Knudsen et al. calculations: the observed

temperature and water vapor trends used in the analysis

are uncertain, and most important, it is unlikely that the

past trends will (as they assumed) continue unchanged into

the future (see Austin, 2004; Pitari, 2004). Given the likelihood of changes in trends in the Arctic, we put more

weight on the CCM projections, which attempt to capture

these changes, than the Knudsen empirical predictions.

6.6.5 Uncertainties in Model Projections

and Open Questions 

Sources of uncertainty in model projections of future

ozone range from those intrinsic to the models, such as

uncertainties in parameterizations and adequate incorporation of feedbacks, to those external, such as uncertainties

in the emissions scenarios and sea surface temperature

datasets used to drive the models. Many of the intrinsic

sources of uncertainties in CCMs are discussed in Chapter

5 and the uncertainties resulting from projections of future

ODS emissions are discussed in Chapter 8. The purpose

of this section is to discuss some of the processes that are

not included in the models that may contribute to uncertainty in ozone projections.

The model results presented above have focused on

one WMGHG emissions scenario (see Appendix 6A). The

IPCC Special Report on Emissions Scenarios (SRES)

(IPCC, 2000) A1B scenario is by no means the most likely

scenario, and the use of different, but possibly equally likely,

scenarios is expected to cause different projections of future

ozone, see Figure 6-11. Furthermore, the SRES scenarios

define only changes in anthropogenic emissions and not

concurrent changes in natural emissions due, for example,

to changes in surface climate. Climate change-induced

increases in emissions of halogenated very short-lived

sources gases, changes to their transport to the tropical

tropopause layer (TTL), and their degradation to bromine,

chlorine, and iodine (see Section 2.4 of Chapter 2) have the

potential to change tropical ozone abundances in the future.

Given the current resolution of the TTL region in CCMs,

we cannot fully assess this question. Currently our understanding from transport studies would indicate that only a

small fraction of air is directly lifted vertically through the

TTL into the lower tropical stratosphere, but there are indications that the “sideways” transport into the extratropical

stratosphere might be significant and sometimes fast

(Levine et al., 2006).

At present, the prescribed model boundary conditions are not the emissions scenarios themselves but concentration scenarios derived from the emissions scenarios

(Prather and Ehhalt et al., 2001). This decoupling of the

processes linking changes in emissions to changes in concentrations of WMGHGs and ODSs neglects feedbacks

that could be important, e.g., future increases in stratospheric ozone may change tropospheric photolysis rates

through enhanced absorption of solar UV, and change tropospheric ozone levels by increasing the stratospheretroposphere exchange source of ozone (Hauglustaine et

al., 2005). These in turn affect tropospheric oxidizing

capacity, which affects WMGHG and ODS lifetimes. The

effects of such changes on lifetimes have been ignored in

the construction of ODS and WMGHG concentration scenarios; see Section 8.3.1 of Chapter 8 and Prather and

Ehhalt et al. (2001), respectively. Natural processes, such

as volcanic eruptions (see Section 6.3.6), could also affect

tropospheric oxidizing capacity and/or change the effectiveness of stratospheric ODSs in depleting ozone.

New anthropogenic source gases, currently

excluded from emissions scenarios, may become impor21st CENTURY OZONE LAYER

6.36
tant in the future, e.g., fugitive emissions of hydrogen from

a hydrogen economy. Tromp et al. (2003) estimated that a

hydrogen economy could cause stratospheric water vapor

increases of up to 35% and decrease ozone by up to 10%

in polar regions during spring. However, Warwick et al.

(2004) found a much reduced effect on stratospheric ozone

on the order of 0.5% or less. This reduced estimate is the

result of an assumed smaller leakage rate of hydrogen combined with the inclusion of reduced CO, NOx, CH4, and

nonmethane hydrocarbons due to the reduction of fossil

fuel usage. Furthermore, the timing of the conversion to a

hydrogen economy with respect to the reduction of halogen

loading is a major factor, because the coupling to chlorine

chemistry causes most of the ozone loss. If the shift to a

hydrogen economy occurs primarily after 2020 (Kammen

and Lipman, 2003), the potential harm due to the increase

in anthropogenic hydrogen emission can be prevented.

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Appendix 6A

MODEL SCENARIOS

Table 6A-1 lists the forcings used in the 2-D model and Chemistry-Climate Model (CCM) simulations. Two reference and two sensitivity simulations along with a set of model forcings have been proposed as part of the CCM Validation

Activity for the Stratospheric Processes and their Role in Climate (SPARC) project of the World Climate Research

Programme (Eyring et al., 2005) to assess the near-term and long-term evolution of stratospheric ozone. Unless otherwise

stated, both types of models used changes in halocarbons as prescribed in Table 4B-2 of WMO (2003), WMGHGs based

on the IPCC (2000) Special Report on Emissions Scenarios (SRES) A1B scenario, and sulfate aerosols for 1979 to 1999

based on the climatology of David Considine (NASA Langley Research Center; see http://www.pa.op.dlr.de/CCMVal/lForcings/CCMVal_Forcings.html) with background values thereafter.

The sea surface temperatures (SSTs) in the CCMs are prescribed from observations (HadSST1) in past simulations

(Rayner et al., 2003). For REF2/SCN2 simulations they come either from the underlying IPCC coupled-ocean model

simulation or from the UK Meteorological Office HadGEM1 simulations using IPCC SRES scenario A1B. In the reference past simulations (REF1), the quasi-biennial oscillation (QBO) is either prescribed (Giorgetta and Bengtsson, 1999)

or internally generated, and the influence of the 11-year solar cycle on photolysis rates is parameterized according to the

intensity of the 10.7 cm radiation of the Sun (Lean et al., 1997). In future runs, the solar cycle and an external QBO

forcing are only considered in SCN2, but not in REF2.

In the CCMVal halogen file (i) The data from Table 4B-2 of WMO (2003), which were given every 5 years, were

linearly interpolated into monthly values; (ii) The data for each year in Table 4B-2 were interpreted as the midpoint of

each year rather than the correct 1st of the year; and (iii) Halon-2402 was mistakenly not included. Assumptions (i) and

(ii) result in a slightly smaller peak organic chlorine (CCly) than observed (3.56 ppb instead of 3.58 ppb) and the peak

occurs 2.4 years too late, while assumption (iii) results in organic bromine (CBry) that is around 1 ppt too low at the peak.

To test the sensitivity of these assumptions on ozone, one 2-D model has been used to rerun the simulation with a corrected Ab halogen scenario. The difference between the two simulations is very small (especially compared with the differences between models), and none of the conclusions from the analysis of the 2-D models or CCMs are affected by the

differences in the halogen forcing.

Table 6A-1. Model simulations.

Run 1 Period Motivation Halocarbons WMGHGs SSTs 2 Solar QBO

Variability

P1 1979- Past WMO (2003) IPCC (2000) N/A None None

2004 Trends Table 4B-2 A1B

F1 2005- Future, WMO (2003) IPCC (2000) N/A None None

2100 Control Table 4B-2 B1

REF1 1960/ Past WMO (2003) IPCC (2000) Observed Observed Forced or

1980- Trends Table 4B-2 A1B (HadISST1) (MAVER) internally

2004 generated

REF2 1960/ Future, WMO (2003) IPCC (2000) Modeled None Only

1980- Control Table 4B-2 A1B internally

2100 generated

SCN2 1980- Solar WMO (2003) IPCC (2000) Modeled Observed + Forced or

2025 Cycle and Table 4B-2 A1B repeating in internally

QBO future generated

NCC 1970- Fixed WMO (2003) Fixed at start Observed As in REF1/ Forced or

2050 GHGs Table 4B-2 of simulation 1970-79 REF2/SCN2 internally 

repeating generated

or modeled

1 P1 and F1 runs are for 2-D models, and remainder are for CCMs. 2 Applies to CCM simulations only.
CHAPTER 7

Surface Ultraviolet Radiation:

Past, Present, and Future

Lead Authors:

A.F. Bais

D. Lubin

Coauthors:

A. Arola

G. Bernhard

M. Blumthaler

N. Chubarova

C. Erlick

H.P. Gies

N. Krotkov

K. Lantz

B. Mayer

R.L. McKenzie

R.D. Piacentini

G. Seckmeyer

J.R. Slusser

C.S. Zerefos

Contributors:

U. Feister

V. Fioletov

J. Gröbner

E. Kyrö

H. Slaper

Final Release: February 2007

From Scientific Assessment of Ozone Depletion: 2006
SURFACE ULTRAVIOLET RADIATION

7.2

• New methods to quantify the aerosol optical properties have been developed. These properties have important

influences on UV radiation. By combining spectral irradiance and radiance measurements, the effective single scattering albedo and optical depth of aerosols can be derived with reasonable accuracy at locations with moderate to

high aerosol optical depth. Under low aerosol conditions, accurate determination of the wavelength dependence of

the aerosol optical depth in the UV-B is restricted by calibration uncertainties and by the interference of ozone and

SO2.

• Although the Total Ozone Mapping Spectrometer (TOMS) instrument is no longer available, continuity of

satellite-derived global UV data is maintained. The new Ozone Monitoring Instrument (OMI) onboard the

National Aeronautics and Space Administration (NASA) Earth Observing System (EOS) Aura spacecraft was

launched in July 2004 for continued global monitoring of ozone, other trace gases, and surface UV irradiance.

• Tropospheric aerosols are responsible for the overestimation of UV irradiance from satellite instruments

(e.g., TOMS) that use solar backscattered ultraviolet radiation to derive surface UV irradiance. Although at

clean sites the agreement with ground-based measurements is good, over more polluted locations the bias can be as

large as 40% because the lowermost atmosphere containing the absorbing aerosols is not adequately probed. The

presence of clouds, and snow or ice cover, can also lead to significant biases. New algorithms have been developed

to improve the parameterization of aerosol and snow and ice effects on satellite-derived surface UV irradiance, as

well as of cloud effects using Advanced Very High Resolution Radiometer (AVHRR) and METEOSAT images,

showing on average good agreement with ground-based UV observations.

• Further improvements have been made in UV measuring instruments and techniques. For example, a transportable spectroradiometer was compared with instruments at more than 25 sites in Europe. The uncertainty of

well-maintained spectroradiometers, however, could not be significantly reduced in recent years, mainly due to the

remaining difficulties with lamp calibrations. Diode array and CCD spectrographs record the entire spectrum in a

fraction of a second, allowing for better spectral characterization of cloud effects of UV radiation, but their intrinsic

stray-light problem limits their use in the UV-B. Narrowband multifilter radiometers are now used in several networks providing more information than broadband radiometers and are a useful supplement to well-maintained

spectroradiometers. Those that include rotating shadow-bands provide, in addition, estimates of the aerosol optical

properties.

• Algorithms have been developed for converting spectral irradiance to actinic flux, which is more relevant to

atmospheric chemistry. This enables us to derive historic actinic fluxes from the 1990s, when the irradiance measurements became widely available. The resulting uncertainties in the actinic fluxes are between 5% and 15% under

all sky conditions.

• Model calculations incorporating only ozone projections show that UV levels will decrease over the next few

decades. These calculations imply that UV irradiance is currently close to maximum and under this scenario they

will revert to pre-1980 levels at midlatitudes between about 2040 and 2070, but later at southern high latitudes.

However, other factors that influence UV are likely to dominate over these time scales.

• Climate change will also influence surface UV irradiance through changes induced mainly to clouds and surface reflectivity. Aerosols and air pollutants are also expected to change in the future. These factors may lead to

either increases or decreases in surface UV radiation, through absorption or scattering. If the projections for future

ozone are correct, these factors are likely to dominate future changes in UV radiation.
CHAPTER 7

SURFACE ULTRAVIOLET RADIATION: PAST, PRESENT, AND FUTURE

Contents

SCIENTIFIC SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.1

7.1 INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.3

7.2 FACTORS AFFECTING UV RADIATION: NEW FINDINGS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.3

7.2.1 The Extraterrestrial Solar Spectrum . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.4

7.2.2 Stratospheric Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.5

7.2.3 Clouds . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.5

7.2.4 Aerosol Scattering and Absorption . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.7

7.2.5 Tropospheric Ozone and Other Gaseous Air Pollutants . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.8

7.2.6 Surface Albedo . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.9

7.2.7 Altitude . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.10

7.2.8 UV Irradiances on Inclined Surfaces . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.11

7.2.9 UV Under Water, Ice, and Snow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.11

7.3 RESOURCES FOR STUDYING SURFACE UV RADIATION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.13

7.3.1 Ground-Based Measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.13

7.3.1.1 Spectroradiometers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.13

7.3.1.2 Broadband Filter Radiometers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.14

7.3.1.3 Narrowband Multifilter Radiometers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.14

7.3.1.4 Advancements in Instrumentation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .7.15

7.3.1.5 Data Quality and Databases . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.15

7.3.2 Satellite-Derived UV Estimates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.16

7.3.2.1 Differences Between Ground-Based and TOMS-Derived UV Irradiances . . . . . . . . . . . . . . . . 7.16

7.3.2.2 UV Retrieval Algorithms and Related Uncertainties . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.17

7.3.3 Advancements in UV Radiation Modeling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.17

7.3.4 Short-Term UV Forecasting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.20

7.4 UV CLIMATOLOGY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.21

7.4.1 Derived from Ground-Based UV Measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.21

7.4.1.1 Long-Term Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.21

7.4.1.2 Short-Term and Spatial UV Variability . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.25

7.4.2 Estimates from Satellite Data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.26

7.4.3 Reconstruction of Past UV Records . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.26

7.4.4 Polar UV Variability . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.28

7.5 EXPECTATIONS FOR THE FUTURE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.29

7.5.1 Links with the Recovery of the Ozone Layer . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.29

7.5.2 Predictions of UV Long-Term Changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.29

7.5.3 Factors Related to Climate Change . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.30

7.6 REMAINING QUESTIONS AND UNCERTAINTIES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.32

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.32

APPENDIX 7A: SPECTRAL DATA AVAILABLE FROM DATABASES

7A.1 World Ozone and Ultraviolet Radiation Data Centre (WOUDC) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.49
7A.2 European Ultraviolet Database (EUVDB) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.51

7A.3 U.S. National Science Foundation UV Monitoring Network . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.52

APPENDIX 7B: INTERNET ADDRESSES FOR UV SITES

7B.1 General UV Information . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.53

7B.2 International UV Projects . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.53

7B.3 Radiative Transfer Codes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.53

7B.4 Extraterrestrial Spectra . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.53

7B.5 Internet Sites with UV Information . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.54
SURFACE ULTRAVIOLET RADIATION

7.1

SCIENTIFIC SUMMARY

• Ultraviolet (UV) measurements from some stations in unpolluted locations indicate that the Montreal Protocol

is working, because at these locations further increases have not been observed since the late 1990s. Outside

polar regions, increases in UV due to ozone depletion have been relatively small and in many places they are difficult to separate from those due to other causes, such as changes in cloud and aerosol. In unpolluted locations, especially in the Southern Hemisphere, UV irradiances have been decreasing in recent years, as expected from the

observed increases in ozone. 

• At most midlatitude stations in the Northern Hemisphere, surface UV irradiance continued to increase at

rates of a few percent per decade. The observed increases and their significance depend on location, wavelength

range, and the period of measurements. These increases cannot be explained solely by ozone depletion and could

be attributed to a decreasing tendency in aerosol extinction and air pollution since the beginning of the 1990s and

partly to decreasing cloudiness, as estimated from satellites. The longest reconstructed series of erythemal irradiance (in Switzerland back to 1920s) revealed that high UV levels occurred in the mid-1940s and early 1960s due to

reduction in cloudiness. Increased cloudiness in the mid-1970s resulted in reduced erythemal irradiance over several locations in Europe.

• In polar regions, high UV irradiances lasting for a few days have been observed, associated with low total

ozone episodes. Erythemal irradiance averaged over several days has been increased by ~70% over southern

Argentina during vortex overpasses in October. Instantaneous enhancements up to a factor of 6 have been observed

over Antarctica. Over northern Europe and Alaska, the observed enhancements were smaller.

• At present, clouds generally influence surface UV irradiance more strongly than any other atmospheric variable, including ozone. Further studies that quantify cloud effects on surface UV irradiance are now available.

Under overcast conditions, reductions can exceed 90%. Reductions of UV irradiance by clouds are 15-45% smaller

than in the visible part of the spectrum. Recent measurements have shown intermittent enhancements of up to 40%

in UV actinic flux over clear-sky values for cloudy conditions when the solar disk is unoccluded, whereas enhancements of 25% were found for irradiance under similar conditions.

• Cloud variability is the major factor limiting our ability to detect long-term changes in surface UV radiation

due to ozone. At most sites, even if ozone trends were linear, at least 10-15 years of measurements would be

needed to detect a trend in UV radiation.

• The high surface reflectivity (close to unity) in Antarctica can lead to enhancements in clear-sky UV irradiance up to 50%. Under such conditions, the high albedo compensates the UV attenuation by clouds, resulting in

surface irradiance similar to clear skies, while zenith radiance is strongly enhanced (even tripled). For the first time

the UV surface albedo was measured in Antarctica with high spectral resolution, confirming that albedo in the UV

is slightly lower than in the visible, but still close to unity. New observations in the European Alps showed that a

snow-covered terrain may increase the clear-sky erythemal and UV-A irradiance by up to 22% and 15% respectively, with respect to low-albedo conditions. Spatial variations of snow coverage may result in large (up to 40%)

variations in sky radiance from different directions. The effect of snow albedo is stronger on tilted surfaces.

• Air pollutants may counterbalance the UV radiation increases resulting from ozone depletion. Observations

confirmed that surface UV irradiance at locations near the emission sources of ozone (O3), nitrogen dioxide (NO2),

or sulfur dioxide (SO2) in the lower troposphere is attenuated by up to ~20%. Air pollution exerts stronger attenuation in UV compared with total solar irradiance.

• Three-dimensional radiative transfer models have improved the ability to model UV radiation under broken

cloud and over complex terrains. For most measuring sites, these are the prevailing conditions. However, the

realization of this improved capability is still limited by the lack of the necessary input parameters.
SURFACE ULTRAVIOLET RADIATION

7.2

• New methods to quantify the aerosol optical properties have been developed. These properties have important

influences on UV radiation. By combining spectral irradiance and radiance measurements, the effective single scattering albedo and optical depth of aerosols can be derived with reasonable accuracy at locations with moderate to

high aerosol optical depth. Under low aerosol conditions, accurate determination of the wavelength dependence of

the aerosol optical depth in the UV-B is restricted by calibration uncertainties and by the interference of ozone and

SO2.

• Although the Total Ozone Mapping Spectrometer (TOMS) instrument is no longer available, continuity of

satellite-derived global UV data is maintained. The new Ozone Monitoring Instrument (OMI) onboard the

National Aeronautics and Space Administration (NASA) Earth Observing System (EOS) Aura spacecraft was

launched in July 2004 for continued global monitoring of ozone, other trace gases, and surface UV irradiance.

• Tropospheric aerosols are responsible for the overestimation of UV irradiance from satellite instruments

(e.g., TOMS) that use solar backscattered ultraviolet radiation to derive surface UV irradiance. Although at

clean sites the agreement with ground-based measurements is good, over more polluted locations the bias can be as

large as 40% because the lowermost atmosphere containing the absorbing aerosols is not adequately probed. The

presence of clouds, and snow or ice cover, can also lead to significant biases. New algorithms have been developed

to improve the parameterization of aerosol and snow and ice effects on satellite-derived surface UV irradiance, as

well as of cloud effects using Advanced Very High Resolution Radiometer (AVHRR) and METEOSAT images,

showing on average good agreement with ground-based UV observations.

• Further improvements have been made in UV measuring instruments and techniques. For example, a transportable spectroradiometer was compared with instruments at more than 25 sites in Europe. The uncertainty of

well-maintained spectroradiometers, however, could not be significantly reduced in recent years, mainly due to the

remaining difficulties with lamp calibrations. Diode array and CCD spectrographs record the entire spectrum in a

fraction of a second, allowing for better spectral characterization of cloud effects of UV radiation, but their intrinsic

stray-light problem limits their use in the UV-B. Narrowband multifilter radiometers are now used in several networks providing more information than broadband radiometers and are a useful supplement to well-maintained

spectroradiometers. Those that include rotating shadow-bands provide, in addition, estimates of the aerosol optical

properties.

a Algorithms have been developed for converting spectral irradiance to actinic flux, which is more relevant to

atmospheric chemistry. This enables us to derive historic actinic fluxes from the 1990s, when the irradiance measurements became widely available. The resulting uncertainties in the actinic fluxes are between 5% and 15% under

all sky conditions.

a Model calculations incorporating only ozone projections show that UV levels will decrease over the next few

decades. These calculations imply that UV irradiance is currently close to maximum and under this scenario they

will revert to pre-1980 levels at midlatitudes between about 2040 and 2070, but later at southern high latitudes.

However, other factors that influence UV are likely to dominate over these time scales.

• Climate change will also influence surface UV irradiance through changes induced mainly to clouds and surface reflectivity. Aerosols and air pollutants are also expected to change in the future. These factors may lead to

either increases or decreases in surface UV radiation, through absorption or scattering. If the projections for future

ozone are correct, these factors are likely to dominate future changes in UV radiation.
7.1 INTRODUCTION

The potential for increased solar ultraviolet (UV)

radiation reaching the Earth’s surface in response to ozone

(O3) reduction has been a major concern since the first

signs of ozone depletion in the early 1980s. Although UV

is a small fraction of the total radiant solar energy, it may

produce detrimental effects on the ecosystem and

degrading effects on materials, and therefore knowledge

of its variability in time and space has high priority in scientific research. Oxygen, ozone, and nitrogen molecules

in the upper atmosphere absorb the UV-C (wavelength

< 280 nanometers (nm)) part of the spectrum effectively,

whereas ozone in the stratosphere strongly absorbs the

UV-B (280-315 nm) part and consequently determines the

spectral shape of UV radiation at the surface. UV-B is of

great biological importance because photons in this region

may damage deoxyribonucleic acid (DNA) molecules and

some proteins of living organisms. On the other hand, UVB is essential for the synthesis of vitamin D in the human

body, which has beneficial health effects and helps in prevention of some diseases. Vitamin D can not be produced

in sufficient amounts when incident UV-B radiation is very

low. UV-A (315-400 nm) is less affected by ozone and

may also contribute to some biological effects, especially

if received in high doses. UV in both spectral ranges is

important for tropospheric chemistry because it is involved

in important photochemical reactions, such as the photolysis of nitrogen dioxide (NO2), ozone, and formaldehyde.

Changes in tropospheric composition induced by UV

changes may be important for the stratosphere through

interactions between the troposphere and stratosphere.

Finally, there are linkages between changes in UV-A radiation and climate change.

Monitoring of UV radiation has been a challenging

task because of the great difficulties in conducting accurate measurements and proper quality control, and because

UV is highly variable both in time and space. Assessment

of the UV levels in the last two decades is addressed

through a limited number of measurement records from

instruments operating at the ground or onboard satellites.

Most of the unresolved issues that were pointed out

in the previous Assessment (Kerr and Seckmeyer et al.,

2003) have been further studied. Some progress has been

made with respect to the effects of absorbing aerosols on

satellite UV retrievals, and different methods have been

proposed for determining the aerosol single scattering

albedo from radiation measurements and modeling.

Improvements have also been made in quantifying the

cloud effects and the influence of high surface albedo on

the ground-based and satellite-derived measurements.

Effects from inhomogeneities in the surface reflectivity of

the terrain surrounding a monitoring site have also been

studied, although marginal progress has been made in

investigating the representativeness of measurements

made at a monitoring site for the surrounding areas.

While propagating through the atmosphere, solar

UV radiation is influenced by complex processes of scattering and absorption before reaching the Earth’s surface.

Sufficient understanding and quantification of the effects

of these processes, particularly the tropospheric, are essential for the simulation of the UV radiation field and for

explaining its variations in time and space. Improvement

of our knowledge and more quantitative information on

these processes since the previous Assessment (WMO,

2003) are presented in Section 7.2. In particular this section describes effects arising from the Sun-Earth geometry, clouds, atmospheric gases and aerosols, as well as

effects from the surface albedo and altitude. Finally, new

findings about the transfer of UV under water and ice are

also included. Section 7.3 discusses the progress made in

instruments and procedures that are used to measure solar

UV radiation and other relevant parameters, from the

ground and space. The differences between ground-based

and satellite-derived surface UV measurements are also

discussed in connection with further developments in

satellite UV retrieval algorithms. The section also includes

developments in radiative transfer (RT) modeling and

other modeling approaches. Finally, this section discusses

the short-term predictions of surface UV and the use of

these predictions to construct the UV index. Changes in

surface UV irradiance on different time scales derived

from ground-based and satellite instruments, as well as

from reconstructed past UV records are discussed in

Section 7.4. Particular emphasis is given to long-term

changes and to the attribution of these changes to different

atmospheric parameters. The link between surface UV

and climate change, and prediction of future UV levels in

relation to ozone changes, are discussed in Section 7.5.

Finally, open issues that need further investigation are outlined in Section 7.6.

7.2 FACTORS AFFECTING UV RADIATION:

NEW FINDINGS

The non-uniform distribution of the predominant

factors that interact with UV radiation (clouds, ozone,

aerosols, and surface albedo) makes detailed calculations

of UV radiation a nontrivial task. The impact of these

factors cannot always be treated independently, since in

most cases they act synergistically (e.g., Kerr, 2005). The

following sections describe the progress that has been

achieved during the last four years toward understanding

and quantifying the role of these factors in modifying the

SURFACE ULTRAVIOLET RADIATION

7.3
solar UV radiation received at the surface. Due to the

large natural variability of surface UV, ground-based

measurements that are made in many locations worldwide

provide results that are valid mainly locally, and cannot

be easily extrapolated to global scales (see also Figure 7-

4). On the other hand, satellite estimates of surface UV

radiation provide global coverage, but the sampling frequency is typically only once per day, with a coarse spatial resolution that represents average conditions over

large areas. Further, these estimates are based on model

calculations that include assumptions that are not always

realistic.

7.2.1 The Extraterrestrial Solar Spectrum

UV radiation entering the Earth’s atmosphere is primarily controlled by the variations in the Sun emittance

and in the orbital position of the Earth. Uncertainties in

the spectral solar irradiance at the top of the atmosphere

translate directly into uncertainties in the spectral radiance

(and irradiance) at all levels in the atmosphere and at the

surface due to the linearity of the radiative transfer equation. Thus, accurate knowledge of the solar spectrum at

the top of the atmosphere (see Appendix 7B.4) is very

important for accurate prediction of the radiation levels at

the surface with radiative transfer models.

Variations in the orbit of the Earth, the so-called

Milankovitch cycles (obliquity, eccentricity, and precession), have very long periods and produce changes in radiation at the top of the atmosphere of up to ~30% on time

scales of hundreds of thousands of years. Shaffer and

Cerveny (2004) calculated surface irradiance changes in

the UV band 300-325 nm taking into account variations in

the extraterrestrial radiation reaching the atmosphere and

the solar zenith angle (SZA), but assuming invariant latitudinal ozone concentrations (see example in Figure 7-1).

Although over such long time scales there should have

been significant changes in oxygen, Segura et al. (2003)

showed that even for large changes in oxygen, changes in

ozone are relatively small. Variations in surface UV radiation caused by these orbital changes are expected to be

very small (less than ~1%) in the next few centuries. Thus

solar activity (sunspots, and 11-year and 27-day cycles) is

the dominant factor responsible for the variability in the

extraterrestrial irradiance spectrum in the near future.

Solar activity in the last 70 years has been exceptionally strong, and the previous period of equally strong

activity occurred more than 8,000 years ago, as suggested

by Solanki et al. (2004) based on dendrochronologically

dated radiocarbon concentrations. This strong solar

activity marginally influences the UV radiation that

reaches the surface (UV-B and longer wavelengths), but it

is important for the UV-C wavelengths that are involved

in stratospheric ozone production.

Spectroradiometers onboard Earth-orbiting satellites allow investigations of temporal variations of solar

irradiance in different wavelength ranges. Two new composite solar irradiance reference spectra extending from

0.1 to 2400 nm were constructed using recent space measurements for two distinct time periods during solar cycle

22 (1985-1997). From data gathered with instruments

onboard the Atmospheric Laboratory for Applications and

Science (ATLAS) 1 and 2 and the Upper Atmosphere

Research Satellite (UARS), the accuracy of the measured

solar spectra was determined to be ~3% in the UV and visible ranges (Thuillier et al., 2004). Gurlit et al. (2005) presented a moderate resolution solar irradiance spectrum in

the spectral range 316.7-652 nm derived from the azimuthcontrolled Limb Profile Monitor of the Atmosphere/Differential Optical Absorption Spectroscopy

(LPMA/DOAS) balloon gondola at around 32 km float

altitude. This spectrum is in very good agreement with

previously reported spectra (Kurucz et al., 1984; Thuillier

et al., 1997, 1998; Harder et al., 2000) in the visible, but it

SURFACE ULTRAVIOLET RADIATION

7.4

Figure 7-1. Percent deviations in daily near-UV

(300-325 nm) solar irradiance on the March equinox

at 60°N between 250,000 before present and 50,000

after present, caused by Milankovitch cycles. Larger

variations, ranging between about ±30%, have been

calculated for different latitudes and for the solstices.

Reprinted from Shaffer and Cerveny, 2004, with permission from Elsevier.
is 1.4 to 6.2% lower in the UV-A. In agreement with

Skupin et al. (2003), this study emphasizes that the present

level 1 calibration of the Scanning Imaging Absorption

Spectrometer for Atmospheric Chartography (SCIAMACHY) of the European Space Agency (ESA) is systematically 15% higher in this wavelength range (316.7-

652 nm) with respect to all other available solar irradiance

measurements, and suggests using the recalibration proposed by the University of Bremen for SCIAMACHY.

Finally, a new reference solar spectrum (given in steps of

0.5 nm) was constructed by combining measurements

derived from different satellite platforms and from the

ground, and this spectrum was compared with historical

reference spectra (Gueymard, 2004, 2006). It was concluded that reference solar irradiance spectra in the UV-B

and UV-A that were published after 1985 are in very good

agreement (within a few percent) and that the ATLAS 3

spectrum is probably the most suitable one for radiative

transfer calculations. The latter is in agreement with the

conclusion of Bais et al. (2003) that was derived from comparison of measured and modeled surface UV spectra.

The variability of the solar irradiance spectrum for

the duration of solar cycle 23 has been investigated using

data from different satellites. Several studies quantified

the amplitude of the solar cycle in the wavelength band

245-250 nm and found that it ranged between 6% and 8%:

DeLand et al. (2004a) used data in the period from March

1985 to May 1997 from the Solar Backscatter Ultraviolet

Radiometer 2 (SBUV/2) onboard the National Oceanic

and Atmospheric Administration (NOAA) 9 satellite;

DeLand et al. (2004b) used data from the Solar Ultraviolet

Spectral Irradiance Monitor (SUSIM) and the Solar Stellar

Irradiance Comparison Experiment (SOLSTICE) instruments aboard UARS; and Floyd et al. (2003) used the

SUSIM dataset alone. In the band 200-205 nm, the amplitude is larger (~9%). All three analyses reveal that no irradiance variations due to solar cycle variability could be

detected in the UV-B and UV-A ranges at these spectral

resolutions. Therefore surface UV radiation is not

expected to respond directly to solar variations, but only

indirectly through changes in stratospheric ozone induced

by the variability of solar irradiance in the far ultraviolet

(see Chapter 3, Section 3.4.4). A weaker solar activity will

result in less ozone in the stratosphere and thus in more

UV-B radiation at the surface, but also in less risk of

sudden ozone decreases due to solar flares.

7.2.2 Stratospheric Ozone

The influence of stratospheric ozone on the transmission of UV radiation through the atmosphere is well

understood and has been extensively discussed in previous

Assessments (Herman and McKenzie et al., 1999). The

community now has access to better measurements and

radiative transfer models; and the effects of molecular

absorptions are now better quantified through field measurements and supporting model calculations (e.g.,

Burrows et al., 1999; Mohamed-Tahrin et al., 2001). New

evidence for increases of surface UV-B radiation due to

stratospheric ozone depletion in Antarctica, particularly

during the period of the ozone hole in spring, has been

reported from observations (Bernhard et al., 2004, 2006).

Summertime enhancements of surface UV were reported

also at northern high latitudes during episodes of low

ozone in the stratosphere caused by chemical destruction

and transport processes (Orsolini et al., 2003).

Surface UV radiation depends on the vertical profile of ozone and on the vertical profile of temperature

because of the temperature dependence of the molecular

absorption cross sections. McKenzie et al. (2003) showed

that differences in the shape of the vertical profile of ozone

amplify the seasonal differences in UV irradiance between

the northern (NH) and southern (SH) hemispheres. The

summertime UV irradiance at southern midlatitudes

greatly exceeds that at northern midlatitudes. Kazantzidis

et al. (2005) quantified the differences in surface UV irradiances that are calculated by RT models when substituting

the standard Air Force Geophysics Laboratory (AFGL)

ozone and temperature profiles with measured ones. The

differences in the UV-B irradiance are ~2-3% at small

SZA, increasing by a few percent at larger SZA. Krzy ́scin

(2004) reported that for large SZA, the UV irradiance at

the surface is more sensitive to the ozone changes in the

midstratosphere than to changes in the lower stratosphere,

while the opposite occurs for small SZA. Under the current state of the atmosphere, the dependence of surface

UV on the vertical profiles of temperature and ozone is

relatively weak, but it might become stronger in the future,

if the temperature of the lower stratosphere (where the

bulk of ozone resides and most of the UV absorption

occurs) changes (see Chpater 5, Section 5.3). Generally,

the effect of changes in the ozone profile on surface UV

radiation is larger than that of changes in the profile of

temperature.

7.2.3 Clouds

Clouds dominate any other atmospheric variable as

a source of surface UV variability, resulting in either a

reduction or an increase in UV radiation. Indirectly, clouds

may affect absorption of UV by other atmospheric constituents, while their variability usually masks the effects

of other variables. For example, clouds are the major

factor that limits the detectability of ozone-induced trends

SURFACE ULTRAVIOLET RADIATION

7.5
in UV radiation (den Outer et al., 2005; Glandorf et al.,

2005). The higher the variability of clouds, the longer the

record of ground-based UV data required for detecting an

ozone-induced trend in UV radiation. Cloud effects on

UV radiation are complex, thus restricting our ability to

describe and predict accurately the UV radiation field at

the ground. In practice, the necessary parameters needed

to calculate local cloud effects are rarely available. Even

if they were, the complexity of cloud geometry would

require three dimensional (3-D) model calculations.

Climate modeling suggests that the temporal and spatial

distribution of clouds is likely to change in the future

(IPCC, 2001 and Section 7.5.2). If ozone changes are

small, the anticipated changes in cloudiness will have a

much greater impact on UV radiation than changes in

ozone. Despite the inherent difficulties in investigating

the effects of clouds on surface UV, new case studies have

been published and have broadened our understanding.

Results from different studies of cloud effects (e.g.,

Calbó et al., 2005) can be compared using the cloud modification factor (CMF), which is defined as the ratio

between the measured UV radiation in a cloudy sky and

the calculated radiation for a cloudless sky. Typical CMF

values for overcast skies range from 0.3 to 0.8, depending

both on cloud type and characteristics (Cede et al., 2002;

Calbó et al., 2005). Den Outer et al. (2005) have investigated the effects of clouds on UV and visible solar radiation in the Netherlands by comparing measurements with

model calculations. These studies confirm that clouds

attenuate UV radiation by 15-45% less than total solar

radiation, as a result of Rayleigh scattering that redirects

toward the surface more efficiently the UV part of the radiation reflected by the clouds. Nichol et al. (2003) found a

tendency toward greater cloud attenuation with increasing

solar zenith angle.

Clouds may also have an enhancement effect, manifested by increased UV irradiance at the surface when the

solar disk is unoccluded. Recent studies show enhancements of up to 25% and reconfirm earlier findings that the

enhancements are significant (Herman and McKenzie et

al., 1999; Kerr and Seckmeyer et al., 2003) and can last

for hours (Cede et al., 2002). The enhancements were

found to be most pronounced for large cloud cover of 5 to

7 octas (Cede et al., 2002) and were smaller in the UV than

in the visible and infrared (Pfister et al., 2003).

Simultaneous occurrence of a small solar zenith angle and

a cumulus cloud close to the Sun’s position in the sky at a

high-altitude inter-tropical desert region produces very

high erythemal doses. Piacentini et al. (2003) reported

mean percentage increases due to the cloud effect of 6%

for UV and 12% for total solar irradiance, relative to no

clouds near the Sun. Cloud enhancements in UV actinic

flux can be up to 40% over clear-sky values when the solar

disk is unoccluded (Crawford et al., 2003). When the solar

disk is occluded, reductions in actinic flux below the cloud

appear to vary inversely with cloud fraction in some

instances. Wavelength dependence has been observed

under broken clouds, with shorter wavelengths generally

exhibiting lower variability in irradiance for both enhancements and reductions (e.g., 20-30% less variability at 320

versus at 420 nm). For actinic fluxes, the wavelength

dependence due to clouds is stronger for cloud-induced

enhancements than for cloud-induced reductions with

respect to clear skies. At large SZA, a wavelength dependence appears to occur even for overcast conditions. The

observed behavior (reductions, enhancements, and wavelength dependences) could be explained through the difference in the impact of clouds on the direct and diffuse

components of actinic flux. Integration over longer

periods results in weaker wavelength dependence, suggesting that differences in wavelength response may be

localized phenomena with impacts that are minimized

when integrated over the surrounding area.

Simultaneous measurements of actinic flux, irradiance, and aerosol and cloud properties were made from

four ground stations and by aircraft (Kylling et al., 2005).

The actinic flux above the cloud is increased by between

60-100% when compared with a cloudless sky, with the

largest increase occurring for the optically thickest cloud.

Similarly, the actinic flux below the cloud is decreased by

about 55-65%. Just below the cloud top, the downwelling

actinic flux reaches a maximum.

Mie scattering from cloud particles interacts with

Rayleigh scattering in the atmosphere and produces a complex wavelength dependence in the top-of-the-atmosphere

reflectances measured by satellite instruments that operate

in the ultraviolet part of the spectrum (Ahmad et al., 2004).

It was also shown that clouds can perturb the absorption

by tropospheric ozone in complex ways that cannot be

explained by models that treat them as reflecting surfaces

rather than as volume scatterers. Winiecki and Frederick

(2005) confirmed that the cloud transmission in the ultraviolet decreases with decreasing wavelength as a consequence of increasing absorption by tropospheric ozone in

the cloud, and they suggested that this mechanism results

in enhanced production of excited-state oxygen atoms,

O(1D).

The UV-B sky radiance distribution under partial

and overcast stratified cloud fields was studied using

azimuthally averaged radiance measurements at Hobart,

Australia (Kuchinke et al., 2004) and modeling. It was

shown that inhomogeneities within realistic cloud fields

have a significant effect on the radiation distribution at the

ground, even on an overcast day.

SURFACE ULTRAVIOLET RADIATION

7.6
7.2.4 Aerosol Scattering and Absorption

The influence of aerosols on the transmission of UV

radiation has important consequences for stratospheric and

tropospheric photochemistry, human and plant biology,

remote sensing of column ozone, and surface UV mapping. There is a large variation in aerosol properties in

both space and time caused either by natural or anthropogenic activities, and quantifying their effect remains a

challenge. Uncertainties in aerosol particle microphysical

data, particularly the particle number concentration, cause

variability of up to 10% in predictions of actinic flux and

UV irradiance (Früh et al., 2003). Advances in quantifying the effect of aerosols are being made from various

fronts: laboratory and field measurements of aerosol

optical properties in the UV, detection of aerosol amounts

and optical properties from the ground and space, and theoretical modeling of scattering and absorption of UV radiation. Detecting and predicting aerosol absorption

properties in the UV have received particular attention

in recent years.

The optical properties of aerosols in the ultraviolet

have been measured for different aerosol types in the laboratory and in the field. Black carbon/soot aerosols were

characterized during the Aerosol Interactions and

Dynamics in the Atmosphere (AIDA) Soot Aerosol

Campaign (Schnaiter et al., 2003) using an optical technique based on the integrating sphere method (Mogo et

al., 2005). Biomass burning aerosols were studied using a

long path extinction spectrometer (LOPES) (Schnaiter et

al., 2005) and fluorescence from a UV lidar (Immler et al.,

2005). Desert dust, Arctic industrial pollution, and marine

aerosols were investigated by Wetzel et al. (2003).

Vaglieco et al. (2002) and Kirchstetter et al. (2004) found

that organic carbon from motor vehicles and biomass

burning contributes significantly to absorption of UV (and

visible) radiation, although Myhre and Nielsen (2004)

found that absorption of UV by organic acids is not significant. Kirchhoff et al. (2002) found a wavelength dependence of aerosol optical depth (AOD) in the UV-B that

opposes the Angstrom law; they examined the possible

influence of additional UV-absorbing species, such as

formaldehyde, but found no influences on the AOD

retrieval. On the other hand, Arola and Koskela (2004)

studied various sources of errors in direct irradiance measurements due to the field of view of Brewer spectroradiometers that can partly explain this behavior. The

properties of desert dust in the UV are also under debate.

Some measurements have indicated enhanced absorption

in the UV compared with the visible (Mattis et al., 2002;

Meloni et al., 2004).

The optical properties of aerosol optical depth,

single scattering albedo (SSA), and phase function are

used to quantify the transmission of UV through turbid

atmospheres and may be obtained by both satellite and

ground radiometric data. Direct measurements of aerosol

absorption present a great challenge even for ground-based

remote sensing techniques, and the uncertainties have been

recently discussed (e.g., Petters et al., 2003; Bais et al.,

2005b; Goering et al., 2005; Krotkov et al., 2005a, 2005b).

Using spectral direct irradiance measurements, AOD was

retrieved by Lenoble et al. (2002) in the Bavarian Alps

with absolute uncertainties of 0.03 to 0.05 and by Gröbner

and Meleti (2004) in Ispra, Italy using a Brewer spectrophotometer. During the period 1991-2002, the AOD at

320 nm in Ispra varied between 0.05 and 2, showing a pronounced seasonal variability with high values (~0.6) in

spring and summer and low values (~0.3) in winter. The

AOD in the UV-A region determined from global and diffuse irradiance in the French Alps showed seasonal variations between 0.05 and 0.2 (Lenoble et al., 2004a).

Techniques for retrieval of SSA have been developed using

multiangle measurements of sky radiance (e.g., Dubovik

et al., 2002; Qin et al., 2002), direct, diffuse, and global

irradiance in conjunction with RT models (e.g., Petters et

al., 2003; Bais et al., 2005b; Goering et al., 2005; Krotkov

et al., 2005a, 2005b), and by combining Raman lidar, total

ozone, and UV-B irradiance measurements with a RT

model (Balis et al., 2004). The aerosol SSA is reported

routinely by the Aerosol Robotic Network (AERONET)

but only at visible wavelengths, so SSAat UV wavelengths

can only be inferred by extrapolation. Petters et al. (2003)

used an ultraviolet multifilter rotating shadowband

radiometer (UVMFRSR) and RT model calculations to

estimate the SSA, which ranged from 0.65 to 0.91 at 300

nm and 0.80 to 0.99 at 368 nm at Black Mountain, N.C.

Bais et al. (2005b) estimated the SSA by combining

Brewer irradiance measurements with a RT model and

found that the ratio of direct to diffuse irradiance gives the

highest accuracy compared with global irradiance, as long

as the aerosol optical depth at 340 nm is greater than 0.2.

From five years of irradiance measurements, the SSA in

Thessaloniki, Greece, was determined to be between 0.85

and 0.99 when the AOD is larger than 0.8, and to cover

the entire range from 0.64 to 0.99 for AOD between 0.2

and 0.8 (Bais et al., 2005b). The very low minimum SSA

values (~0.65) reported in these two studies may have

resulted from additional absorption by atmospheric gases,

for example by NO2 (Krotkov et al., 2005c) that was not

taken into account in the retrieval of SSA.

Recently, an inversion technique has been developed to convert the unique Total Ozone Mapping

Spectrometer (TOMS) spectral signature generated by the

SURFACE ULTRAVIOLET RADIATION

7.7
interaction of molecular scattering and aerosol absorption

into a quantitative estimate of AOD and SSA, with accuracies of ±30% and ±0.03 respectively (Torres et al., 2002,

2005).

New studies demonstrated the strong influences of

variations in aerosol concentration and composition on

long- and short-term variations in surface UV radiation:

Chubarova et al. (2002) for Moscow, Russia; Barnard et

al. (2003) for California, USA; Palancar and Toselli (2002)

for Cordoba, Argentina; Micheletti et al., (2003) for

Buenos Aires, Argentina; Ogunjobi and Kim (2004) for

Kwangju, South Korea; Latha et al. (2004) for Northern

India; Arola et al. (2003b) for Sodankylä, Finland, and

Thessaloniki, Greece; Kambezidis et al. (2005) for Athens,

Greece, and di Sarra et al. (2002) and Meloni et al. (2003a;

2003b; 2003c; 2005) for the island of Lampedusa in the

Mediterranean. Jaroslawski and Krzy ́scin (2005) showed

that aerosol forcing on the erythemally weighted UV at

Belsk, Poland, under clear skies is as important as the

forcing due to changes in ozone amount, both on longterm and short-term time scales.

The long-term variations of surface UV irradiance,

which are discussed in Section 7.4, may strongly be influenced by temporal changes in the optical properties of

aerosols, especially the SSA, which defines their absorption efficiency. In addition, the wavelength dependence

of the two main aerosol optical properties, AOD and SSA,

may result in different long-term variation of surface UV

radiation at short and long wavelengths, if the type of

aerosols over a particular site changes with time. Apart

from their natural sources, aerosols are also a consequence

of anthropogenic activities leading to atmospheric pollution, and thus they may change with time, depending on

the effectiveness of pollution abatement measures.

Finally, a reduction in SSA may result in less scattered

radiation available for absorption by atmospheric gases,

thus affecting indirectly the photochemistry in the troposphere.

7.2.5 Tropospheric Ozone and Other

Gaseous Air Pollutants

UV radiation can be efficiently absorbed by manmade or natural gases in the troposphere, and at the same

time it controls the concentration of some of them. For

example, tropospheric concentrations of ozone may be

significantly affected by changes in photodissociation

rates induced by stratospheric ozone changes (Isaksen et

al., 2005), which in turn affects the intensity of UV radiation at the surface. Under certain conditions, the concentration of gases resulting from photochemical reactions

can increase simultaneously resulting in additional absorption of UV radiation near the Earth’s surface. On the other

hand, UV radiation is regulated by stratospheric ozone, as

well as by the influence of aerosol and cloudiness. Thus,

UV radiation interacts with tropospheric gases and

aerosols in complex ways through a variety of atmospheric

processes.

Tropospheric absorbers of UV radiation include

ozone (O3), sulfur dioxide (SO2), nitrogen oxides (NO2,

dinitrogen pentoxide (N2O5), NOx), some bromine and

chlorine compounds (e.g., chlorine dioxide (OClO),

molecular chlorine (Cl2), dichlorine monoxide (Cl2O),

bromine monoxide (BrO), hypobromous acid (HOBr)),

and the oxygen dimer (O2-O2) (Atkinson et al., 1997;

Sander et al., 2003). In addition, many organic species

can also absorb UV radiation, including different types of

aldehydes (i.e., formaldehyde, benzaldehyde, acetaldehyde), acetone, peroxyacetyl nitrate (PAN), propanal,

glioxal, nitrated aromatics, certain organic acids, and

others (Atkinson et al., 1997; Jacobson, 1999; Sander et

al., 2003; Myhre and Nielsen, 2004). The abundance of

most of these substances in the troposphere usually is not

large enough to produce measurable effects on surface UV,

except near the sources of emission in polluted areas, or

during natural hazards, like forest fires, volcanic eruptions, etc. For example, over large industrial centers on

the eastern coast of the United States, the additional photochemically-produced ozone in the lower 2 km was up to

14 parts per billion by volume (ppb) (corresponding to a

column contribution of 2.5 Dobson units) (Cooper et al.,

2005) and was occasionally much higher (Jacobson, 1999).

The content of NO2 in the lower troposphere over polluted

areas can be up to 1-2 DU, or even more (Richter and

Burrows, 2002; Richter et al., 2005; Wang et al., 2005a;

Cede et al., 2006a). Tropospheric SO2 can be high in

regions with volcanic activity and over industrial areas,

with concentrations of 2 DU or more, as revealed from

SCIAMACHY and ground measurements (Wang et al.,

2005b). Finally, formaldehyde concentrations can exceed

0.6 DU (Heckel et al., 2005). The attenuation of UV radiation at the surface from these gases can be important, particularly if they coexist in the same area. Table 7-1 shows

the effects of the main air pollutants in the boundary layer

on the attenuation of UV irradiance, expressed by the

absorption sensitivity (AS) parameter, defined as the relative change in UV irradiance due to 1 DU change in gas

concentration. The absorption sensitivity of NO2 and SO2

can be even higher than for O3 for the same change in concentration.

There is an increased awareness of the possible

roles of these tropospheric trace gases on surface UV radiation, particularly in more polluted environments

(Chubarova et al., 2002; Koronakis et al., 2002; Palancar

SURFACE ULTRAVIOLET RADIATION

7.8
and Toselli, 2002; Chubarova, 2004). Monitoring over

several years of UV-B and total solar irradiance in Moscow

and at a site located 50 km upwind to Moscow has revealed

that in more than 75% of the cases, the monthly-mean UVB values at the rural site are higher, with differences

reaching 18% in winter and 9% in summer (Chubarova,

2002). For Moscow conditions, ozone and nitrogen

dioxide are the most important gases in tropospheric UVB absorption. For example, the observed attenuation of

UV-B irradiance due to NO2 is up to 6%, being ~2-3% on

the average (Chubarova, 2004). During extremely high

pollution episodes caused by forest fires near Moscow in

2002, the UV reduction only due to NO2 was about 11-

14%. Koronakis et al. (2002) reported reductions in UVA irradiances of similar magnitude due to increased concentrations of tropospheric ozone and nitrogen oxides in

Athens, Greece. Progress has also been made, using suites

of ground-based instruments, to separate the effects of

extinctions by NO2 and aerosols (Krotkov et al., 2005c).

The future concentrations of tropospheric gases will

depend on increased population, international agreements,

such as the Kyoto Protocol and its successors, and the continued affordability of fuels. These changes in concentrations can affect UV radiation in various ways. Some gases,

such as tropospheric ozone, can have a direct influence

through absorption of UV. Others, such as nitrous oxide

(N2O) and methane (CH4), can influence UV through their

role in stratospheric chemistry, which results in changes

to ozone concentrations. Further, all the greenhouse gases

contribute to cooling of the stratosphere, which affects

ozone chemistry, and to warming at the surface. The latter

leads to changes in many variables that are capable of

modifying directly the UV radiation at the surface. These

include changes in glaciation, ice/snow cover, clouds,

rainfall, and run-off. Presently it is difficult to predict the

future concentrations of tropospheric gases that attenuate

UV radiation (Isaksen et al., 2005), and consequently their

importance for surface UV in the future. It is clear, however, that UV is reduced under polluted atmospheric conditions and, in some cases, these reductions may balance,

to a certain degree, the UV enhancements produced by

stratospheric ozone depletion.

7.2.6 Surface Albedo

UV radiation is enhanced by surface albedo, particularly over strongly reflecting surfaces, such as snow, ice,

or sand, and even more in the presence of cloud layers that

trap the reflected radiation between the clouds and the surface. The wavelength-dependent scattering and absorption processes control the enhancement of UV radiation

by surface albedo, and hence they determine the area

around a site that can contribute to these enhancements.

Changes in surface albedo due to global warming may thus

become important for future surface UV radiation,

through, for example, reduction of snow- and ice-covered

areas, deforestation, or expansion of deserts. New studies

have focused on quantifying the effects of high-albedo

surfaces on UV irradiance under both clear skies and

cloudy conditions, and on estimating surface albedo by

combining measurements and modeling. Increases in erythemal irradiance of up to ~20% were reported in these

studies.

Variations of erythemal and UV-A irradiances due

to surface albedo changes were investigated during 15

cloudless days at Briançon (French Alps) in February 1998

with variable snow cover and a small snowfall in the

middle of the campaign. The enhancement of erythemal

SURFACE ULTRAVIOLET RADIATION

7.9

Table 7-1. The absorption sensitivity (AS) of different integrals of UV irradiance due to a 1-DU change

in the column of O3, SO2, and NO2 in the lower troposphere. Calculations are for local noon in January and

July at 45°N, assuming 300 DU of ozone column (X) above 2 km. For a different ozone column, the absorption sensitivity for O3 and SO2 can be approximately estimated from the expression: , whereas

NO2 is practically unaffected. Adapted from Chubarova (2006).

Absorption Sensitivity (%/DU)

O3 NO2 SO2

July January July January July January

Erythemal irradiance −0.6 −0.4 −0.8 −1.8 −1.0 −0.9

UV-B (280 - 315 nm) −0.3 −0.4 −0.8 −1.5 −0.6 −0.8

UV-A (315 - 400 nm) Negligible −1.6 −3.0 Negligible

UV (280 - 400nm) Negligible −1.6 −2.9 Negligible

ASX 5 AS300

300

X
irradiance was found to change from 15% to 5% (7% to

5% for UV-A) as a result of changing surface albedo, with

a jump to 22% (15% for UV-A) after the snowfall. The

corresponding values of the effective surface albedo are

respectively 0.4, 0.1, and 0.5 for erythemal, and 0.25, 0.1,

and 0.4 for UV-A irradiances (Smolskaia et al., 2003).

Measurements at the same station in winter/spring 2002

showed that the enhancement of surface UV irradiance

depends more strongly on the snow distribution around

the site than on the actual topography. The maximum

enhancement of erythemal irradiance (~22%) for clear

skies agrees with results of a 3-D RT model, assuming a

snow albedo of 0.3 between the snow line and the tree line

and of 0.8 above the tree line. Retrieving an effective surface albedo from the enhancement of UV irradiance measurements is challenging, because small uncertainties in

the enhancement (e.g., ±2%) lead to large uncertainties

(±0.05) in effective albedo (Lenoble et al., 2004b).

Quantitative estimates of the enhancement of surface UV

irradiance due to surface albedo changes were derived also

for Sodankyla, Finland, by analyzing a time series of

monthly erythemal dose measurements (Arola et al.,

2003b). It has been shown that for this location the

enhancement is largest in May, being 21% at its maximum

and 7% on average. 

The effect of surface albedo on UV sky radiance

was investigated at the High Alpine Research Station

Jungfraujoch, Switzerland, 3576 meters above mean sea

level (amsl), located at the boundary between snowcovered and snow-free terrain (Huber et al., 2004). Toward

the northwest, the effect of snow-free terrain causes a significant reduction of UV sky radiance (between 20% and

40% depending on SZA) relative to a modeled simulation

assuming a homogeneous snow albedo of 0.4. Finally, at

Neumayer, Antarctica, it was found that UV zenith sky

radiance may increase by more than a factor of 2, relative

to clear sky, due to the combined effect of clouds and high

snow albedo (see example in Figure 7-2) (Wuttke and

Seckmeyer, 2006). On the contrary, no change has been

observed in the total irradiance, which was ~670 W m-2,

both for overcast and clear-sky conditions.

Using model calculations, Parisi et al. (2003) estimated the increases in daily erythemal UV exposures on

horizontal and on inclined planes over water, concrete,

and sand relative to grassland. Under cloud-free skies, the

estimated additional daily UV exposure on a horizontal

plane was up to 3% higher over sand relative to grassland,

while on a vertical plane the enhancement was up to 11%,

for the range of SZA in the Australian sub-tropical location of this study.

Finally, measurements of spectral surface albedo

from aircraft over flat arable land in East Anglia, Great

Britain, show increasing albedo with increasing wavelength, from about 2% in the UV-B to about 5% at 500 nm

(Webb et al., 2004). Effects of scattering and absorption

within the layer beneath the flight level were removed by

nonlinear extrapolation of the airborne albedo measurements to the ground (Wendisch et al., 2004).

7.2.7 Altitude

UV radiation increases with altitude, because of

reduction of scattering and absorption by the less-dense

overhead atmosphere, and, at high altitudes, because of

additional reflections from the surface and the clouds

below. At high altitudes, the surface is usually covered

with snow, thus irradiance increases further due to

enhanced surface albedo. As these processes are wavelength dependent, the increase of UV radiation varies also

with wavelength.

New studies have investigated the altitude effect on

UV-B irradiance using simultaneous measurements at different altitudes. The reported increases in erythemal irradiance are 7% per kilometer in Bolivia (Zaratti et al.,

2003), 10.7% per kilometer in the Swiss Alps (Schmucki

and Philipona, 2002), 7-16% per kilometer in Germany

and between 5 and 23% per kilometer at different altitude

pairs in Bolivia (Pfeifer et al., 2006), and ~11% per kilometer in the western Himalayas (Singh and Singh, 2004).

The latter study also reported increases in direct and difSURFACE ULTRAVIOLET RADIATION

7.10

Figure 7-2. Spectra of zenith radiance measured

under overcast and cloudless skies at the German

Antarctic Neumayer Station (70.65°S, 8.25°W) on 19

December 2003 and 22 January 2004, respectively.

The overcast spectrum exceeds the clear-sky spectrum by up to a factor of 10. In addition, the wavelength with maximum radiance is shifted from the UV

(330 nm) for the clear-sky case to the visible (450

nm) for the overcast situation. Adapted from Wuttke

and Seckmeyer, 2006.
fuse irradiance, respectively, of 17.4 and 8.5% per kilometer. The variability in the estimates derived by these

studies confirms that the altitude effect of UV irradiance

is not linear, depending strongly on atmospheric and surface parameters, and varies also with wavelength.

Piacentini et al. (2003) measured extremely high

values of solar irradiances at Cerro Cruz Azul (3900 m

amsl) at the Puna of Atacama high-altitude inter-tropical

desert, Argentina, under conditions favoring enhancements from reflections at the cloud edges and from

multiple scattering in the cloud layer. The maximum shortwave (300-3000 nm) irradiance was 1528 W m-2, about

8% higher than the solar constant, adjusted to the actual

Sun-Earth distance, while the maximum UV irradiance

integral (295-385 nm) was 69.5 W m-2 at 2.4° SZA.

Seasonal variations of the altitude effect on UV

radiation are mainly influenced by changes in solar elevation, albedo, and turbidity. Allen and McKenzie (2005)

measured with personal dosimeters the UV exposure of

skiers at the ski field of Mount Hutt, New Zealand (2080

m amsl), and found the UV Index maxima greater by ~23%

in mid-October and ~30% in mid-September than at sea

level at the same time of year. Thus the exposure depends

strongly on season.

7.2.8 UV Irradiances on Inclined Surfaces

Erythemal UV irradiance incident on a horizontal

surface is not always the best way of estimating the real

dose received by humans or animals, which is usually

much higher, considering that nonhorizontal parts of their

bodies may receive direct solar radiation at smaller angles

of incidence. As solar radiation measurements are referenced traditionally on a horizontal surface, often personal

UV radiation exposure studies and anatomical distributions of UV radiation relate the measured relative exposures to ambient UV radiation on a horizontal surface. A

dataset of more than 3,000 days of irradiance measurements at 27 different orientations and on a horizontal plane

reveals that irradiances on inclined surfaces are up to 1.7

times higher at the Schneefernerhaus high-altitude observatory (2650 m amsl) and 1.4 times higher in Munich (530

m amsl) compared with irradiances on a horizontal plane

at the same locations (Oppenrieder et al., 2004, 2005).

Simulations of these effects with a RT model were found

to be in good agreement with measurements (Mech and

Koepke, 2004). From clear-sky irradiance measurements

with different detector inclinations in the plane of the Sun

at Rosario, Argentina, it was found that the highest intensity of erythemal irradiance does not occur when the

detector is pointed directly to the Sun, but at an angle larger

than the solar elevation (i.e., closer to zenith). This

enhancement, which is stronger at large SZA (~30% at 80°

SZA) and diminishes at solar zenith angles smaller than

~50°, is attributed to differences in the contribution of diffuse and direct irradiances (Piacentini and Cede, 2004).

The irradiance incident on inclined planes depends more

strongly on ground reflectivity than irradiance on a horizontal plane (Koepke and Mech, 2005). Maximum

increases of 57% in erythemal UV irradiance from horizontal to normal incidence were calculated by Weihs

(2002) for altitudes above 3000 m.

7.2.9 UV Under Water, Ice, and Snow

Recent years have seen an expansion of research

into UV under water, ice, and snow, along the lines of

assessing variability in the context of climate change as

well as ozone depletion, and in discovering the wide variability in UV penetration that exists among various types

of aquatic ecosystems. Considerable new research has

emerged on the dependence of UV in a wide variety of

aquatic systems on bio-optical constituents such as dissolved organic matter (DOM), dissolved organic carbon

(DOC), which is part of DOM, chromophoric (i.e., optically active) DOM (CDOM), particulate organic carbon

(POC), and chlorophyll a, and on inorganic suspended particles. Table 7-2 summarizes many of these new results,

listing the transmission depth, Z(1%), of UV-B radiation

to 1% of the surface value for several aquatic systems. In

addition, satellite methods for globally mapping UV penetration into oceanic waters have progressed from developmental stages to more viable algorithms (Vasilkov et

al., 2005).

Häder et al. (2003) emphasize that for polar aquatic

ecosystems, factors such as climate warming and changes

in the sea ice season, surface albedo, and CDOM, may

influence exposure to UV-B radiation even more than the

springtime ozone depletion. As one example, Perovich

(2002) reported that 10 cm of snow cover reduces UV-B

transmittance by a factor of 40, thus protecting biota from

UV-B while allowing a substantially larger fraction of PAR

(i.e., Photosynthetically Active Radiation) to penetrate.

During Arctic spring conditions, Z(1%) for UV-B radiation incident on bare sea ice may be as deep as 1.75 m, but

with 5 cm of overlying snow, Z(1%) may be reduced to 30

cm. This implies a risk to non-UV-adapted organisms

within and under the ice, if snowmelt were to occur earlier

in the spring season as a result of climate warming.

Similarly, Cockell and Córdoba-Jabonero (2004), who

measured spectral UV transmittances through snow and

ice on the Mars Oasis, a deglaciated section of Alexander

Island, Antarctica, reported that an overlying snow depth

of 4 cm is sufficient to offset a UV-B increase caused by

SURFACE ULTRAVIOLET RADIATION

7.11
60% ozone depletion. The penetration depth of UV radiation into snow may have also relevance for photochemical

processes, through the release of NO2 during the photolytic

destruction of nitrate anions in the snow pack (Fisher et

al., 2005). The expected decrease in snow cover in many

areas due to climate warming suggests a strong linkage

between climate change and UV-B variability at high latitudes.

Johannessen et al. (2003) have developed an empirical method to parameterize UV attenuation and CDOM

absorption in oceanic waters using satellite ocean color

measurements (i.e., SeaWiFS (Sea-viewing Wide Fieldof-view Sensor) images of normalized water-leaving radiance). Their method is based on in situ measurements

from a variety of sources, including the Mid-Atlantic and

Georgia Bights, the Bering Sea, and inshore waters of the

Delaware and Chesapeake Bays. Vasilkov et al. (2005)

developed and validated a method for retrieving the radiation field in oceanic waters that uses backscatter UV

(BUV) measurements from TOMS and OMI in conjunction with SeaWiFS data and full radiative transfer computations of irradiance at depth. Ocean Raman scattering,

which is a function of absorption by particulate matter and

CDOM in the upper water column, contributes to the

filling-in of Fraunhofer lines in space-based BUV measurements over cloud-free open ocean (Vasilkov et al.,

2002). There is the potential to use BUV measurements

to estimate chlorophyll and CDOM concentration for

waters, where CDOM and particulate matter are well correlated. A BUV-based method would offer the advantages

of (1) good sensitivity at low chlorophyll concentrations,

and (2) a high-frequency spectral structure, thus being less

affected by uncertainties associated with absolute calibration errors. However, complete development of such a

method requires more in situ measurements of UV biooptical properties throughout the world’s oceans than

presently exist.

From spectral transmission measurements in relatively transparent oceanic waters (characterized by small

concentrations of particulate matter, such as silt, and

yellow substance), Piazena et al. (2002) found that UV-B

penetrates up to ~25% of the photic zone in these waters,

while UV-A penetrates up to 93%. Rasmus et al. (2004)

found that Southern Ocean waters are quite transparent to

UV-B radiation during January, and that under an unperturbed ozone layer, 305 nm doses influence the first 4.0-

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7.12

Table 7-2. Values of the 1% transmission depth (ZUV-B(1%)) for solar radiation at single wavelengths and

wavelength bands in the UV spectral region from recent studies in a variety of aquatic systems.

Location ZUV-B(1%) (m) Wavelength Range Source

N. Michigan stream, DOC = 2.0 mg l-1 0.45 305-320 nm Frost et al., 2005

N. Michigan stream, DOC = 24 mg l-1 0.04 305-320 nm Frost et al., 2005

Boreal forest lake, DOC = 3.0 mg l-1 2.00 300-320 nm Donahue et al., 2003

Boreal forest lake, DOC = 6.7 mg l-1 0.40 300-320 nm Donahue et al., 2003

Humic lake, Finland, DOC = 5.9 mg l-1 0.32 310-320 nm Huovinen et al., 2003

Ultraoligotrophic high Arctic lake 7.00 308 nm Ørbæk et al., 2002

Lake Huron, central 8.40 325 nm Smith et al., 2004

Lake Erie and Lake Ontario, central 2.50 325 nm Smith et al., 2004

Lake Huron tributaries 0.19 325 nm Smith et al., 2004

Oceanic water, Jerlov type OI 30.50 310 nm Piazena et al., 2002

Oceanic water, Jerlov type OIA 24.70 310 nm Piazena et al., 2002

Oceanic water, Jerlov type OIB 20.60 310 nm Piazena et al., 2002

Oceanic water, Jerlov type OII 12.40 310 nm Piazena et al., 2002

Oceanic water, Jerlov type OIII 6.60 310 nm Piazena et al., 2002

Atlantic water, Nordic Seas 31.00 310 nm Aas et al., 2002

Northern Barents Sea 11.80 310 nm Aas et al., 2002

Northern Greenland Sea 8.80 310 nm Aas et al., 2002

Norwegian Coastal Current 7.80 310 nm Aas et al., 2002

Kara Sea 3.60 310 nm Aas et al., 2002

North Water Polynya, pre-bloom 9.40 320 nm Vasseur et al., 2003

North Water Polynya, bloom 7.90 320 nm Vasseur et al., 2003
9.5 meters of the upper mixed layer, between 51°S and

65°S. They determined that the levels of UV and PAR

were high enough that Southern Ocean phytoplankton is

often not light limited.

Aas et al. (2002) synthesized several decades of in

situ spectral UV attenuation measurements in Arctic

marine waters and found considerable variability, related

to particular classes of Arctic surface water, including

Atlantic water, polar water influenced by river runoff, or

Arctic water found in central gyres. Barron and Barron

(2005) studied the waters of Prince William Sound, Alaska,

in the context of glacial retreat due to climate warming,

and reported that glacial melt water and suspended rock

sediment cause a tenfold reduction in UV transmission at

depth, compared with waters in the same basin distant from

glacial influences.

In freshwater ecosystems, optical properties can

vary widely due to factors such as water renewal time, the

proportion of wetlands in the catchment, and the influence

of tributaries, thus influencing the transmission of UV

radiation in the water column (e.g., Frenette et al., 2003;

Kjeldstad et al., 2003; Bracchini et al., 2005; Erga et al.,

2005). Smith et al. (2004) found for the Great Lakes in

the USA, much higher UV attenuation by CDOM in tributaries (89-94% of the attenuation) as compared with the

lakes (37-77%). Arctic freshwater ecosystems are often

characterized by low nutrient concentration and low DOC,

implying particular vulnerability to UV stress (Häder et

al., 2003). UV penetration in many boreal lakes is

increasing, as DOC concentrations decrease due to influences of climate change and acidification. In Scandinavian

humic lakes (60-63°N) and boreal forest lakes, DOC is a

strong regulator of UV-B transmission (Donahue et al.,

2003; Huovinen et al., 2003). In the shallowest aquatic

ecosystems, there appears to be strong attenuation of UV

radiation in the water column, correlated with DOM concentrations (e.g., Palen et al., 2002; Frost et al., 2005).

7.3 RESOURCES FOR STUDYING SURFACE UV

RADIATION

The variability of surface UV irradiance in time

and space is studied using measurements from groundbased instruments, as well as from estimates that are

derived from a combination of satellite radiance measurements and radiative transfer modeling. Significant

advancements have been reported in the last few years

in measurement and calibration procedures of instruments and in algorithms used both in ground-based and

satellite-borne platforms. Particular effort has been

devoted to the determination and quantification of the

errors and uncertainties involved.

7.3.1 Ground-Based Measurements

Quality control and quality assurance procedures

have received considerable attention in the last decade,

leading to significant improvements in instrument stability

and data quality. Many monitoring stations have adapted

widely accepted calibration and measurement procedures

and protocols. In addition to irradiance, spectrally

resolved measurements of actinic flux, direct irradiance,

radiance, and polarization are now frequently made.

7.3.1.1 SPECTRORADIOMETERS

Spectroradiometers are still the most reliable type

of instruments to measure solar UV radiation at the Earth’s

surface (Seckmeyer et al., 2001). In most cases, the measured quantity is the global irradiance, but in recent years,

spectral measurements of actinic flux and of direct irradiance are gaining particular attention, for more accurate

determination of photolysis frequencies, spectral aerosol

optical depth, or column densities of different atmospheric

constituents (e.g., Hofzumahaus et al., 2002; Webb et al.,

2002a; Houët and Brogniez, 2004; Lenoble et al., 2004a;

Kazadzis et al., 2005; Kylling et al., 2005).

Empirical and semi-empirical methods have been

derived to determine actinic flux and photolysis frequencies from irradiance measurements, in experimental campaigns where the derived products could be validated by

measurements (Webb et al., 2002a; Bais et al., 2003;

Hofzumahaus et al., 2004; Monks et al., 2004). For clearsky conditions, the conversion with empirical models

agrees better than 10% with measurements (McKenzie et

al., 2002; Webb et al., 2002b; Kazadzis et al., 2004;

Topaloglou et al., 2005); for cloudy conditions the agreement is within ~15%. Use of model calculations yields a

similar agreement under clear and overcast conditions, but

is less accurate under broken clouds (Kylling et al., 2003b).

A semi-empirical approach that combines measured

actinic flux and irradiance spectra with model calculations

results in similarly good agreement for all sky conditions

(Schallhart et al., 2004). Considering that long-term measurements of actinic fluxes and photolysis frequencies are

few and sparse, these methodologies constitute the only

way to assess the variability of these quantities and their

effects on tropospheric photochemistry in the last 10 to 15

years.

The distribution of actinic flux in the vertical

dimension and through a cloud layer was measured during

a field campaign and was well described by a RT model

(Kylling et al., 2005). The interpretation of such measurements and simulations is important for improving our

understanding of the radiative and chemical processes that

SURFACE ULTRAVIOLET RADIATION

7.13
occur in the troposphere, especially under and above the

clouds.

7.3.1.2 BROADBAND FILTER RADIOMETERS

Broadband radiometers represent instruments that

measure integrated irradiance over a large wavelength

range, such as UV-A or UV-B, but typically have been

designed to measure irradiance weighted by the action

spectrum for erythema as defined by the Commission

Internationale de l’Éclairage (CIE) (McKinlay and Diffey,

1987). These instruments are suitable for many applications because of their low cost and ease of operation in

remote locations, but they do have limitations in terms of

quality assurance and quality control (QA/QC). They are

a major source of ultraviolet radiation information for

larger geographical areas and have been in continual use

for longer periods than many other types of UV instruments (Nunez et al., 2002; Hicke et al., 2004; Kimlin et

al., 2005). The UV instrumentation working group of the

World Meteorological Organization (WMO) prepared a

report (Seckmeyer et al., 2005) which describes and

defines the specifications of this type of instruments, provides guidelines for their calibration and characterization,

and outlines QA/QC procedures based on current understanding of these instruments and UV science objectives.

These guidelines have been applied, for example, in the

Austrian UV monitoring network (Blumthaler, 2004). A

detailed uncertainty analysis of broadband radiometers

operating in Norrkoping has been presented by Josefsson

(2006).

A new method for calculating correction factors for

the non-ideal angular response of the broadband radiometers was developed using actual measurements of the

anisotropy of the UV-B sky (Kuchinke and Nunez, 2003).

The usual assumption of isotropic sky distribution overestimates the diffuse component by 4-6% for clear skies. 

Knowledge of the relative spectral response of erythemal radiometers is essential for correcting their measurements for column ozone and SZA variations.

Measurements of the response of a sample of instruments

at three different laboratories were found to agree within

20% with typical dynamic ranges of approximately 104

(Schreder et al., 2004). The dependence of the output

signal of Solar Light Co.SL501 erythemal radiometers on

ambient temperature and relative humidity was investigated (Huber et al., 2002, 2003). It was shown that despite

the internal temperature stabilization at 25°C, increasing

the ambient temperature by 20°C reduced the instrument’s

spectral response by up to 10% in the UV-B and by a factor

of two in the UV-A, depending on the instrument and its

internal relative humidity. By varying the relative

humidity from 9 to 70%, the instrument’s sensitivity was

reduced by 5% in the UV-B and by a factor of two in the

UV-A. There is yet no evidence that such effects occur

also in Yankee Environmental Systems, Inc. (YES) model

UVB-1 radiometers that are stabilized above 40°C. The

dependence of the SL501 radiometers on the ambient conditions calls for further investigation and quantification of

the effects on the actual measurements and of this

radiometer’s suitability for long-term studies.

7.3.1.3 NARROWBAND MULTIFILTER RADIOMETERS

Narrowband multifilter radiometers measure solar

radiation in several (typically 4 to 7) UV channels with

bandwidths ranging from 2 to 10 nm. They provide more

information than broadband radiometers and require less

maintenance than spectroradiometers. All channels can

be measured at high frequency, making the instruments

well suited for investigating short-term variation of spectral UV irradiance caused, for example, by rapidly

changing cloud conditions (Lovengreen et al., 2005). High

sampling rates also allow deployment on balloons for

measuring vertical profiles of UV (Kylling et al., 2003a).

For accurate calibration, knowledge of the spectral

response of the different channels is required (Høiskar et

al., 2003; Bernhard et al., 2005; Lakkala et al., 2005) and

corrections for the effect of ozone and SZA variation

should be applied (Diaz et al., 2005). Methods for proper

quality control of Norwegian Institute for Air Research

NILU-UV multichannel radiometers, operating under

harsh conditions in SH high latitudes, are described by

Lakkala et al. (2005), who have also reported large (up to

~35%) drifts in their sensitivity.

Data from multifilter radiometers allow the calculation of secondary data products such as reconstructed

high-resolution solar spectra, erythemal irradiance, total

ozone, and cloud transmission, with the aid of empirical

or RT models (Høiskar et al., 2003). Such products are

regularly included in the UV-monitoring programs of

Norway (Mikkelborg et al., 2000), the U.S. National

Science Foundation (NSF) (Bernhard et al., 2005), and the

U.S. Department of Agriculture (Davis and Slusser, 2005).

Feister et al. (2005) use a neural network technique to

reconstruct solar irradiance spectra. For solar zenith

angles smaller than 80°, total ozone derived from multifilter radiometers can have accuracy similar to data

obtained from Dobson spectrophotometers (Dahlback et

al., 2005). Seroji et al. (2004) demonstrated the feasibility

of calculating photolysis rates of O3, NO2, and formaldehyde (CH2O) from filter instrument data.

If the instruments are equipped with an automated

shadow band, direct solar irradiance can be retrieved,

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7.14
which allows the calculation of aerosol optical depth (e.g.,

Krotkov et al., 2005a) and the optical depth of thin clouds

(Min et al., 2004). By combining direct and diffuse measurements, single scattering albedo and other aerosol

parameters can be derived (Petters et al., 2003; Goering et

al., 2005; Krotkov et al., 2005b).

The improved technical characteristics of these

instruments and the possibility to derive additional products constitute their main advantages and make them

important supplements to spectral measurements of solar

irradiance in the global UV monitoring system.

7.3.1.4 ADVANCEMENTS IN INSTRUMENTATION

Further progress has been made toward uncovering

instrumental errors in solar UV radiation monitoring and

in the development of methods for improving the measurement techniques and the quality of the data.

A temperature dependence in the transmission of

polytetrafluoroethylene (PTFE), which is a widely used

material in the diffusers of UV spectroradiometers, has

been quantified by laboratory studies (Ylianttila and

Schreder, 2005). At ~19°C, the transmission of most types

of PTFE increases abruptly by 2-3%, whereas below and

above this point a temperature coefficient of about

−0.1%/°C is observed. The influence on UV irradiance

was quantified through comparisons of measured and

modeled UV irradiances at an unpolluted site in New

Zealand, where all necessary modeling input parameters

were available (McKenzie et al., 2005). These authors

concluded that all data measured at ambient temperatures

of less than 15°C are typically underestimated by ~2%.

The temperature sensitivity of PTFE diffusers introduces

an additional uncertainty in the UV measurements, especially at sites where the ambient temperature fluctuates

around this range of temperatures (i.e., ~15-19°C), which

may have also a seasonality. It is rather unlikely that longterm variations of surface irradiance at a particular site

may be influenced, unless the ambient temperature has

changed over a certain period from temperatures below

this threshold to temperatures above it. Finally, during the

calibration of the instruments, the diffuser is usually heated

by the lamp; thus normally its temperature should be above

this threshold.

Knowledge of the angular response error of a radiation instrument allows, with the aid of RT modeling, the

calculation of correction factors. A transportable unit that

was developed in the frame of the Quality Assurance of

Spectral Solar UV Measurements in Europe (QASUME)

project was used to measure the angular response of 11

spectroradiometers operating in Europe (Bais et al.,

2005a), thus allowing back correction of datasets recorded

at these sites. Angular response errors of ~2% can now be

reached in Brewer spectrophotometers by modifying their

entrance optics (Gröbner, 2003), compared with errors of

~10% found in the standard Brewer configuration (e.g.,

Bais et al., 2005a).

Due to the combined effect of two polarizationsensitive elements, the entrance window and grating, the

sensitivity of Brewer spectrophotometers for direct-Sun

and sky-radiance measurements changes with solar zenith

angle. This introduces systematic errors in Langley

extrapolations, aerosol optical depth retrievals, and aerosol

single scattering albedo retrievals. The need for appropriate corrections has been shown by Cede et al. (2006b).

New methodologies were proposed for improving the

measurements of O3, SO2, and NO2 columns with Brewer

spectrophotometers (Cede and Herman, 2005; Cede et al.,

2006a).

In recent years, diode array and charged-coupled

device (CCD) single monochromator instruments have

become available, allowing measurement of the full solar

spectrum from UV to infrared within seconds or milliseconds (Edwards and Monks, 2003; Jäkel et al., 2005).

However, in the UV-B range, these instruments suffer from

stray-light contamination, and many attempts are

underway to reduce this problem.

7.3.1.5 DATA QUALITY AND DATABASES

The importance of quality assurance of UV measurements has been recognized since the beginning of the

measurement records, although procedures for implementing quality assurance were developed later, mainly

through intercomparison campaigns (e.g., Bais et al., 2001;

Lantz et al., 2002; Metzdorf et al., 2003; Webb et al.,

2003). Most of the spectroradiometers deployed today are

radiometrically calibrated with tungsten halogen standard

lamps that are traceable to standards maintained by

National Metrology Institutes (NMI) (e.g., Gröbner and

Sperfeld, 2005; Wuttke et al., 2006). Differences between

the UV spectral irradiance scales realized and disseminated by different NMIs in the past are one of the most

critical problems in UV radiometry (Kübarsepp et al.,

2002; Metzdorf et al., 2003). Even standards of the same

Institute may disagree beyond their stated accuracy, and

deviations exceeding 4% in the UV have been observed

(Bernhard and Seckmeyer, 1999; Kiedron et al., 1999;

Gröbner et al., 2002). Alternative methods for reducing

the uncertainties in the calibration standards and for

improving the comparability of the irradiance scales of

different NMIs are in progress (Yoon et al., 2002; Metzdorf

et al., 2003; Sperfeld et al., 2003; Durak and Samadov,

2004).

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7.15
In the framework of the QASUME project, a transportable spectroradiometer was developed (Gröbner et al.,

2005; Gröbner and Sperfeld, 2005). On-site comparisons

with 26 spectroradiometers deployed at various locations

in Europe revealed irradiance differences ranging between

4 and 14% (Gröbner and Sperfeld, 2005). Similarly, a new

spectroradiometer to act as traveling standard for the

Network for the Detection of Atmospheric Composition

Change (NDACC) (formerly Network for the Detection

of Stratospheric Change, NDSC) was developed (Wuttke

et al., 2006).

The uncertainties in UV irradiance measurements

vary between different instrument types and stations,

depending on the availability of resources for achieving

accurate characterization of instruments and for applying

proper quality control procedures. These procedures are

improving with time following the progress of the relevant technology. It is therefore expected that measurements conducted in early stages of monitoring have larger

uncertainties, which in some cases are difficult to quantify. Different studies in the last 5-10 years have reported

uncertainties between 5-7% for global or direct spectral

irradiance measurements in the UV-B (Bernhard and

Seckmeyer, 1999; Bernhard et al., 2004; Kazadzis et al.,

2005; Bernhard et al., 2006), which are dominated by the

uncertainty of the calibration standards and, for global

irradiance, the cosine-error correction.

There are two major databases of spectral UV radiation measurements: the World Ozone and Ultraviolet

Radiation Data Centre (WOUDC) in Toronto, Canada

(http://www.woudc.org) with spectral irradiance data from

29 stations, and the European Ultraviolet Database

(EUVDB) in Helsinki, Finland (http://uv.fmi.fi/uvdb/)

containing data from 43 stations. A list of stations and currently available data are given in Appendix 7A. New techniques have been pursued to uncover errors in spectral

measurements at the databases EUVDB and WOUDC, and

tools have been developed (Slaper et al., 2004; Engelsen

and Kylling, 2005) for this purpose. In EUVDB, all data

are now automatically checked and flagged for their

quality.

Several UV monitoring networks maintain independent UV databases. The U.S. National Science

Foundation’s UV monitoring network includes qualitycontrolled spectra measured between 1988 and 2006 at

seven high-latitude sites located in Antarctica, South

America, Alaska, and Greenland (see Appendix 7A.3). A

new data edition named “Version 2” is currently in preparation; it will offer higher accuracy and additional data

products such as total ozone, effective albedo, and cloud

optical depth (Bernhard et al., 2004, 2006).

7.3.2 Satellite-Derived UV Estimates

Instruments onboard satellites provide global maps

of surface UV irradiance by combining backscattered radiance measurements with RT models. The accuracy of the

models is limited mostly by uncertainties in input parameters representing the atmosphere and the Earth’s surface.

Issues that require particular attention are the detection

and correction of long-term calibration drifts and the establishment of interpolation methods when combining different satellite datasets.

7.3.2.1 DIFFERENCES BETWEEN GROUND-BASED AND

TOMS-DERIVED UV IRRADIANCES

Surface UV retrievals from the measurements of

the Total Ozone Mapping Spectrometer (TOMS) instruments onboard different satellites (1978-2002) is probably

the only dataset that has been so extensively compared

with ground-based measurements. Recently, the TOMS

UV irradiance database (http://toms.gsfc.nasa.gov) has

been expanded to include five new products (noon irradiance at 305, 310, 324, and 380 nm, and noon erythemally

weighted irradiance), in addition to erythemal daily exposure, allowing direct comparisons with ground-based

measurements from UV spectroradiometers (Krotkov et

al., 2004).

Since the previous Assessment (WMO, 2003),

advances have been made in understanding the factors

causing differences between satellite- and ground-based

UV data. Sabburg et al. (2002) compared TOMS daily

erythemal doses on clear days against measurements at

four sites of the Brewer network of the U.S. Environmental Protection Agency, and concluded that TOMS UV

retrievals have a positive bias between 1.4 and 12.5%, with

an average of ~5%. Fioletov et al. (2004) produced two

UV Index climatologies over Canada and the United States

for the period 1980-1990 using TOMS data, and these were

validated against Brewer UV measurements from several

sites. With snow on the ground, TOMS retrievals were up

to ~60% lower, due to insufficient treatment of the snow

albedo. In summer, the TOMS UV estimates were 10 to

30% higher, caused most likely by air pollution and

absorbing aerosols. Differences between daily erythemal

irradiances derived by TOMS (version 7 and 8) and by a

Brewer spectrophotometer at Lampedusa have been found

to increase with increasing aerosol optical depth, often

associated with Saharan dust outbreaks (Meloni et al.,

2005). Arola et al. (2005) found positive bias of 20-30%

by comparing irradiance data over Thessaloniki, Greece,

and Ispra, Italy, with TOMS version 8 products.

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7.16
These studies confirmed that insufficient information on aerosols near the surface is responsible for the

observed biases in satellite UV irradiances. Arola et al.

(2005) and Krotkov et al. (2005b) have shown that knowledge of the absorbing properties of aerosols may reduce

appreciably these biases, and suggested that taking into

account even climatological estimates of AOD and SSA

would greatly improve the satellite UV retrievals. Aerosol

absorption corrections were found to improve TOMSderived UV over Moscow during the period MaySeptember (Chubarova et al., 2005). Erythemal irradiances recorded at eight stations in Argentina agreed with

TOMS estimates, which were first corrected for the effects

of aerosols, to within ±10% under snow-free conditions

(Cede et al., 2004). These newly reported differences are

much smaller than those previously reported (Kerr and

Seckmeyer et al., 2003). Therefore the use of satellitederived UV estimates without accounting for effects from

aerosols at locations with large aerosol abundances may

lead to false conclusions.

New approaches have been suggested to model the

effects of clouds over snow-covered terrain. The method

proposed by Arola et al. (2003a) improved the agreement

between TOMS-retrieved erythemal daily doses and

ground-based measurements over two Finnish and five

Canadian sites. Tanskanen et al. (2003) developed a

methodology by analyzing the TOMS Lambertian

Equivalent Reflectivity (LER) data using a moving timewindow technique, accounting also for autumn and spring

transitions during snow accumulation and snow melt.

7.3.2.2 UV RETRIEVAL ALGORITHMS AND RELATED

UNCERTAINTIES

Further studies were devoted to the development of

new algorithms for producing global coverage of surface

UV irradiances, involving information from other satellite

products. Apart from the effect of aerosols (discussed in

Section 7.3.2.1), the treatment of the cloud and surface

reflectivity by the UV-retrieval algorithms is still a challenging issue.

The influence of clouds is one of the major sources

of uncertainty in surface UV irradiances retrieved from

satellite instruments, particularly over areas with high surface reflectivity, and different methodologies to parameterize the cloud transmittance have been proposed.

Meerkötter et al. (2003) compared two independent

approaches using data from the AVHRR instrument

onboard the polar-orbiting NOAA satellites and from

MVIRI (METEOSAT Visible and Infrared Instrument)

onboard the geostationary METEOSAT satellite. The

latter offers particularly good temporal resolution, but was

expected to be problematic at high latitudes where their

study was focused. However, they found that METEOSAT

(or the new METEOSAT Second Generation (MSG) satellite) may be used to account sufficiently for cloud effects

in UV at latitudes up to 70°. By combining cloud information from METEOSAT, ozone column from TOMS and

GOME (Global Ozone Monitoring Experiment), and RT

modeling (Surface Irradiation Derived from a Range of

Satellite-Based Sensors-SIDES), UV climatologies over

Europe from 1984 to 2003 were generated by Verdebout

(2004a; 2004b) with spatial resolution of ~5 km.

Improvements to the Joint Research Center (JRC) algorithm SIDES were applied (Wuttke et al., 2003), resulting

in better agreement with measurements in the UV-B. A

new satellite UV algorithm for producing UV-B and UVA irradiances and the UV Index using the Surface and

Atmosphere Radiation Budget (SARB) products of Clouds

and the Earth’s Radiant Energy System (CERES) was

developed (Su et al., 2005). Differences when compared

with a high-resolution, multistream radiative transfer code

range from −10% to +4% for UV-B and UV-A irradiances,

and from −26% to +16% for the UV Index. Ciren and Li

(2003) derived a long-term (1983-2000) global climatology of surface UV-B and erythemal irradiance by combining TOMS ozone and International Satellite Cloud

Climatology Project (ISCCP) D1 3-hour reflectance measurements, thus accounting more efficiently for the diurnal

variability of clouds. On the other hand, Williams et al.

(2004) have shown that using the ISCCP climatology for

estimating surface UV irradiance results in significant bias,

suggesting instead to use the TOMS reflectivity at 360 nm.

The TOMS UV algorithm (Krotkov et al., 1998;

Herman and McKenzie et al., 1999; Krotkov et al., 2001)

was successfully applied to the new Ozone Monitoring

Instrument (OMI) onboard the Aura spacecraft for continued monitoring of ozone, other trace gases, and surface

UV irradiance (Levelt et al., 2006; Tanskanen et al., 2006).

OMI is a contribution of The Netherlands’s Agency for

Aerospace Programs (NIVR) in collaboration with the

Finnish Meteorological Institute (FMI) to the NASA Earth

Observing System (EOS) Aura mission and, among other

products, provides global maps of surface UV irradiance

(see example in Figure 7-3).

7.3.3 Advancements in UV Radiation

Modeling

Considerable progress has been made in radiative

transfer modeling, in particular on the interaction between

clouds and radiation, the impact of aerosols, and threedimensional radiative transfer. Important aspects for the

ultraviolet spectral region are: (1) the strong Rayleigh scatSURFACE ULTRAVIOLET RADIATION

7.17
tering and absorption by ozone; (2) the approximately

wavelength-independent optical properties of clouds and

their negligible absorption unless they are polluted by

absorbing aerosol; and (3) the importance of short wavelengths due to the strong biological and chemical effects

of UV radiation, despite their small contribution to the

total energy.

A variety of radiative transfer codes are freely available (see Appendix 7B.3) (Madronich and Flocke, 1997;

Evans, 1998; Ricchiazzi et al., 1998; Key, 1999; Koepke

et al., 2005; Mayer and Kylling, 2005; Rozanov et al.,

2005). In addition to being useful for calculating irradiance at the surface, most of them can be used to calculate

other quantities, such as radiance and actinic flux, also at

different altitudes. Some (e.g., LibRadTran) allow calculations throughout the whole solar and thermal spectral

range.

SURFACE ULTRAVIOLET RADIATION

7.18

OMUVB Erythemal Daily Dose

180 120 60 60 20 180

LONGITUDE, degrees

60

60

LATITUDE degrees

10

11

12

180 120 60 60 20 180

LONGITUDE, degrees

60

60

LATITUDE degrees

10

11

12

Figure 7-3. Global distribution of the cloud-corrected daily integral of erythemal irradiance in kJ m-2 for 28

June 2005 (upper panel) and 28 December 2005 (lower panel) derived from OMI measurements.
Radiative transfer is used in: (1) research applications, such as the interpretation of campaign data, requiring

the highest degree of sophistication including, for

example, three-dimensional solvers or spherical geometry; (2) routine applications, such as the quality control

of observations or provision of cloudless-sky reference

data to routinely estimate the influence of clouds from

observed spectra; or (3) applications where large numbers

of calculations are needed, such as the computation of photolysis frequencies in chemistry-transport models (CTMs)

or the calculation of global time series of UV data. For

most routine applications, the above-mentioned freely

available codes are suitable, although they require significant computational power. The use of fast parameterizations like SMARTS2 (Gueymard, 2001) or lookup-table

based approaches like FASTRT (Engelsen and Kylling,

2005) may be more appropriate whenever a large number

of modeled spectra are required. The calculation of photolysis frequencies in chemistry or chemistry-transport

models is probably one of the most demanding applications, in terms of the number of spectra to be calculated,

requiring specially tailored solutions like Fast-J (Wild et

al., 2000). Finally, statistical methods have been derived

by relating a large number of observations at one or a few

sites to a set of ancillary measurements, such as total

ozone, cloud cover, or pyranometer observations (e.g.,

Fioletov et al., 2001; Schwander et al., 2002). While these

are usually very good in predicting the radiation for a given

location, it may be problematic to transfer the results to

other locations or to calculate radiation for time periods

other than the one used to build the statistical model.

After numerous model-model or model-measurement intercomparisons, the confidence in the results of

radiative transfer codes in the UV is high if the input

parameters are chosen carefully. Model-model intercomparisons have shown that different radiative transfer codes

may agree to well below the 1% level if they are operated

with identical optical properties. This is even true for complex three-dimensional codes, as has been shown in the

Inter-comparison of 3-D Radiation Codes (I3RC) (Cahalan

et al., 2005), where a core group of radiative transfer codes

agreed to within 0.1-1% for a variety of complex cases,

neglecting polarization, spherical geometry, and inelastic

scattering.

Uncertainty of model calculations can have different sources. First, uncertainties in parameters like extraterrestrial irradiance (Gueymard, 2004) and the absorption cross section of ozone (Orphal and Chance, 2003)

translate directly to uncertainties in irradiance. The largest

source of uncertainty is usually the lack of knowledge of

the actual state of the atmosphere or surface, in particular

for cloudy, polluted, or partially snow-covered conditions.

Clouds pose the most challenging modeling problem, due

to their complex geometry and high variability in space

and time. The effect of clouds on radiation is basically

understood (at least for spherical droplets) and can be simulated with a 3-D radiative transfer code. For theoretical

case studies, various cloud models have been used,

including simple geometrical shapes, statistical assumptions (Cahalan et al., 1994), cloud-resolving models

(Trautmann et al., 1999; Brasseur et al., 2002), and highresolution observations by satellites and aircraft

(Meerkötter and Degünther, 2001; Scheirer and Schmidt,

2005). However, the simulation of an actual situation is

complicated because a sufficiently complete description

of the clouds is generally not available. Information such

as cloud cover or pictures by sky cameras give a first hint

but are not sufficient to simulate the irradiance at the

ground. The situation for aerosols is similar; while the

optical depth is easily obtained by observations, the irradiance is critically determined by the single scattering

albedo, which is not easily measured (see Section 7.2.4).

Sometimes, the missing information is replaced by an

effective parameter, such as an effective single scattering

albedo (Kylling et al., 1998; Kazantzidis et al., 2001;

Petters et al., 2003; Bais et al., 2005b). If more observations are available to constrain the actual situation, the

concept of effective cloud parameters may fail (Kylling et

al., 2005), but enough data are usually not available to

verify that.

Three-dimensional radiative transfer effects in the

ultraviolet spectral range caused by clouds have been

recently studied both experimentally and theoretically

(Brasseur et al., 2002; Kylling et al., 2005; Marshak and

Davis, 2005). Topography also plays a role. For example,

mountains may either block part of the diffuse-sky radiance, or cause an increase by reflection of extra radiation

at snow-covered mountain slopes (see Sections 7.2.6 and

7.2.7). Moreover, air or clouds below an observation site

enhance the backscattered radiation and lead to further

increase in irradiance (McKenzie et al., 2001a). Considerable research efforts are spent on three-dimensional

radiative transfer applications, to gain a detailed understanding of the processes that control the interaction of

radiation with the atmosphere and the surface. The main

objective of the Radiation Transfer Model Intercomparison

(RAMI) was to improve understanding of the important

interactions between radiation and vegetation (Pinty et al.,

2004).

Validation campaigns have mostly concentrated on

cloudless conditions. Comparisons of spectral irradiance

(e.g., Mayer et al., 1997; van Weele et al., 2000) and

actinic flux (Bais et al., 2003; Früh et al., 2003; Shetter et

al., 2003; Hofzumahaus et al., 2004) for a cloudless sky

SURFACE ULTRAVIOLET RADIATION

7.19
generally found differences of about 5-10% between

observations and simulations, which were usually within

the combined uncertainty of measurement and model.

Comparisons for cloudy sky are rare, due to the difficulty

in obtaining cloud parameters independent of the irradiance observations. In the European project INSPECTRO

(Influence of Clouds on the Spectral Actinic Flux in the

Lower Troposphere), simultaneous observations of radiation and clouds using ground-based, aircraft, and satellite

instrumentation provided a first dataset to evaluate threedimensional radiative transfer calculations. Kylling et al.

(2005) showed for a broken cloud case that the onedimensional approach was not able to simultaneously

reproduce the measured enhancement of the actinic flux

above the cloud and the reduction below the cloud.

Further evaluation of the campaign data is in progress.

In summary, radiative transfer in the ultraviolet

spectral range has reached a mature state. If the input

parameters are known, the radiation can be simulated with

reasonable accuracy for many applications. However,

some of these input parameters (e.g., a reasonable description of clouds or the single scattering albedo of the aerosol)

are not easily attained. While satellite observations of

clouds allow derivation of average radiation levels, they

usually fail to provide information about the small-scale

variability in space and time, which could be essential for

biological systems. While the main aim of research in the

ultraviolet spectral region remains the study of irradiance

and actinic flux for applications in biology and chemistry,

observations of radiance give much deeper insights into

the interactions between radiation and the atmosphere or

surface (Huber et al., 2004; Kuchinke et al., 2004). It may

also help to validate and improve radiation codes and to

decide which input parameters are required for specific

applications.

7.3.4 Short-Term UV Forecasting

In the four years since the previous Assessment

(WMO, 2003), the UV Index (WMO, 1994, 1998) has

become the internationally accepted standard for disseminating information on solar UV levels to the general

public. Forecasts, as well as measurements of the UV

Index, now appear on numerous web-sites in many countries (Appendix 7B.5) and are used by TV, radio, and print

media across the globe in an effort to change population

behavior and moderate exposure to solar UV radiation.

UV Index forecasts are a direct application of the scientific work on understanding the factors such as absorption

and scattering affecting the transmission of UV radiation

through the atmosphere to the Earth’s surface.

Forecasts of the UV Index have been provided by

numerous agencies since the 1990s using different computational methods (Austin et al., 1994; Burrows et al.,

1994; Long et al., 1995; Burrows, 1997; Rikus, 1997; Bais

et al., 1998; Staiger et al., 1998; Lemus-Deschamps et al.,

1999; Renwick and McKenzie, 2001; Lemus-Deschamps

et al., 2004; Staiger and Koepke, 2005). Forecasting the

column ozone amount and estimating the atmospheric

aerosol contribution to absorption remain the major difficulties in improving the accuracy of calculating the clearsky surface UV irradiance and in subsequently forecasting

UV indices. Computed UV irradiance for clear-sky conditions is generally overestimated in comparison to measurements (Kalliskota et al., 2000; McKenzie et al., 2001b),

mainly due to lack of information on AOD (De Backer et

al., 2001). Long-term correlations of aerosols with measured UV irradiance may improve AOD forecasts locally

(Krzy ́scin et al., 2001; Petters et al., 2003; Gröbner and

Meleti, 2004; Goering et al., 2005).

Including clouds in UV Index forecasting requires

both accurate forecasts of cloud cover and cloud types, as

well as more accurate data on cloud UV transmittance.

Forecasts of UV Index have used empirical relations

between forecast cloud types and measured UV radiation

to adjust clear-sky forecasts. More recent approaches use

information on cloud and surface albedo provided by

numerical weather prediction (NWP) models (Staiger and

Koepke, 2005). Work on improving agreement between

calculations utilizing clouds and their effects on UV transmittance and surface measurements continues (Arola et

al., 2003b; Krzy ́scin et al., 2003; Lane-Veron and

Somerville, 2004; Su et al., 2005).

The major parameter affecting the forecast of the

clear-sky UV Index is the estimated ozone amount, which

can be determined either using the persistence of satelliteobserved ozone or from a range of NWP meteorological

variables via statistical regressions (e.g., Plets and

Vynckier, 2000). Some forecasting centers such as the

National Centers for Environmental Prediction (NCEP),

the Koninklijk Nederlands Meteorologisch Institute

(KNMI), the European Centre for Medium-Range Weather

Forecasts (ECMWF), and the Australian Bureau of

Meteorology (BOM) assimilate ozone into their forecast

models using satellite data to produce forecasts of the

global (or continental) ozone fields a number of days in

advance (Lemus-Deschamps et al., 2004). Research validating the results of the BOM numerical weather prediction (NWP) system’s forecasts of UV Index continues with

comparisons against surface observations (Gies et al.,

2004; Lemus-Deschamps et al., 2004) in an effort to

improve agreement. This ongoing validation study is coupled with research to improve the presentation of the UV

SURFACE ULTRAVIOLET RADIATION

7.20
Index in communications to the public, with the aim of

increasing its impact and making it more effective in

changing behavior (Dixon and Armstrong, 1999; Dixon et

al., 2002; Blunden et al., 2004). A range of different UV

Index presentations utilizing the recommendations of the

World Health Organization (WHO, 1994; Rehfuess, 2002)

is being tested in Australia by the Cancer Councils, with

planned follow-up assessments of their impact.

7.4 UV CLIMATOLOGY

Long- and short-term UV variability has been

studied extensively during the last few years by analyzing

measurements of surface UV irradiance that are now available for more than 10 years at several locations worldwide. Routine spectral measurements started only in the

late 1980s and thus long-term changes that are calculated

from these datasets should be smaller compared with what

they would have been if UV records were available before

the beginning of the ozone depletion in the early 1980s.

Satellite estimates are also available for long periods (since

1979) and provide global coverage, which allows the derivation of UV estimates for any location on the Earth.

However, gaps in the satellite records in the mid-1990s

increase the uncertainty and complicate the calculation of

long-term changes. In addition to the relative shortness of

the UV records, surface UV exhibits a large variability

from natural factors, such as clouds, ozone, aerosols, surface albedo, and UV-absorbing atmospheric constituents

(e.g., SO2 and NO2), which reduce the significance of the

estimated long-term changes. Extreme values of UV radiation at the surface as a consequence of these factors have

been reported and were discussed in Section 7.2. The attribution of the causes for the observed changes is among

the present research priorities, especially as it pertains to

the response of surface UV to the recovery of the ozone

layer.

7.4.1 Derived from Ground-Based UV

Measurements

7.4.1.1 LONG-TERM CHANGES

The magnitude and statistical significance of the

long-term changes in surface UV irradiance derived from

analysis of ground-based measurements are of the order

of a few percent per decade, depending on location, season,

wavelength range, and the period of measurements. The

sign of these changes varies also with location and depends

strongly on the length of the measurements series. As

measurements suitable for calculation of long-term

changes are sparse, there are difficulties in projecting the

derived results to other time periods or other locations.

An example is shown in Figure 7-4, which presents time

series of monthly mean anomalies in noon erythemal irradiance under all-sky conditions at 11 UV monitoring sites

distributed worldwide. Linear regressions over the period

from the early 1990s to the mid-2000s indicate negative

changes at stations in the Southern Hemisphere and in the

Arctic. This pattern is especially evident if linear regressions are calculated since 1998, when the first signs of a

slowdown in the ozone depletion were observed (see

Chapter 3, Section 3.2), suggesting that surface UV radiation has responded to ozone increases, at least at these

sites. In the northern midlatitudes, the observed UV irradiance changes are positive. The significance of the calculated changes is affected by the uncertainties of the

measurements, which may range between 5% and 10%,

depending on the site and the period. Data in the beginning of the records have usually higher uncertainties (see

Section 7.3.1.5). Statistically significant linear changes

(neglecting the measurement uncertainties) for the entire

record of measurements were found only at Lauder and

Palmer (−8.4% per decade at the 99% confidence level)

and at Barrow (−4.4% per decade at 95% confidence

level). Observations in Lauder prior to 1994 were performed only under fair-weather conditions, and this may

have influenced the calculated linear changes. Statistically

significant positive changes are found at all seasons in the

station of Lindenberg, produced by two abrupt changes in

UV irradiance that occurred in 1996 and 2003, respectively. The change in 2003 is shown also at Bilthoven.

For the spring months the linear changes are insignificant

and negative at all stations, except at Thessaloniki,

Lindenberg, and Bilthoven. In the summer months, half

of the stations show negative and the other half positive

changes that are all insignificant, except from Lauder

(–8.9% per decade at the 95% confidence level). The UV

irradiance changes at these stations are dominated mainly

by the variability in cloudiness and, to a lesser extent, by

ozone and aerosols. The variability of clouds and aerosols

has masked the detection of increases in surface UV irradiance that were expected to result from the observed

decreases in total ozone in the early 1990s that were caused

by stratospheric aerosols from the Mt. Pinatubo eruption.

At most stations, the use of measurements made only under

clear skies results in too little data to allow reliable statistical analysis.

Glandorf et al. (2005) analyzed in more detail 12

years of UV measurements at Thessaloniki, Greece, and

and 14 years of UV measurements at Sodankylä, Finland,

and concluded that no unambiguous upward trend could

be constituted at either station, due to the high natural variability of the data. It was suggested that for the given variSURFACE ULTRAVIOLET RADIATION

7.21
7.22

Figure 7-4. Long-term changes in monthly mean erythemal irradiance recorded by UV spectroradiometers within ±1 hour around local noon at 11

sites distributed worldwide, shown as departures from the long-term averages. Linear regressions on the data (red lines) were used to estimate the

corresponding linear changes after 1998, when the first indications of a slowdown in total ozone decreases have been observed. The trends are

listed in the upper right of each plot, with color-coding to indicate the statistical significance (red = 99%; blue = 95%; black = negligible). This figure

was prepared using updated series of published UV datasets. The statistical significance of the changes was calculated from the data variability

only, neglecting the uncertainties in the measurements.
ability in total ozone and clouds at the two sites, more than

15 and 12 years of data, respectively, would have been

needed for the detection of persistent statistically significant trends in surface UV radiation, in accordance with

previous findings by Weatherhead et al. (1998). However,

due to the complexity and the self-regulating mechanisms

of our climate system, quasi-periodic changes in UV radiation are more likely to occur than long-lasting linear

changes. Using two different periods of the Sodankylä

dataset (1993-2001 vs. 1990-1997), Lakkala et al. (2003)

calculated substantially different linear changes in spectral UV irradiance at 305 nm, even of opposite sign.

At northern midlatitude sites, surface UV irradiance was found to be increasing since the beginning of

1990s, while elsewhere there are signs of leveling. The

updated record of clear-sky spectral UV irradiance at 305

nm and 325 nm and at 63° solar zenith angle over

Thessaloniki (Kerr and Seckmeyer et al., 2003) indicates

that UV has continuously increased since the beginning of

the record (1990) at both wavelengths (Figure 7-5).

Although not very clear, there are indications that after

about 2000, surface irradiance has increased with smaller

rates or has even been constant, following the stabilization or a slight increase of total ozone over Thessaloniki.

The upward trend during the 1990s at wavelengths unaffected by ozone absorption could only be explained by an

increase in atmospheric transmission. Aerosol optical

depth and SO2 column over Thessaloniki have decreased

since 1997 (Zerefos, 2002; Kazadzis et al., 2005), while a

widespread brightening of the atmosphere has been

observed since the late 1980s (Wild et al., 2005).

Increasing tendencies of surface UV, unexplained by the

total ozone decline alone, were reported also for

Hohenpeissenberg, Germany (Trepte and Winkler, 2004).

Reduced sunshine duration and increased cloudiness lead

to a UV decrease that partially compensates for the UV

increase caused by ozone loss in spring; in the autumn,

these effects overcompensate for the effect of ozone loss.

The increasing peak values of UV irradiances that

were reported in the previous Assessment (Kerr and

SURFACE ULTRAVIOLET RADIATION

7.23

Figure 7-5. Long-term variability in

monthly mean solar spectral irradiances at 324 nm (upper panel) and at

305 nm (middle panel) measured at

Thessaloniki, Greece, under clear

skies at 63° solar zenith angle, shown

as departures from the long-term

(1990-2006) averages. The lower

panel shows the corresponding

departures in the ozone column.

Linear regressions on the data and the

derived trends per decade are also

shown for each parameter. Update

from Zerefos et al., 2001.
Seckmeyer et al., 2003) in response to decreasing total

ozone at New Zealand show a clear decreasing tendency

in the recent four years (Figure 7-6). The magnitude of

this decrease is larger than that expected by the increase in

total ozone over the area, indicating that aerosol extinction may have also increased during this period. Aerosol

data from GOME show increasing aerosol absorption

between 45°S and 60°S over the period 1995 to 2000 (de

Graaf et al., 2005).

Measurements of surface UV irradiance with broadband radiometers are available for longer periods compared with spectral measurements, although they usually

have higher uncertainty (Seckmeyer et al., 2005). Ten

years (1990-2000) of solar UV-B irradiance (290 to 320

nm) measurements at Hiratsuka, Japan (35.35°N,

139.27°E), show an increase at a rate of 15.7% per decade

(Sasaki et al., 2002). After removing the effect of the

Quasi-Biennial Oscillation (QBO) with a 26-month

moving average filter, and removing the effect of cloudiness by normalizing with shortwave irradiance measurements, the rate of increase was reduced to 12.2% per

decade, corresponding to the observed total ozone changes

at the same location. Statistically significant changes of

monthly averages are found only for winter (16.7% per

decade). Long-term (1968-2003) UV irradiance (300-380

nm) measurements at the Meteorological Observatory of

Moscow State University (Figure 7-7) show a distinct

interannual variability, with low levels in the 1970s and

1980s and an overall increase from late 1970s toward the

late 1990s (Chubarova et al., 2005). This increase is

believed to be due to the decrease in cloud amount and

aerosol optical depth (since ~1994), with contributions of

up to 10% and 2%, respectively. A decline in atmospheric

turbidity of 8-10% during the period 1994-2002 as comSURFACE ULTRAVIOLET RADIATION

7.24

01 80

01 8201 8401 8601 8801 9001 9201 9401 9601 9801 0001 02

01 0401 06

Summer Year (December - February) Deviation % Erythemal Irradiance Anomaly (A %) Figure 7-7. (a) Interannual variability of simulated

erythemal irradiance and measured irradiance integral from 300 nm to 380 nm (QUV380) for the period

May-September in Moscow, Russia, shown as deviations from the 1968-1997 mean. (b) Anomalies in erythemal irradiance (A) due to different atmospheric

factors: effective cloud amount A(N_eff), total ozone

A(X), aerosol optical thickness A(τaerosol), and cloud

optical thickness A(τcloud). Adapted from Chubarova

and Nezval, 2000; Chubarova et al., 2005.

Figure 7-6. Long-term changes in summertime

ozone (upper part) and in peak UV Index (lower part)

deduced from assimilated ozone data (lines) and from

UV spectral measurements at Lauder, New Zealand

(symbols). Update from McKenzie et al., 2004.
pared with the 1976-1995 period is observed at many other

sites in Russia (Makhotkina et al., 2005). Increases in erythemal irradiance at Norrköping, Sweden, of ~5% per

decade were reported by Josefsson (2006) for the period

1983-2003.

Increases in surface UV irradiances during the last

decade are evident from several datasets. These changes

are attributable to the increasing of atmospheric transmission because of ozone depletion, global reduction of cloud

optical thickness, and reduction of the amount or nature of

aerosols. Long-term satellite observations (World Climate

Research Programme Global Energy and Water Cycle

Experiment (GEWEX) International Satellite Cloud

Climatology Project, ISCCP) reveal an overall increase of

solar irradiance at the surface from 1983 to 2001 (Figure

7-8) at a rate of about 1% per decade, despite the finding

that up to about 1990, solar irradiance measured at the surface was decreasing (Pinker et al., 2005). With respect to

clouds, Meerkötter et al. (2004) found no significant longterm trend in cloud amount between 1990 and 2003 from

the analysis of NOAA/AVHRR data and from Surface

Synoptic Observations (SYNOP), although there are indications of decreasing cloud cover after 2000. Decreases

in cloud cover of less than 1%/decade, increases in clearsky days of ~0.6%/decade, and decreases in overcast days

of ~0.8%/decade have been observed over most of China

during the last half of the 20th century (Qian et al., 2006).

For Barrow, Alaska, a statistically significant upward trend

in cloud occurrence frequency, from 76 to 82% between

1976 and 2001, was reported (Dutton et al., 2004).

Evidently, only in some areas may an increase in atmospheric transmission be attributed to clouds.

An indication of cleaning of the atmosphere is supported by the long-term aerosol optical depth data at Uccle,

Belgium (Cheymol and De Backer, 2003), which were

retrieved from Brewer direct-Sun measurements. They

reported a statistically significant negative trend of 24.6 ±

3.7%/decade over the period 1989-2002 for the aerosol

optical depth at 306.3 nm; the annual mean was ~0.8 in

1992 and decreased to ~0.6 in 2002. Significant seasonal

trends at all wavelengths (in the range 306.3-320.1 nm)

were found only in winter months, while in the summer

period only at 306.3 nm, a decrease in aerosol optical depth

of 12.4 ± 5.5%/decade is significant at the 2σ level.

Similarly, Jaroslawski and Krzy ́scin (2005) reported that

aerosol optical depth in Belsk, Poland, shows a decreasing

tendency in the beginning of 1990s followed by an

increase in surface UV irradiance (UV Index) of about 4%.

7.4.1.2 SHORT-TERM AND SPATIAL UV VARIABILITY

Apart from the long-term changes, the variability

of surface UV irradiance has been extensively studied

using datasets measured all over the globe and at different

periods. Some datasets are sufficiently long to be used for

the establishment of a UV climatology that is representative for the region around the measurements site(s)

(Tarasick et al., 2003; den Outer et al., 2005; Liao and

Frederick, 2005). Most of the observed variability is

caused by the influence of well-known factors, such as

ozone, clouds, aerosols, surface albedo, SO2, and NO2,

which were discussed in Section 7.2.

A climatology of surface erythemal irradiance,

expressed as UV Index, was developed by Tarasick et al.

(2003) using spectral measurements from 14 UV monitoring stations, spanning from 82.5°N to 69.0°S, but

mainly distributed in Canada. The most prominent feature at all stations is the high degree of variability that is

primarily due to clouds and thus is more pronounced at

cloudy sites, such as Halifax and Goose Bay, and much

less at Mauna Loa. Smaller, positive or negative, variations are due to ozone variability. Also apparent, at most

of the Canadian sites as well as San Diego, is that maximum clear-sky UV Index values in the 1990s were higher

than those calculated for the pre-1980 period, by about

10%, in accordance with the total ozone decline of about

6-7% during this period. Dramatically higher UV Index

values are seen at Syowa during the Antarctic spring, due

to the appearance of the ozone hole, with peak values similar to those in June at San Diego.

SURFACE ULTRAVIOLET RADIATION

7.25

Figure 7-8. Satellite-derived surface solar irradiance

(from 1983 to 2001) averaged over the globe, after

removal of the mean annual cycle. The linear regression (solid line) on the data reveals a positive trend of

~0.1% per decade. The second-order polynomial

(dashed line) indicates a small decrease from 1983 to

1992, with a reversal around 1992. Both fits are significant at the 99% level of confidence. Reprinted with

permission from Pinker, R.T., B. Zhang, and E.G.

Dutton, Do satellites detect trends in surface solar radiation? Science, 308 (5723), 850-854, 2005. Copyright

2005, AAAS.
Liao and Frederick (2005) presented a climatology

of erythemal and DNA-damaging solar irradiance at high

latitudes in the Southern Hemisphere, derived from data

recorded between 1990 and 2001 by four spectroradiometers located between 55°S and 90°S. Monthly integrals

show a distortion of the normal annual cycle in irradiance,

with enhancements occurring in October and November

due to the springtime ozone loss. In some cases, these irradiances exceed those near summer solstice in December.

At Palmer Station, near 65°S, the monthly erythemal irradiance in November 1997 was more than double that

observed five years earlier. The combined influence of

the ozone column and cloudiness variability results in a

large interannual variability of surface erythemal irradiance. Monthly averages vary between about ±20 and

±50% for the period from October to December.

Geographical differences in surface UV irradiance

were investigated using measurements from multifilter

radiometers (Diaz et al., 2006) and from erythemal

radiometers (Pazmiño et al., 2005) at stations located at

different latitudes from tropical to sub-Antarctic regions.

Extremely high UV-B irradiances were recorded when part

of the polar vortex with low ozone moved over the measurement sites. McKenzie et al. (2006) have found that

peak erythemal UV irradiances in New Zealand are ~40%

higher than at locations of similar latitude in North

America, as a result of the lower aerosol and ozone

amounts in New Zealand and the smaller Earth-Sun separation during the Southern Hemisphere summer.

7.4.2 Estimates from Satellite Data

Global UV climatology and trends based on TOMS

data for the period 1979-1991 were reported in the WMO

(1999) Assessment, and since then no further studies on

UV changes have been conducted using the extended

dataset, although inferences can be made on the basis of

the leveling off in ozone depletion reported by Reinsel et

al. (2005). One of the problems for extending the satellite

UV data in the recent years is the gap between Nimbus-7

and EP TOMS. Final EP/TOMS calibrations are not available after 2001, which currently prevents updating global

TOMS UV climatology and trends. Updated global satellite UV trends will be a subject of future work, most likely

combining TOMS, SeaWifs (Herman, 2004), SBUV/2

(Chapter 3, Appendix A1.1), and the new Ozone

Monitoring Instrument (OMI) on NASA Aura spacecraft

data (Tanskanen et al., 2006).

However, regional UV climatologies have been

developed based on RT modeling and various input parameters. Daily maps of surface UV doses over Europe since

1984 were developed by Verdebout (2004a; 2004b) based

on total ozone from TOMS, Total Ozone Vertical Sounder

(TOVS) and GOME, and taking into account cloud effects,

using METEOSAT/MVIRI images, and effects from tropospheric aerosols, snow cover, and altitude. Fioletov et

al. (2004) produced a regional UV climatology by calculating the UV Index over Canada and the United States for

the period 1980-1990 using total ozone and reflectivity

from TOMS and UV Index derived from ground-based

global solar radiation, total ozone, snow, and dew point

temperature measurements.

Similar climatologies were derived for Argentina

(Luccini et al., 2006), revealing high values of UV irradiances at the northwestern tropical high-altitude Andean

plateau. For December-January, the monthly mean UV

Index was higher than 18 and the daily erythemal dose was

higher than 10 kJ m-2. For most of the continental region,

clouds attenuate the clear-sky erythemal irradiance by less

than 20% on the average throughout the year.

7.4.3 Reconstruction of Past UV Records

Assessment of UV radiation changes in the past is

important because many biological and health-related

effects depend on UV dose accumulation during long

periods. Because reliable spectral UV measurements are

only available since the end of the 1980s, several attempts

have been made and various methods have been proposed

to reconstruct past UV irradiance records.

Erythemal irradiances or irradiance integrals in

other UV spectral ranges are calculated from statistical

relationships between actual measurements and various

geophysical parameters (e.g., ozone, total solar irradiance,

snow cover, etc.) recorded in recent years (e.g., McArthur

et al., 1999; Gantner et al., 2000; Fioletov et al., 2001;

Diaz et al., 2003; Trepte and Winkler, 2004), or using

neural-network techniques (Janouch, 2004). Another

hybrid approach uses a radiative transfer model to estimate UV levels under cloud-free conditions, followed by

correction for cloud effects based on either total solar irradiance (Kaurola et al., 2000; Bodeker et al., 2002; den

Outer et al., 2005), or sunshine duration measurements

(Lindfors et al., 2003; Lindfors and Vuilleumier, 2005).

Reuder and Koepke (2005) used RT model calculations

and algorithms for deriving the required input parameters,

total ozone, surface albedo, and clouds from routine observations and aerosol properties from climatological information and the Optical Properties of Aerosols and Clouds

(OPAC) database. Chubarova and Nezval (2000) reconstructed UV series by calculating UV anomalies due to

different geophysical factors. Coupled ChemistryClimate Models (CCMs) have also been used for deriving

past UV series (Kaurola et al., 2004; Reuder et al., 2005).

SURFACE ULTRAVIOLET RADIATION

7.26
In particular, UV trends since 1980 were obtained for the

entire globe from the Unified Model with Eulerian

Transport and Chemistry (UMETRAC) climate model,

including coupled chemistry, and from the FinROSE

chemistry-transport model; they demonstrated good agreement with TOMS UV estimates (Kaurola et al., 2004).

Reconstructed UV series allow the assessment of

UV trends, as well the estimation of climatic UV irradiances over large areas and long periods. For example, the

UV Index climatology over Canada and the United States

has been derived from reconstructed UV data combined

with TOMS UV estimates (Fioletov et al., 2004). The

longest time series to date of UV irradiance reconstructed

from total ozone and sunshine duration datasets has been

obtained for Switzerland (Lindfors and Vuilleumier,

2005). In the period 1926-2003, distinct interannual

changes in erythemal irradiance have been demonstrated,

with high values in the middle 1940s, in the early 1960s,

and in the 1990s (Figure 7-9). The irradiance changes

prior to 1980 are mainly explained by variations in relative sunshine duration, caused mainly by changes in

cloudiness, while the increase in the 1990s is due to the

negative trend in total ozone. Reconstructed daily UV

doses between 1968 and 2001 at Hohenpeissenberg,

Germany, show positive changes of 2-5% in the summer

months (between March and August) and negative changes

of similar magnitude in winter (September to October)

(Trepte and Winkler, 2004). The former are due to

decreases in cloudiness and in ozone column, whereas the

latter are attributable mainly to increases in cloud optical

thickness.

Annual integrals of erythemally weighted UV doses

were derived for the period 1979-2003 at Bilthoven, The

Netherlands, based on model calculations, total ozone, and

pyranometer data. These data show increasing UV to the

mid-1990s followed by a small decrease, except for a relatively high value in 2003 due to smaller reductions by

clouds. Linear regressions over the entire 25-year period

revealed increases in UV dose by 3.1% per decade under

clear skies and 5.5% per decade under all-sky conditions

(den Outer et al., 2005; Lindfors and Vuilleumier, 2005).

For the period 1936-2003 in Northern Norway, Engelsen

et al. (2004) found trends in monthly reconstructed UV-B

doses of 4.5%, 2.8%, and 1.3% per decade, respectively,

for March, April, and May. UV-A and PAR doses, which

are unaffected by total ozone, reveal smaller trends (1-2%

per decade) in March and April, while in May the trend is

comparable with that in the UV-B.

The reconstructed UV time series at European

observatories in Central and Eastern Europe since the mid1960s (Krzy ́scin et al., 2004; Chubarova et al., 2005)

demonstrate a pronounced decrease in UV irradiance in

the late 1970s and beginning of the 1980s and confirm the

increase in the 1990s that is reported from measurement

records (see Section 7.4.1). The decreases are mainly

explained by negative UV anomalies, up to −10% due to

increases in cloud amount, and up to −5% due to total

ozone increases (Figure 7-7). It should be noted, however,

that sometimes variations of clouds and total ozone are

not completely decoupled, both being affected by circulation patterns. In such cases it is difficult to separate their

influence on surface UV radiation.

Changes in surface UV radiation during the last century derived from reconstruction methods appear comparable in magnitude to those observed in the 1990s and

2000s. Although short-term UV irradiance variations

imposed by different geophysical factors vary between

regions, the similarity in the main features of reconstructed

erythemal irradiance series confirms the prevalent effect

of global processes, such as for example, the changes in

atmospheric processes over Europe in the mid-1970s

(IPCC, 2001) and the ozone decrease in the midlatitudes

of the Northern Hemisphere in the 1990s. Surface UV

radiation exhibits high values in the last 30 years, low

values in the 1950s and 1960s, and high values again in

the 1940s and early 1950s. Most of the past changes have

been caused by changes in cloud amount and, to a certain

extent, in aerosols; whereas ozone changes have been

SURFACE ULTRAVIOLET RADIATION

7.27

1930 1940 1950 1960 1970 1980 1990 2000 -10

-5

10

deviation (%) deviation (%)

-10

-5

10

15

DEVIATION FROM 1940 - 1969 LEVEL

deviation (%) deviation (%)

UVozon - 5-year running mean

UVsunshine - 5-year running mean 5-year running mean

UVsno

 - 5-year running mean

UVest - yearly

UVest - 5-year running mean

UVest - linear trend 1979-99

Figure 7-9. Upper panel: Time series of reconstructed erythemal irradiance over Switzerland

shown as deviations (changes in %) from the 1940-

1969 mean. Lower panel: Percentage contribution

of changes in ozone, sunshine duration, and snow

depth to the UV variability. Adapted from Lindfors

and Vuilleumier, 2005.
responsible for the increases of surface UV only in the

1980s and 1990s.

7.4.4 Polar UV Variability

Studies of polar UV variability have recently benefited from in situ measurement networks becoming more

robust (Diaz et al., 2003; Bernhard et al., 2004; Lakkala et

al., 2005; Pazmiño et al., 2005; Bernhard et al., 2006) and

now provide regionally specific information about the

effects of clouds and surface albedo. On the high Antarctic

Plateau (Bernhard et al., 2004), most cloud cover is optically thin, and 71% of clouds sampled by the NSF UV

Monitor at the South Pole had optical depth less than 1.

Due to small optical depth and high surface albedo at the

South Pole, the average cloud attenuation of UV radiation

was only 5%, and the maximum observed UV cloud attenuation was 23%. This is consistent with radiative transfer

studies and measurements by Nichol et al. (2003).

Measurements reported by Luccini et al. (2003) for

Almirante Brown Station (64.9°S, 62.9°W) show that the

average cloud attenuation for erythemal UV was 46%

under overcast skies without precipitation, and 73% under

clouds with precipitation. Luccini et al. (2003) and Lubin

et al. (2002), both studying summertime irradiance levels

on the Antarctic Peninsula, show high (short-term) temporal variability for cloud UV attenuation in their measurements.

Liao and Frederick (2005) analyzed the interannual

variability in springtime UV-A and UV-B irradiance at all

four of the NSF UV monitoring program’s high southern

latitude sites over 1990-2001, and determined that Palmer

Station exhibits the greatest interannual variability in both

UV-A and erythemal irradiances due to cloud and albedo

effects. They also examined trends in monthly integrated

erythemal irradiance, and found upward trends during

November and December at McMurdo and South Pole,

but no upward trend during October. This result may

reflect the springtime ozone decrease already reaching its

maximum severity during October over the continent by

1990, but gradually increasing in duration since 1990.

Diaz et al. (2003) used the NSF UV monitoring program

data to develop an empirical model that allows estimation

of historical trends in UV-B if total column ozone and

ground-based pyranometer data are available. Their

results show increases in UV-B at the South Pole during

October since 1983, consistent with the beginning of the

anthropogenic springtime ozone depletion.

Lubin et al. (2004) used satellite data to study the

multidecadal variability and trends in Antarctic UV-B, in

the context of exposure by Southern Ocean phytoplankton.

These authors also confirm that elevated UV-B irradiances,

related to ozone depletion, had already occurred prior to

the ozone “hole” discovery (before 1985), and that these

enhanced irradiances had the potential to impact the most

sensitive organisms years before dedicated biological field

work began and before such damage could be detected.

These authors also show that the most frequent occurrences of enhanced UV-B over waters containing significant phytoplankton biomass occurred in the Weddell Sea

and Indian Ocean sectors of Antarctic waters, where there

has historically been very little related biological field

work. These two factors may help explain why reports of

ecological impact due to the springtime ozone decrease

have been modest compared with initial speculations.

When the polar vortex breaks up, the dilution of

ozone-poor air from within the vortex can contribute to

ozone decrease over midlatitudes, leading to enhancement

of surface UV. Konopka et al. (2003) studied the fate of

vortex remnants and the chemistry occurring within them

during the spring of 1997 and 2000, using a Lagrangian

chemical transport model, and found different behavior in

the lower stratosphere and in the midstratosphere. Above

20 km, remnants from the polar vortex (Orsolini, 2001)

remain long-lived in the summer westward circulation.

Balloonborne in-situ measurements of water vapor and

methane (Durry and Hauchecorne, 2005) provided evidence for such vortex remnants in the midlatitude summer

stratosphere. Below 20 km, the subtropical jet bounds the

meridional propagation of those remnants (Piani et al.,

2002), and their lifetimes are considerably reduced due to

enhanced stirring by synoptic eddies. Ajtic et al. (2004)  ́

noted that despite the regional differences observed in the

springtime impact of Antarctic ozone depletion at midlatitudes, when the Antarctic polar vortex breaks down in late

December and January, the resulting ozone reduction at

southern midlatitudes is essentially zonally constant and

decreases toward lower latitudes.

The expansion of the Antarctic polar vortex during

the 1990s, both in spatial extent and temporal extent

through early summer, has increased the frequency of elevated UV-B episodes over sub-Antarctic populated areas

(Casiccia et al., 2003; Pazmiño et al., 2005). These

episodes are no longer just small pockets of ozonedepleted stratospheric air coming from the breakup of the

polar vortex, but now include occasional excursions of the

polar vortex edge over Ushuaia and Punta Arenas. This

occurred 44 times in the years 1997, 1998, and 2000 combined, and some episodes lasted 3-4 days. Surface measurements show average erythemal UV increases of 68%

over Ushuaia since 1997, and episodic total UV-B

increases of up to 80% over Punta Arenas.

Diaz et al. (2003) show that Barrow, Alaska, has

experienced UV-B increases related to springtime ozone

SURFACE ULTRAVIOLET RADIATION

7.28
depletion in March and April, but these increases are a

factor of ten smaller than those observed at the South Pole.

Summertime low-ozone episodes in the Arctic also affect

surface UV-B irradiances (Orsolini et al., 2003). These

summertime events result from gas-phase chemistry

involving nitrogen and hydrogen cycles, which become

very efficient during the 24-hour insolation that occurs in

the Arctic summer. During summer 2000, two low-ozone

episodes brought about erythemal UV increases of order

10-15%, each lasting more than five days.

7.5 EXPECTATIONS FOR THE FUTURE

7.5.1 Links with the Recovery of the Ozone

Layer

All other UV modulating geophysical variables

being stable, it is expected that decreases in stratospheric

ozone will lead to increases in UV radiation at the surface.

However, in practice it is difficult to detect such changes

because of the high variability in UV radiation that is

caused by factors other than ozone. During the period of

most-rapid ozone decline (from ~1980 to ~1997), the situation was exacerbated by a lack of high-quality calibrated

UV data. In recent years, data quality and quantity have

improved. However, there are still only a few sites globally with sufficient long-term data to detect the effects of

the observed changes in ozone. Spectroradiometric data

from unpolluted NDACC sites located in regions of ozone

depletion (e.g., Lauder, New Zealand, and Antarctica)

offer the greatest hope for detecting these changes over

the natural variability caused by clouds and aerosols.

However, to date, no studies have clearly demonstrated

any reductions in surface UV irradiance that may be attributable to possible increases in ozone since the late 1990s.

The most convincing studies to demonstrate these

long-term changes all used some sort of filtering to minimize cloud effects (Calbó et al., 2005). Other confounding

factors include systematic changes in surface albedo and

air quality (aerosols and tropospheric ozone). Even with

modern state-of-the-art equipment at our disposal, it

remains a challenge to measure UV radiation to sufficient

accuracy to unequivocally detect future changes in UV

radiation that may result from future changes in ozone.

At Lauder, New Zealand, the observed increases in

peak summer UV irradiance due to ozone depletion have

ceased (McKenzie et al., 1999). Since 1999, the summer

ozone levels have leveled off or started to increase, and in

the most recent published update, it has been noted that

the observed peak UV Index values now appear to be

diminishing (McKenzie et al., 2004). Others factors may

also be important, because the UV reductions in recent

years are larger than those expected from changes in ozone

alone, and there are apparent reductions in UV-A that

cannot be attributed to ozone changes. These reductions

may therefore be due in part to increases in absorbing

aerosols over this region, which have been inferred from

satellite data (de Graaf et al., 2005). Satellite-derived

ozone measurements show that the leveling off (or even

turnaround) in summertime ozone occurs throughout the

midlatitude region of the Southern Hemisphere (McKenzie

et al., 2004; Reinsel et al., 2005).

7.5.2 Predictions of UV Long-Term

Changes

Several attempts have been made to predict future

changes of ultraviolet radiation. Such predictions are often

based on radiative transfer simulations considering predicted changes of the relevant input parameters. All predictions made to date have concentrated on the effect of

ozone changes, which is considered as the most prominent

of these parameters, especially for the lowermost UV-B

wavelengths and in the polar regions. One of the first estimates of how ultraviolet radiation might evolve in the 21st

century was provided by Slaper et al. (1996), who based

their calculations on ozone predictions for different scenarios under cloud-free conditions. Later studies included

clouds, which were assumed to be constant in time, in the

model calculations (Taalas et al., 2000; Reuder et al.,

2001). While calculations including clouds are more realistic, any changes of UV radiation were still due to ozone,

which is neither the only, nor the main, factor influencing

ultraviolet radiation. Clouds, aerosols, and surface albedo

are also very important (see Section 7.2). These are subject to change in a future climate, and they have already

changed during the last decades. This is manifested by a

worldwide decrease in total solar irradiance of ~4%

between 1961 and 1990 (Liepert, 2002), which seems to

have been followed by an increase in the 1990s (Wild et

al., 2005). These changes are attributed to a modification

in cloudiness and aerosol, which also affects ultraviolet

radiation. A variation of this magnitude is comparable to

the effect of ozone and may thus strongly reduce or

enhance ozone-related trends. Based on the A2 and B2

IPCC emission scenarios, Giorgi et al. (2004) predicted

decreases in the mean cloudiness over the period 2071-

2100 in Europe, with respect to the period 1961-1990. For

the summer months, these changes range from approximately −2 to −15%. This would result in an increase in

the daily UV dose of ~10% instead of a reduction of ~5%

due to the recovery of the ozone layer (Reuder et al., 2005).

Another indirect effect on long-term trends in cloudiness

SURFACE ULTRAVIOLET RADIATION

7.29
involves ocean plankton and dimethyl sulfide (DMS) production, which feeds back to clouds via production of

cloud condensation nuclei (CCN). Cloud formation

depends on the availability of CCN, which may be produced locally through sea-to-air ventilation of DMS. PAR

is important for DMS production, whereas UV photolysis

leads to DMS loss (Deal et al., 2005). If long-term changes

of clouds and aerosols cause an increase in surface UV

irradiance, then any ozone reductions in the future will

become more important for UV radiation levels.

Depending on the application, interest might be in

either cloudless sky trends (e.g., if human exposure at the

beach is investigated) or in all-sky changes (e.g., if exposure of the biosphere is investigated). Both are affected

by changes in aerosol content, and the all-sky calculations

must also take into account changes in cloudiness. Future

scenario calculations should therefore also include the

effect of clouds and aerosols. However, available data

must be handled with care because predictions of clouds

and aerosols are highly uncertain. In particular, IPCC

(2001) states that “as has been the case since the first IPCC

Assessment Report in 1990, probably the greatest uncertainty in future projections of climate arises from clouds

and their interactions with radiation.” In fact, different

climate models disagree about the magnitude and even the

sign of the changes in cloudiness. As a consequence, the

inclusion of clouds in UV predictions would at present add

a large uncertainty range to the results. In any case, giving

a range for the potential impact of changes in cloudiness

on predictions of ultraviolet radiation is better than

excluding them completely. If clouds, as predicted by climate models, are included into the analysis, great care

must be taken in the UV calculation to treat them consistently with the cloud-radiation parameterization of the

respective CCM.

The availability of CCM simulations for ozone

column response to Montreal Protocol controls (Chapter

6) allows simulations of the corresponding multidecadal

response in surface UV irradiance, assuming that the rest

of the atmosphere and surface remain the same. By use of

ozone simulations by the Atmospheric Model with

Transport and Chemistry (AMTRAC) model (see Chapter

6 for details) as input to a UV radiative transfer algorithm

designed for satellite UV mapping (Lubin et al., 1998), it

is possible to estimate when biologically active UV irradiances might be expected to peak and to return to levels

that existed prior to anthropogenic influences on the ozone

layer. AMTRAC is currently the only Chemistry-Climate

Model with results up to 2100 (Austin and Wilson, 2006).

The ozone losses calculated by this model are larger than

observed, especially in the Southern Hemisphere (see the

blue dashed curves corresponding to the “AMTRAC”

model in Figures 6-10 and 6-12 of Chapter 6), and hence

more emphasis is given on the timing of the projected

changes of UV radiation, rather than on their absolute magnitude. Figure 7-10 shows an example of predicted erythemal irradiance at local noon for summertime at 45°N

and 45°S, where a large part of the Earth’s population lives,

and for October at 65°S. The UV calculations are based

on the average ozone column, respectively, at the latitude

bands 35°N-60°N, 35°S-60°S, and 60°S-90°S. Erythemal

irradiance increases from a reference value (calculated as

the average of the period 1970-1980) for twenty years

beginning in 1980, and this increase reaches a maximum

during ~2005. The predicted increases in erythemal irradiance are twice as large in the Southern Hemisphere, compared with the Northern Hemisphere. The return of the

calculated irradiance to the pre-1980 levels is expected by

about 2040 at the Northern Hemisphere and by about 2070

at the Southern Hemisphere, while at the southern high

latitudes this return is expected to occur later. These simulations are focused on the effect of future ozone changes

on surface UV radiation, thus it was assumed that all other

UV-influencing parameters (Section 7.2) remain

unchanged since 1980. In reality many other atmospheric

variables (e.g., cloudiness, surface albedo, aerosols, and

temperature) are expected to change during the period of

study, as a result of climate change. Changes in the vertical profiles of ozone and temperature are also expected

to influence the solar UV irradiance received at the surface. Therefore the changes shown in Figure 7-10 could

be different in magnitude and shape when predictions of

the other parameters become available and are included in

the UV calculations.

7.5.3 Factors Related to Climate Change

If the ozone layer in the next ~50 years returns to

its normal state, as predicted by models (in Chapter 6, see

Figure 6-10 and Section 6.6), then climate change will

likely dominate the surface UV variability in the future.

While most of the discussion of climate forcing generally

involves visible and infrared wavelengths, where most of

the Sun’s energy lies, there are a number of linkages

between ultraviolet radiation and climate change.

Included among the elements that force the Earth’s climate on different time scales is the total solar irradiance

incident at the top of the atmosphere, which varies by

0.08% or 1.1 W m-2 in annual mean over the 11-year solar

cycle (Section 6.11.2.1 in IPCC, 2001). The amplitude of

variability in UV-C radiation is even higher, while in the

UV-B and UV-A, only isolated spectral lines are modulated by the solar activity (see also Section 7.2.1). Such

variations in UV irradiance cause a number of climateSURFACE ULTRAVIOLET RADIATION

7.30
related changes: a change in downward shortwave flux at

the tropopause; a change in stratospheric heating (Larkin

et al. (2000) estimated the additional UV heating to be a

factor of 8-10) leading to a change in downward infrared

flux; or a change in ozone concentration (e.g., Tourpali et

al., 2003; Egorova et al., 2004; Langematz et al., 2005)

affecting also the balance of absorption of solar UV radiation in the troposphere. A change in the stratospheric

heating may also affect stratospheric zonal wind structure

and in turn affect vertically propagating planetary waves

in the winter hemisphere. Such anomalies in zonal wind

structure can propagate down into the troposphere in the

same fashion as anomalies arising from heating in the

lower stratosphere due to volcanic aerosol (e.g.,

Stenchikov et al., 2004). All of this may lead, for example,

to shifts in the positions of the sub-tropical jets in the troposphere. A change in ozone concentration feeds back on

the downward shortwave flux at the tropopause and may

thus force further changes in tropospheric dynamics (e.g.,

Shindell et al., 2001; Haigh, 2003; Labitzke and Matthes,

2003; Shindell et al., 2003) or may affect the lifetime of

reactive greenhouse gases (IPCC/TEAP, 2005).

Soon et al. (2000) suggested that the climate may

be particularly sensitive to changes in UV irradiance. They

found a stronger climate response to an imposed UV radiative forcing, preferentially absorbed in the layer above 250

hPa, than to the radiative forcing induced by increases in

CO2 or total solar irradiance. The strong climate response

to UV forcing is a coupled feedback involving vertical

static stability, tropical thick cirrus ice clouds, and stratospheric ozone.

As discussed in Sections 7.2.3, 7.2.4, 7.2.5, and

7.4.4, forcing occurs in the other direction as well.

Variations in atmospheric aerosols, air pollutants in urban

areas (e.g., NO2 and SO2), and cloud content over long

time scales, as well as variations in ozone concentration

as a result of chemical processes, lead to significant variations in UV irradiance at the surface (see also Meloni et

al. (2003c; 2004) and Hatzianastassiou et al. (2004a;

2004b) for radiative forcing by aerosols at UV wavelengths). Extreme UV irradiances that are expected in

tropical high altitude regions are often mitigated by overcast skies in summertime during the rainy season. Changes

of climate patterns in those regions linked to a decrease in

cloud density, as it happens sometimes during El NiñoSouthern Oscillation (ENSO) events, may become important for future surface UV radiation levels.

An indirect effect on the climate that works in this

direction involves ocean plankton. Increased UV penetration to the troposphere also leads to increases in the concentration of tropospheric hydroxyl radicals (OH) and

hence to decreases in the lifetime of CH4, which influences

ozone (IPCC, 2001). Such changes may play a role in

future CO2 and CH4 trends and consequently in global

warming due to greenhouse gases (e.g., Coale et al., 2004).

Moreover, changes in UV radiation are inducing changes

in biogenic emissions, some of which are important drivers

of climate change. The overall scheme of such interactions is complicated, and our understanding of the effects

on UV is still poor (EC, 2003). Although several mechanisms that couple UV radiation variations with climate

variations can be identified, it is hard to determine the

importance and magnitude of these couplings. Progress

in simulating past and future UV radiation levels may be

achieved through the combined use of CCMs, RT models,

and measurements.

SURFACE ULTRAVIOLET RADIATION

7.31

LOCAL NOON ERYTHEMAL IRRADIANCE

Figure 7-10. Estimated changes in erythemally

weighted surface UV irradiance at local noon in

response to projected changes in total column ozone

for the period 1970-2099, using zonal-averages in total

ozone in the latitude bands 35°N-60°N, 35°S-60°S,

and 60°S-90°S, and the solar zenith angle corresponding to 45°N in July, 45°S in January, and 65°S in

October, respectively. At each latitude, the irradiance

is expressed as the ratio to the 1970-1980 average.

The results have been smoothed with a 5-year running mean filter to remove some of the year-to-year

variability in the ozone predictions in the model. The

ozone losses calculated by this model are larger than

observed, especially in the Southern Hemisphere (in

Chapter 6, see the blue dashed curves corresponding

to the “AMTRAC” model in Figures 6-10 and 6-12).

Therefore, the corresponding increases in the calculated UV irradiance are overestimated. Surface UV

irradiance is predicted to peak between about 2000

and 2010 and is expected to return to the pre-1980

levels between 2040 and 2070 at midlatitudes, but later

at the southern high latitudes.
7.6 REMAINING QUESTIONS AND

UNCERTAINTIES

Despite the progress in the last four years, there are

still open issues related to the transfer of UV radiation

through the atmosphere and to interaction of UV radiation

with climate change. These include:

• using assumptions in modeling aerosol effects on surface UV radiation (e.g., asymmetry parameter or

phase function of aerosols);

• quantifying the effects of inhomogeneous clouds and

inhomogeneous topography on UV irradiance and

actinic flux at the surface; 

• investigating the role of UV from the perspective of

introducing changes in tropospheric photochemistry

and, thus, indirectly affecting the stratospheric composition;

• advancing the state of knowledge concerning interactions of climate change and UV radiation in time and

space, for example through changes in surface albedo

(deforestation, decrease of snow-ice cover, etc) and

cloudiness, which have not been extensively studied

so far;

• developing a better understanding of the UV light penetration into the water, which may be underlying partially snow-covered ice, and is important for marine

ecosystems;

• defining better the interaction of radiation and the surface, in particular the reflection of radiation by vegetation and the penetration into canopies; and

• better understanding and quantifying the effects of

pollutant gases and aerosols on surface UV radiation.

With respect to tools and technological issues,

research on UV radiation is expected to progress through

further development in:

• three-dimensional radiative transfer modeling, which

is becoming available but is not yet in widespread use

and requires substantial computing resources and

input parameters that are usually not available;

• more accurate UV radiation measurements that can

help to validate radiative transfer models and satellite-derived UV estimates, while establishment of new

stations, especially in the tropics, would provide better

coverage;

• radiative transfer studies for large solar zenith angles

and viewing angles, which are required for calculations for high latitudes and for the interpretation of

remote sensing observations;

• retrieval of trace gas amounts and the optical properties of aerosols using combinations of measurements

and modeling;

• correct consideration of polarization, which is important for the interpretation of radiance measurements;

and

• quality control procedures of broadband and multichannel narrowband radiometers deployed worldwide, which are not as good as required for addressing long-term changes and for studying radiative

processes in detail.

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SURFACE ULTRAVIOLET RADIATION

7.49

Appendix 7A

SPECTRAL DATA AVAILABLE FROM DATABASES

7A.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 31.58 130.56 01/01/1991

12 Sapporo Brewer Japan JMA 43.06 141.33 01/01/1991

14 Tateno / Tsukuba Brewer Japan JMA 36.06 140.10 01/01/1990

18 Alert Brewer Canada MSC 82.50 -62.32 09/06/1995

21 Edmonton / Stony Plain Brewer Canada MSC 53.55 -114.10 19/03/1992

24 Resolute Brewer Canada MSC 74.72 -94.98 14/03/1991

31 Mauna Loa Brewer United States MSC 19.53 -155.57 24/03/1997

31 Mauna Loa Brewer United States MSC-MLO 19.53 -155.57 01/01/1998

65 Toronto Brewer Canada MSC 43.78 -79.47 01/03/1989

67 Boulder Brewer United States EPA_UGA c 40.08 -105.25 01/01/1997

76 Goose Bay Brewer Canada MSC 53.30 -60.36 01/01/1997

77 Churchill Brewer Canada MSC 58.75 -94.07 30/03/1992

95 Taipei Brewer Taiwan CWBT 25.02 121.47 06/03/1992

101 Syowa Brewer Japan JMA -69.00 39.58 01/01/1993

190 Naha Brewer Japan JMA 26.20 127.68 01/01/1991

241 Saskatoon Brewer Canada MSC 52.11 -106.71 01/01/1991

290 Saturna Island Brewer Canada MSC 48.78 -123.13 03/10/1990

301 JRC Ispra (Varese) Brewer Italy JRC_EU 45.80 8.63 01/01/2001

306 Chengkung Brewer Taiwan CWBT 23.10 121.36 02/01/1992

307 Obninsk Brewer Russia IEM-SPA 55.12 36.30 06/05/1993

315 Eureka Brewer Canada MSC 80.04 -86.17 10/02/2001

319 Montreal (Dorval) Brewer Canada MSC 45.48 -73.75 06/03/1993

320 Winnipeg Brewer Canada MSC 49.90 -97.24 06/07/1992

321 Halifax (Bedford) Brewer Canada MSC 44.70 -63.61 01/07/1992

331 Poprad-Ganovce Brewer Slovakia SHMI 49.03 20.32 01/01/1994

332 Poyang Brewer Korea KMA 36.03 129.38 27/01/1994

338 Bratts Lake (Regina) Brewer Canada MSC 50.20 -104.70 01/01/1995

353 Reading Optronic United Kingdom UMIST 51.45 0.93 27/01/1994

377 Research Triangle Park Brewer United States EPA_UGA c 35.89 -78.87 10/09/1996

379 Gaithersburg Brewer United States EPA_UGA 39.13 -77.21 01/01/1996

380 Atlanta Brewer United States EPA_UGA 33.75 -84.41 30/05/1996

382 Riverside Brewer United States EPA_UGA 34.00 -117.34 01/01/1996

383 Big Bend Brewer United States EPA_UGA 29.31 -103.18 20/02/1997

384 Great Smokey Mountains Brewer United States EPA_UGA 35.60 -83.78 03/01/1997

385 Canyonlands Brewer United States EPA_UGA 38.47 -109.82 31/07/1997

386 Glacier Brewer United States EPA_UGA 48.51 -113.99 18/09/1997

387 Everglades Brewer United States EPA_UGA 25.39 -80.68 24/02/1997

388 Shenandoah Brewer United States EPA_UGA 38.52 -78.44 05/03/1997

389 Acadia Brewer United States EPA_UGA 44.38 -68.26 04/03/1998

390 Denali Brewer United States EPA_UGA 63.73 -148.97 04/10/1997
SURFACE ULTRAVIOLET RADIATION

7.50

391 Virgin islands Brewer Virgin Islands EPA_UGA 18.34 -64.79 25/05/1998

392 Rocky Mountain Natl. Park Brewer United States EPA_UGA 40.03 -105.53 30/05/1998

393 Olympic Brewer United States EPA_UGA 48.14 -123.40 11/12/1997

402 Theodore Brewer United States EPA_UGA 46.90 -103.38 29/09/1998

463 Sequoia Brewer United States EPA_UGA 36.49 -118.82 20/08/1998

465 Kilauea Brewer United States EPA_UGA 19.42 -155.29 18/02/1999

a Website: http://www.msc-smc.ec.gc.ca/woudc/ b JMA, Japan Meteorological Agency; MSC, Meteorological Service of Canada; EPA_UGS, Environmental Protection Agency, University of Georgia,

Athens, USA; CWBT, Central Weather Bureau of Taiwan; JRC_EU, Joint Research Centre of the European Union; IEM-SPA, Institute of Experimental

Meteorology-Scientific Hydrometeorological Institute; SHMI, Slovakian Hydrometeorological Institute; KMA, Korean Meteorological

Administration; UM, University of Manchester. c The Brewer spectrophotometers operated by EPA have ceased operation after 2004.
SURFACE ULTRAVIOLET RADIATION

7.51

7A.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 Feb. 98

3 Uccle Brussels Jobin Yvon Belgium IASB 50.80 4.36 Mar. 93

4 Hradec Kralove Brewer Czech Republic CHI 50.19 15.83 Jan. 94

5 Jokioinen Brewer Finland FMI 60.81 23.50 Mar. 95

6 Sodankylä Brewer Finland FMI 67.37 26.63 Apr. 90

7 Briançon Jobin Yvon France USTL 44.90 6.65 Sep. 99

8 Villeneuve d.Asq Jobin Yvon France USTL 50.37 3.01 May. 97

9 Neuherberg Bentham Germany BfS 48.22 11.58 Mar. 00

10 Offenbach Bentham Germany BfS 50.10 8.75 Jan. 00

11 Hohenpeissenberg Brewer Germany DWD 47.80 11.02 Jan. 95

12 Lindenberg Brewer Germany DWD 52.22 14.12 Jan. 95

13 Potsdam Brewer Germany DWD 52.36 13.08 Jan. 95

14 Garmisch-Partenkirchen Bentham Germany IFU 47.48 11.07 Apr. 94

15 Zugspitze Bentham Germany IFU 47.42 10.98 Jul. 97

16 Thessaloniki Brewer Greece LAP 40.52 22.97 Oct. 89

17 Lampedusa Brewer Italy ENEA 35.50 12.60 Jan. 98

18 Ispra Brewer Italy JRC 45.81 8.63 Aug. 91

19 Rome Brewer Italy URO 41.90 12.52 Feb. 92

20 Andøya Bentham Norway NILU 69.48 16.02 Feb. 98

21 Oesteraas Bentham Norway NRPA 59.92 10.75 Jun. 99

22 Trondheim Optronic Norway NTNU 64.43 10.47 Aug. 97

23 Tromsø Brewer Norway UT 69.66 18.93 May. 95

24 Belsk Brewer Poland IGFPAS 51.83 20.78 Jan. 93

25 Azores Island Brewer Portugal IM 38.66 −27.22 Jul. 92

26 Funchal (Madeira Is.) Brewer Portugal IM 32.64 −16.89 Jan. 90

27 Lisbon Brewer Portugal IM 38.77 −9.15 Jan. 90

28 Penhas Douradas Brewer Portugal IM 40.42 −7.55 Oct. 94

29 Izana Brewer Spain INM 28.49 −16.50 Jan. 95

30 Norrköping Brewer Sweden SMHI 58.58 16.15 Sep. 91

31 Vindeln Brewer Sweden SMHI 64.23 19.77 Jul. 96

32 De Bilt Brewer Netherlands KNMI 52.10 5.18 Jan. 94

33 RIVM (Bilthoven) Dilor XY-50 Netherlands RIVM 52.12 5.20 Jan. 96

34 Reading Optronics United Kingdom UM 51.45 0.93 Jan. 93

Only permanent stations are listed; campaign data are not included.

a Website: http://uvdb.fmi.fi/uvdb/index.html b BOKU, Universität für Bodenkultur Wein; IASB, Institut d’Aéronomie Spatiale de Belgique; CHI, Czech Hydrometeorological Institute; FMI,

Finnish Meteorological Institute; USTL, Université des Sciences et Technologies de Lille; BfS, Bundesamt für Strahlenschutz; DWD, Deutscher

Wetterdienst; IFU, Atmosphärische Umweltforschung; LAP, Laboratory of Atmospheric Physics; ENEA, Ente per le Nuove Technologie, l’Energia e

l’Ambiente; JRC, Joint Research Centre, European Union; URO, University of Rome; NILU, Norsk Institutt for Luftforskning; NRPA, Norwegian

Radiation Protection Authority; NTNU, Norges Teknisk-Naturvitenskapelige Universitet; UT, University of Tromsø; IGFPAS, Institute of Geophysics,

Polish Academy of Sciences; IM, 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; UM,

University of Manchester.
SURFACE ULTRAVIOLET RADIATION

7.52

7A.3 U.S. National Science Foundation UV Monitoring Network a

No. Station Name Instrument Country Institute b Lat. (°N) Long. (°E) Start Date c

1 Arrival Heights Biospherical Antarctica NSF −77.83 166.67 13/12/1989

2 Palmer Biospherical Antarctica NSF −64.77 −64.05 14/03/1990

3 South Pole d Biospherical Antarctica NSF −90.00 80.00 31/01/1991

4 Ushuaia Biospherical Argentina NSF −54.82 −68.32 13/09/1990

5 San Diego Biospherical United States NSF 32.77 −117.20 28/10/1992

6 Barrow Biospherical United States NSF 71.32 −156.68 14/01/1991

7 Summit Biospherical Greenland NSF 72.58 −38.46 15/08/2004

a Website: http://www.biospherical.com/NSF/ b Stations 1-6 are equipped with SUV-100 spectroradiometers. Station 7 (Summit) is equipped with a SUV-150B spectroradiometer. c Date from which date data are available. Instruments may have been installed before this date. d The system at the South Pole was relocated by approximately 200 m in January 1997. Due to the proximity of the instrument to the geographic South

Pole and the movement of the polar ice cap over the bedrock, the azimuth position of the instrument has changed from 84.5°E in January 1997 to

72.5°E in January 2006.
SURFACE ULTRAVIOLET RADIATION

7.53

Appendix 7B

INTERNET ADDRESSES FOR UV SITES

7B.1 General UV Information

WMO UV radiation page http://www.srrb.noaa.gov/UV/

TOMS (ozone and UV satellite data) http://toms.gsfc.nasa.gov/

OMI (Ozone Monitoring Instrument) http://aura.gsfc.nasa.gov/instruments/omi/index.html

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.niwascience.co.nz/services/uvozone/

EPA, U.S. http://www.epa.gov/uvnet/

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/

World Radiation Centre http://wrdc-mgo.nrel.gov/

7B.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 and http://i115srv.vu-wien.ac.at/uv/COST726/Cost726.htm

Technical Research (COST-726)

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/

Stratospheric-Climate Links with Emphasis on the Upper http://www.ozone-sec.ch.cam.ac.uk/scout_o3/index.html

Troposphere and Lower Stratosphere (SCOUT-O3)

UV Index & UV dose based on GOME (TEMIS) http://www.temis.nl/uvradiation/

7B.3 Radiative Transfer Codes

Tropospheric Ultraviolet Visible (TUV) (Madronich and http://cprm.acd.ucar.edu/Models/TUV/

Flocke, 1997)

Library for Radiative Transfer (libRadTran) (Mayer and http://www.libradtran.org/

Kylling, 2005)

System for Transfer of Atmospheric Radiation (STAR) http://www.meteo.physik.uni-muenchen.de/

(Koepke et al., 2005) strahlung/uvrad/Star/starprog.html

Fast and Easy Radiative Transfer (FASTRT) http://zardoz.nilu.no/~olaeng/fastrt/fastrt.html

Santa Barbara Discrete-Ordinate Atmospheric Radiative http://arm.mrcsb.com/sbdart/

Transfer (SBDART) (Ricchiazzi et al., 1998)

Streamer model (Key, 1999) http://stratus.ssec.wisc.edu/streamer/streamer.html

SCIATRAN code (Rozanov et al., 2005) http://www.iup.physik.uni-bremen.de/sciatran/

3-D Spherical Harmonics Discrete Ordinate Method http://nit.colorado.edu/~evans/shdom.html

(SHDOM) model (Evans, 1998)

7B.4 Extraterrestrial Spectra

Solar Ultraviolet Spectral Irradiance Monitor (SUSIM) http://wwwsolar.nrl.navy.mil/susim_atlas_data.html

Solar Stellar Irradiance Comparison Experiment (SOLSTICE) http://lasp.colorado.edu/solstice/

McMath/Pierce at Kitt Peak ftp://ftp.noao.edu/fts/fluxatl/

The New Synthetic Gueymard Spectrum (Gueymard, 2004) http://rredc.nrel.gov/solar/spectra/am0/special.html#newgueymard
SURFACE ULTRAVIOLET RADIATION

7B.5 Internet Sites with UV Information

Argentina http://www.conae.gov.ar/uv/uv.html

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.uv-index.at/

Belgium http://ozone.meteo.be/uv/index.php

Brazil http://www.master.iag.usp.br/ind.php?inic=00&prod=indiceuv

Canada http://www.msc-smc.ec.gc.ca/education/uvindex/forecasts/

index_e.html

Catalonia http://www.meteocat.com/marcs/marcos_previsio/marcs_uvi.htm

Chile http://www.meteochile.cl/

Czech Republic http://www.chmi.cz/meteo/ozon/o3uvb-e.html

Denmark http://www.dmi.dk/vejr/index_sol.html

European Commission (JRC) http://ecuv.jrc.it/

Finland http://www.fmi.fi/research_atmosphere/ atmosphere_2.html

France http://www.soleil.info/main.php

Germany (DWD) http://www.uv-index.de/

Germany (BfS) http://www.bfs.de/uv

Greece http://www.uvnet.gr

Hong Kong http://www.hko.gov.hk/wxinfo/uvfcst/uvfcst.htm

Israel http://www.ims.gov.il/en2.htm#1

Italy http://www.lamma.rete.toscana.it/previ/eng/uvhtm/uvnew0.html

Japan http://www.jma.go.jp/en/uv/index.html

Luxembourg http://www.restena.lu/meteo_lcd/

Netherlands http://www.temis.nl/uvradiation/index.html

New Zealand http://www.niwascience.co.nz/services/uvozone/

Norway http://uvnett.nrpa.no/uv/

Poland http://www.imgw.pl/wl/internet/zz/index.html

Portugal http://www.meteo.pt/en/previsao/uv/prev_uv_d0.jsp

Slovenia http://www.rzs-hm.si/zanimivosti/UV.htm

South Africa http://www.weathersa.co.za/uv/uv.jsp

Spain http://infomet.am.ub.es/uv_i_ozo/uvi.html

Sweden http://www.smhi.se/weather/uvindex/index.htm

Switzerland http://www.uv-index.ch/de/home.php

Taiwan http://www.epa.gov.tw/monitoring/1-1/uv.htm

Uunited Kingdom http://www.m-office.gov.uk/weather/uv/

United States (NOAA) http://www.cpc.ncep.noaa.gov/products/stratosphere/uv_index/

United States (EPA) http://www.epa.gov/sunwise/uvindex.html

World Health Organization (WHO) http://www.who.int/mediacentre/factsheets/fs271/en/index.html

Intersun (WHO) http://www.who.int/uv/en/

7.54
CHAPTER 8

Halocarbon Scenarios, Ozone Depletion 

Potentials, and Global Warming Potentials

Lead Authors:

J.S. Daniel

G.J.M. Velders 

Coauthors:

A.R. Douglass

P.M.D. Forster

D.A. Hauglustaine

I.S.A. Isaksen

L.J.M. Kuijpers

A. McCulloch

T.J. Wallington

Contributors:

P. Ashford

S.A. Montzka

P.A. Newman

D.W. Waugh 

Final Release: February 2007

From Scientific Assessment of Ozone Depletion: 2006

CHAPTER 8

HALOCARBON SCENARIOS, OZONE DEPLETION POTENTIALS,

AND GLOBAL WARMING POTENTIALS

Contents

SCIENTIFIC SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.1

8.1 INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.5

8.2 HALOCARBON LIFETIMES, OZONE DEPLETION POTENTIALS, AND GLOBAL

WARMING POTENTIALS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.5

8.2.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.5

8.2.2 Ozone Depletion Potentials . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.5

8.2.2.1 Atmospheric Lifetimes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.6

8.2.2.2 Fractional Release Factors . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.6

8.2.2.3 Ozone Destruction Effectiveness . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.7

8.2.2.4 ODP Values . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.8

8.2.3 Direct Global Warming Potentials . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.8

8.2.4 Degradation Products and Their Implications for ODPs and GWPs . . . . . . . . . . . . . . . . . . . . . . . . . 8.13

8.3 FUTURE HALOCARBON SOURCE GAS CONCENTRATIONS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .8.13

8.3.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.13

8.3.2 Baseline Scenario (A1) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.15

8.3.2.1 Emissions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.15

8.3.2.2 Mixing Ratios . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.21

8.3.2.3 Equivalent Effective Stratospheric Chlorine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.25

8.3.3 Alternative Projections . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.28

8.3.3.1 Emissions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.28

8.3.3.2 Mixing Ratios . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.28

8.3.3.3 Equivalent Effective Stratospheric Chlorine . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.30

8.3.4 Uncertainties in ODS Projections . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.32

8.4 OTHER PROCESSES RELEVANT TO FUTURE OZONE EVOLUTION . . . . . . . . . . . . . . . . . . . . . . . . . . 8.34

8.5 INDIRECT GWPS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.35

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.36

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.1

SCIENTIFIC SUMMARY

Ozone Depletion Potentials and Global Warming Potentials

• The effectiveness of bromine compared with chlorine on a per-atom basis for global ozone depletion, typically referred to as α, has been re-evaluated upward from 45 to a value of 60. The calculated values from three

independent two-dimensional models range between 57 and 73, depending on the model used and depending on the

assumed amount of additional bromine added to the stratosphere by very short-lived substances (VSLS).

• Semi-empirical Ozone Depletion Potentials (ODPs) have been re-evaluated, with the most significant change

being a 33% increase for bromocarbons due to the update in the estimate for the value of α. A calculation

error in the previous Assessment, which led to a 13% overestimate in the halon-1211 ODP, has been corrected.

• Direct Global Warming Potentials (GWPs) have been updated to account for revised radiative efficiencies

(HFC-134a, carbon tetrafluoride (CF4), HFC-23, HFC-32, HFC-227ea, and nitrogen trifluoride (NF3)) and

revised lifetimes (trifluoromethylsulfurhexafluoride (SF5CF3) and methyl chloride (CH3Cl)). In addition, the

direct GWPs for all compounds have been affected by slight decreases in the carbon dioxide (CO2) absolute Global

Warming Potentials for various time horizons.

• Indirect GWPs have been updated primarily to reflect the later return of ozone-depleting substances (ODSs)

to 1980 levels estimated in this Assessment and to account for the increased bromine efficiency factor. Direct

and indirect GWPs are presented separately due to concern over the appropriateness of combining them into net

GWPs for application in certain climate change issues.

Projections of Halocarbon Abundances and Implications for Policy Formulation 

• Several of the options evaluated for accelerating the future reduction of ODS abundances demonstrate

greater effectiveness than assessed previously. These options are assessed using equivalent effective stratospheric chlorine (EESC) derived from projections of the atmospheric abundances of ODSs based on historic emissions, observations of concentrations, reported production, estimates of future production, and newly available

estimates of the quantities of ODSs present in products in 2002 and 2015.

• The date when EESC relevant to midlatitude ozone depletion returns to pre-1980 levels is 2049 for the baseline (A1) scenario, about 5 years later than projected in the previous Assessment. This later return is primarily due to higher estimated future emissions of CFC-11, CFC-12, and HCFC-22. The increase in CFC emissions

is due to larger estimated current bank sizes, while the increase in HCFC-22 emissions is due to larger estimated

future production.

• For the polar vortex regions, the return of EESC to pre-1980 conditions is projected to occur around 2065,

more than 15 years later than the return of midlatitude EESC to pre-1980 abundances. This later return is

due to the older age of air in the lower stratosphere inside the polar vortex regions. This metric for the polar vortex

regions has not been presented in previous ozone Assessments.

• Three classes of hypothetical cases are presented here to illustrate the maximum potential for reducing midlatitude EESC if anthropogenic production or emission were eliminated after 2006 and if the existing banks

at the end of 2006 were fully eliminated. The sizes of the banks considered for elimination are equal to the total

estimated production through 2006 minus all estimated emissions through 2006. These cases are not mutually

exclusive, and separate effects of the elimination of production, emissions, and banks are not additive.

The table below shows the percentage reductions in integrated EESC relative to the baseline (A1) scenario that can

be achieved in these hypothetical cases (see table footnote 1). EESC is integrated above the 1980 level from 2007

until it returns to the 1980 level (about 2050), and any potential contribution from VSLSs is neglected.
8.2

• The percentage reductions in EESC for halons and CFCs, integrated from 2007 until the 1980 level is

reattained (shown in column B), are larger than previously reported. This is because recent bank estimates

from the Intergovernmental Panel on Climate Change/Technology and Economic Assessment Panel (IPCC/TEAP)

2005 Special Report are significantly larger and likely more reliable than values presented in previous Assessments

for CFC-11, CFC-12, and halon-1211.

• The percentage reductions in integrated EESC for HCFCs shown in hypothetical case A are larger than previously reported. This is because of significantly larger estimates of future HCFC-22 production in Article-5 countries.

• Evidence suggests that anthropogenic emissions have a stronger influence on atmospheric mixing ratios of

CH3Br than reported in the previous Assessment. This enhanced sensitivity leads to a greater impact of an elimination of these emissions on EESC. The effect of an elimination of the anthropogenic methyl bromide emission on

integrated EESC is about 50% larger than it would have been if the same fractional anthropogenic emission were

used as was assumed in the previous Assessment.

• The hypothetical elimination of all emissions of ODSs after 2006 (hypothetical case C) would accelerate the

year EESC is expected to drop below the 1980 value by about 15 years, from 2049 to 2034. The hypothetical

elimination of all emissions from production of ODSs after 2006 (hypothetical case A) would accelerate it by about

6 years, to 2043.

• Two additional hypothetical cases of critical and exempted (quarantine and pre-shipment, QPS) uses of

methyl bromide were considered. In the analysis of both cases, EESC is integrated above the 1980 level from

2007 until it returns to the 1980 level. A methyl bromide phase-out has been in effect since 2005 in developed countries, with critical-use exemptions granted in 2005 and 2006 at levels that are 30-40% of the 2003-2004 production

levels. The size of the critical-use exemptions is similar to the estimated use of methyl bromide for QPS use.

– If critical-use exemptions continue indefinitely at the 2006 level compared with a cessation of these

exemptions in 2010 or 2015, midlatitude integrated EESC would increase by 4.7% or 4.0%, respectively.

Reduction in Integrated EESC Relative to Baseline Scenario A1 (%)

Compound or Hypothetical Case A: Hypothetical Case B: Hypothetical Case C:

Compound Group All Emissions All Emissions All Emissions

Eliminated from Eliminated from Eliminated after 2006

Production after 2006 Existing Banks at

End of 2006

CFCs 0.3 11 11

Halons 0.5 14 14

CCl4 3 (a) 3

CH3CCl3 0.2 (a) 0.2

HCFCs 12 4 16

CH3Br (anthropogenic) 5 (a) 5

1 Hypothetical case A corresponds to the elimination of all emissions from production after 2006. Hypothetical case B corresponds to the elimination

of all emissions from banks existing at the end of 2006 (for example, capture and destruction). Hypothetical case C corresponds to the elimination

of all emissions after 2006 and is approximately equal to the sum of columns A and B.

(a) For these compounds, banks are uncertain and therefore emissions are equated to production in these calculations.

HALOCARBON SCENARIOS, ODPs, AND GWPs
HALOCARBON SCENARIOS, ODPs, AND GWPs

8.3

– If production of methyl bromide for QPS use were to continue at present levels and cease in 2015, midlatitude integrated EESC would decrease by 3.2% compared with the case of continued production at

present levels.

Uncertainties and Sensitivities

• Recent bank estimates from the IPCC/TEAP (2005) report are significantly larger and likely more reliable

than values presented in previous Assessments for some compounds (CFC-11, CFC-12, halon-1211, and

halon-1301). However, there remain potential shortcomings in these new bank estimates that lead to uncertainties

in assessing the potential for reducing integrated future EESC.

• While the future evolution of the ozone layer depends largely on the abundances of ozone-depleting substances, changes in climate arising from natural and anthropogenic causes are likely to also play an important role. Many of these changes are expected to induce spatially dependent chemical and dynamical perturbations

to the atmosphere (Chapters 5 and 6), which could cause ozone to return to 1980 levels earlier or later than when

EESC returns to 1980 levels. Furthermore, the changes may alter the lifetimes of the important ozone-depleting

substances.

8.1 INTRODUCTION

This chapter provides an update of future halocarbon

mixing ratio estimates and of Ozone Depletion Potentials

and Global Warming Potentials. Future scenarios of halocarbons are constructed based on the current Montreal

Protocol and build on the previous Scientific Assessments of Ozone Depletion (WMO, 1999, 2003) and on the

recently published Intergovernmental Panel on Climate

Change/Technology and Economic Assessment Panel

(IPCC/TEAP) Special Report on Safeguarding the Ozone

Layer and the Global Climate System: Issues Related to

Hydrofluorocarbons and Perfluorocarbons (IPCC/TEAP,

2005). A number of hypothetical cases are presented to

demonstrate how halocarbon production and emission, and

halocarbons present in current applications, may contribute

to the future ozone-depleting substance (ODS) loading of

the atmosphere. Uncertainties in future emissions and

atmospheric concentrations are discussed, as are the important differences from previous Assessments.

8.2 HALOCARBON LIFETIMES, OZONE

DEPLETION POTENTIALS, AND GLOBAL

WARMING POTENTIALS

8.2.1 Introduction

Halocarbons released from the Earth’s surface

become mixed in the lower atmosphere and are transported

into the stratosphere by normal air circulation patterns.

They are removed from the atmosphere by photolysis, reaction with hydroxyl (OH) radicals (for compounds containing

carbon-hydrogen bonds), and for some compounds, uptake

by the oceans. Halocarbon molecules that are transported

to the stratosphere deposit their degradation products

directly. A small fraction of the degradation products from

halocarbons destroyed before leaving the troposphere can

also be transported to the stratosphere (see Chapter 2). The

final degradation products are inorganic halogen species

containing fluorine, chlorine, bromine, and iodine atoms.

A significant fraction of inorganic chlorine, bromine, and

iodine are in the form of X (X = Cl, Br, or I) and XO that

participate in efficient ozone destruction in the stratosphere.

Fluorine atoms, in contrast, are rapidly converted into

hydrogen fluoride (HF), which is a stable reservoir and prevents fluorine from contributing to ozone destruction to

any significant degree. Iodine atoms participate in catalytic

ozone destruction cycles, but rapid tropospheric loss of

iodine-containing compounds limits the amount of iodine

reaching the stratosphere (see Chapter 2).

All halocarbons also absorb terrestrial radiation

(long-wavelength infrared radiation emitted from the

Earth’s surface and by the atmosphere) and contribute to

the radiative forcing of climate change. The relative contribution of individual compounds to stratospheric ozone

depletion and global warming can be characterized by their

Ozone Depletion Potentials (ODPs) and Global Warming

Potentials (GWPs), respectively. ODPs and GWPs have

been used in past ozone and climate Assessments (IPCC,

1990; 1995; 1996; 2001; 2005; WMO, 1995; 1999; 2003)

and in international agreements such as the Montreal

Protocol and the Kyoto Protocol.

8.2.2 Ozone Depletion Potentials

Ozone Depletion Potentials are indices that provide

a simple way to compare the relative ability of various

ODSs to destroy stratospheric ozone (Fisher et al., 1990;

Solomon et al., 1992; Wuebbles, 1983). The concept of

the ODP has been discussed extensively in previous WMO

reports (WMO, 1995; 1999; 2003). ODPs are often calculated assuming steady-state conditions with constant

emissions (for compounds that are removed by linear

processes, this is equivalent to assuming an emission

pulse and integrating over the entire decay of the compound (Prather, 1996; Prather, 2002)) and are not

dependent on time. Time-dependent ODPs can also be

calculated (Solomon and Albritton, 1992), which reflect

the different time scales over which the compound and

reference gas (the chlorofluorocarbon CFC-11) liberate

chlorine and bromine into the stratosphere. Compounds

that have shorter (longer) atmospheric lifetimes than CFC11 have ODPs that decrease (increase) with increasing

integration time.

The ODPs considered here are steady-state ODPs,

integrated quantities that are distinct for each halocarbon

species. The ODP of a well-mixed ozone-destroying

species i is given by:

ODPi = (8-1)

This quantity can be calculated using computer models,

with the accuracy depending on the model’s ability to

simulate the distribution of the considered halocarbon

and the ozone loss associated with it. Because ODPs are

defined relative to the ozone loss caused by CFC-11, the

ODP values demonstrate less sensitivity to photochemical

modeling errors than do absolute ozone loss calculations.

Taking advantage of this reduced sensitivity,

Solomon et al. (1992) proposed a semi-empirical approach

to estimate ODPs that approximates and simplifies the

accurate representation of ozone photochemistry and

provides an observational constraint to the model-based

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.5

global O3 loss due to unit mass emission of i

global O3 loss due to unit mass emission of CFC-11
results. In this approach, measurements of correlations

between halocarbons are used to evaluate chlorine and

bromine relative stratospheric release values. These

release rates are used with lifetimes, molecular weights,

and the number and type of halogen atoms per molecule

to estimate the effect of a small source gas pulse emission

on stratospheric ozone depletion. In the case of bromine

and iodine, the catalytic efficiency for ozone destruction

relative to chlorine is needed and differs from unity partly

due to the different partitioning of the halogen chemical

families. For long-lived chlorocarbons and bromocarbons

that are well mixed in the troposphere, the semi-empirical

ODP definition can be expressed by:

(8-2)

where f is the fractional halogen release factor, α is the

relative effectiveness of bromine compared with chlorine

for ozone destruction, τ is the global lifetime, M is the

molecular weight, and nCl (nBr) is the number of chlorine

(bromine) atoms contained in the compound. CFC-11 subscripts indicate quantities for CFC-11, while i subscripts

denote quantities pertaining to the compound for which

the ODP is desired. This equation has been altered slightly

from the form given in the previous Assessment (Equation

1-6, WMO, 2003) because it was less clear how to account

for compounds with both chlorine and bromine atoms

using the previous form. For very short-lived substances

(VSLSs), the location and season of emission affect the

amount of halogen that can make it to the stratosphere,

making Equation (8-2) an inappropriate choice for ODP

estimates of these gases (see Chapter 2 of this Assessment, as well as Chapter 2 of WMO, 2003).

8.2.2.1 ATMOSPHERIC LIFETIMES

The lifetimes of atmospheric trace gases given in

Tables 8-1 and 8-2 have been assessed in Chapter 1. A

detailed explanation of global lifetimes can be found in

WMO (2003). For these reported lifetimes, it is assumed

that the gases are uniformly mixed throughout the troposphere. This assumption is less accurate for gases with

lifetimes <1/2 year, and is but one reason why single values

for global lifetimes, ODPs, or GWPs are less appropriate

for such short-lived gases (see Chapter 2). However, the

majority of ozone-depleting substances and their replacements have atmospheric lifetimes greater than 2 years,

much longer than tropospheric mixing times; hence their

lifetimes, ODPs, and GWPs are not significantly altered

by the location of sources within the troposphere.

8.2.2.2 FRACTIONAL RELEASE FACTORS

The distributions of species in the stratosphere

depend on the competition between local photochemical

removal processes and transport processes that carry the

material from the entry point (mainly at the tropical

tropopause) through, and out of, the stratosphere. Once a

halogen source gas is in the stratosphere, release of a

halogen atom from the source gas can occur through

photolysis or chemical reaction. As already described in

previous WMO reports (WMO, 2003), the fraction, f, of

halocarbon i converted to an inorganic form by some time

at a given location in the stratosphere can be given by:

(8-3)

where ρi(x,y,z,t) denotes the mixing ratio of the halocarbon at a given stratospheric location (x,y,z) at time t,

and ρi,entry is the mixing ratio of species i in the air parcel

when it entered the stratosphere. The value of ρi,entry can

be estimated from knowledge of the age (time since

entering the stratosphere) of the parcel at location (x,y,z)

and the tropospheric time series of species i. With Equation (8-3), measurements of the halocarbon distributions

within the stratosphere then can be used to define fi(x,y,z,t).

The relative fractional release term used in Equation (8-

2) is the ratio fi / fCFC-11, which is a measure of the local

fractional release of inorganic halogen compounds relative to the fractional release of CFC-11. Conceptually,

the fractional release factor should be globally integrated.

In practice, a more limited range of measured correlations representing mid- to high latitudes, where ozone is

highly sensitive to changes in the local photochemical

removal rate, is generally used (e.g., Schauffler et al.,

2003). The fractional release factors used in this report

are given in Table 8-1. Except for the value for CFC114, which is taken from Schauffler et al. (2003), all

values are taken from WMO (2003). Most of these values

have been derived from stratospheric observations, with

models used to estimate the fractional release factors for

a few source gases (Solomon et al., 1992; WMO, 2003).

The fractional release factors for hydrochlorofluorocarbons HCFC-141b and HCFC-142b included in Table

8-1 and used in the previous Assessment differ from the

values derived by Schauffler et al. (2003) from stratospheric observations. These observations suggest values

smaller by a factor of 3.1 and 4.5 for HCFC-141b and

HCFC-142b, respectively. Because of the large growth

rates of these compounds at the time the measurements

were made and the resulting sensitivity of the fractional

release values to age-of-air estimates, the authors of the

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.6

ODP n n f

f

M i Br i Cl i

i

CFC

i

CFC

CFC = + ( ) − −

− α

τ

τ , ,

11 11

11 1

M 3 i

f xyzt xyzt

i

i entry i

i entry

,,, (,,,) ,

,

( ) = ρ ρ −

ρ
previous Assessment did not adopt these new values. With

no additional estimates available since this disagreement

was discussed in the previous Assessment, the discrepancy remains an unresolved issue and we continue to use

the older, model-derived values. Model estimates for all

other gases agree with the values estimated from observations to within 2 times the quoted error bars of the observational estimates (Schauffler et al., 2003) except for

HCFC-22. The HCFC-22 value based on observations is

about 17% lower than calculated in Solomon et al. (1992).

8.2.2.3 OZONE DESTRUCTION EFFECTIVENESS

Although the relative effectiveness of bromine

compared with chlorine for ozone depletion, referred to

as α, is treated as a single, fixed quantity in Equation 8-

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.7

Table 8-1. Lifetimes, relative fractional halogen release factors, and Ozone Depletion Potentials for

halocarbons. ODPs recommended in this Assessment and ODPs adopted in the Montreal Protocol are

included.

Halocarbon * Lifetime Relative Fractional Semi-Empirical ODP in Montreal

(years) Release Factor 1 ODP Protocol

Annex A-I

CFC-11 45 1 1.0 1.0

CFC-12 100 0.60 1.0 1.0

CFC-113 85 0.75 1.0 0.8

CFC-114 300 0.28 ± 0.02 2 1.0 1.0

CFC-115 1700 0.44 † 0.6

Annex A-II

Halon-1301 65 0.62 16 10.0

Halon-1211 16 1.18 7.1 3 3.0

Halon-2402 20 1.22 11.5 6.0

Annex B-II

Carbon tetrachloride 26 1.06 0.73 1.1

Annex B-III

Methyl chloroform 5.0 1.08 0.12 0.1

Annex C-I

HCFC-22 12.0 0.35 0.05 0.055

HCFC-123 1.3 1.11 0.02 0.02

HCFC-124 5.8 0.52 0.02 0.022

HCFC-141b 9.3 0.72 0.12 0.11

HCFC-142b 17.9 0.36 0.07 0.065

HCFC-225ca 1.9 1.1 0.02 0.025

HCFC-225cb 5.8 0.5 0.03 0.033

Annex E

Methyl bromide 0.7 1.12 0.51 0.6

Others

Halon-1202 2.9 1.7 4

Methyl iodide see Chapter 2 see Chapter 2

Methyl chloride 1.0 0.80 0.02

* Chemical formulae for the halocarbons are listed in Table 8-2 and also in Appendix C of this Assessment.

† Model-derived value, WMO (2003). 1 From WMO (2003), Table 1-4, except for the value for CFC-114. For the EESC calculations in Section 1.8 of WMO (2003), slightly different relative

fractional release factors were used by mistake for the halons. The values given here are used for the calculations presented in this Assessment. 2 From Schauffler et al. (2003). 3 The ODP of halon-1211 should have been reported as 5.3 in the previous Assessment (WMO, 2003), but was incorrectly reported as 6.0 due to a

calculation error. 4 WMO (2003), with adjustment for updated α value.
2 and in the calculation of EESC elsewhere in this

Assessment Report, it represents a globally integrated

result with sensitivity to many factors, including the

kinetic parameters for chlorine and bromine species, the

amount of inorganic bromine and inorganic chlorine in

the background atmosphere, and atmospheric transport.

The value of α was assumed to be 45 in the previous

Assessment, based partly on the results from Daniel et

al. (1999) and Ko et al. (1998).

The value of α has been recalculated for this Assessment with the Atmospheric Environmental Research, Inc.

(AER), the Leeds-Bremen, and the University of Illinois

two-dimensional models, which are discussed in Chapters

2 and 6. Using Jet Propulsion Laboratory (JPL) JPL-05

kinetics (preliminary version of Sander et al., 2006), the

AER (D. Weisenstein, private communication) and LeedsBremen (B.-M. Sinnhuber, private communication)

models calculate global values of 61 and 71, respectively.

The University of Illinois (D. Wuebbles, private communication) model suggests a value of 57 using JPL-02

kinetics (Sander et al., 2002). The change in kinetics recommendations from 1997 are thought to play a relatively

minor role in the increased estimates of the α value. The

Leeds-Bremen model, when using JPL-97 rates (DeMore

et al., 1997), calculates a reduced α value of 64 compared

with the 71 with the JPL-05 kinetics. The AER model

value is also higher than its previously calculated value

when using JPL-97 rates; however, changes have been

made in the model in addition to kinetic rates that also

could have affected the α value. The AER value would

have been 52 with the JPL-97 rates if a methyl bromide

(CH3Br) stratosphere fractional release value of 1.12 relative to CFC-11 was applied to the results of Ko et al.

(1998); however, the calculated release value from the current AER model is larger than 1.12, perhaps implying a

value for α lower than 52 (M. Ko, private communication). Thus, there is no simple way to be certain of the precise AER α value with JPL-97 kinetics.

The reasons for the variations in the calculated α

values from these models apparently arise not from

kinetics changes alone, but also from differences in other

model processes, such as transport, the ozone loss spatial distribution, etc. The differences have not yet been

explained in the literature. Nevertheless, due to the consistently larger values recently calculated with these welldocumented models compared with the previously

assumed value of 45, and because the Leeds-Bremen

model suggests an even higher value when additional stratospheric bromine is considered to account for VSLSs (a

value of 73 for an additional 6 ppt of stratospheric

bromine), we now recommend an α value of 60 for global

ozone destruction. However, we emphasize the relatively

large model-dependent range of values discussed in the

previous paragraph.

In Section 8.3, a value of 65 is used for α when

simulating Antarctic conditions, based on the results of

Chipperfield and Pyle (1998) for Arctic conditions. These

calculations assumed DeMore et al. (1994) kinetic rates; a

calculation using updated rates is not currently available.

For fluorine, the relative effectiveness compared

with chlorine for ozone destruction is negligibly small

based on the results from Ravishankara et al. (1994) and

Wallington et al. (1995).

8.2.2.4 ODP VALUES

The most significant change in ODPs since the

previous Assessment is for the bromocarbons, due to the

increase in the recommended value of α from 45 to 60.

This α increase leads directly to an increase in the semiempirical ODPs of all bromocarbons of 33%. The ODP

of halon-1211 was also incorrectly reported in the previous Assessment to be 6.0 due to a calculation error;

it should have been reported as 5.3. The ODPs of the

chlorocarbons remain the same as those reported in

WMO (2003), except for CFC-114. In WMO (2003), the

ODP for CFC-114 was derived from a model; because

the fractional chlorine release and lifetime are available,

the semi-empirical ODP is reported here. Although the

recommended value of the methyl chloride (CH3Cl) lifetime has been decreased from 1.3 to 1.0 (Chapter 1), the

ODP remains unchanged to the reported precision.

8.2.3 Direct Global Warming Potentials

Halocarbons absorb terrestrial radiation (long

wavelength infrared radiation emitted from the Earth’s

surface) and contribute to the radiative forcing of the climate system. They generally have strong absorption features in the atmospheric window region (at approximately

8-12 micrometers) where there is little absorption by

atmospheric gases. This absorption reduces the amount

of outgoing energy from the Earth-atmosphere system and

leads to a direct radiative forcing. It is this forcing that

plays an important role in the calculations of the direct

GWPs discussed in this section.

The change in net radiation at the tropopause caused

by a given change in greenhouse gas concentration or mass

is referred to as radiative efficiency. Radiative efficiency

has units of W m-2 ppb-1 or W m-2 kg-1; it is calculated using

radiative transfer models of the atmosphere and depends

upon the strength and spectral position of a compound’s

absorption bands. The Absolute Global Warming Potential

(AGWP) has units of W m-2 ppb-1 yr or W m-2 k-1 yr and

quantifies the future integrated radiative forcing of a unit

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.8
mass pulse emission of a greenhouse gas; it can be defined

as:

(8-4)

where Fx is the radiative forcing per unit mass of species

x, x(t) describes the decay with time of a unit pulse of compound x, and t9 is the time horizon considered. To compare the relative integrated effects of various compounds

on climate, the Global Warming Potential concept was

developed. The Global Warming Potential (IPCC, 1990;

2001) can be defined as:

(8-5)

where FCO2 is the radiative forcing of carbon dioxide

(CO2), R(t) is the response function that describes the

decay of an instantaneous pulse of CO2, and the decay of

the pulse of compound x has been rewritten assuming it

obeys a simple exponential decay curve determine by a

response time of τx. Both Fx and FCO2 are generally given

in units of W m-2 kg-1. The unit pulse response terms lead

to a dependence of GWPs on the integration time horizon;

compounds that decay more quickly (slowly) than the reference (CO2) have GWPs which decrease (increase) with

increasing time horizon. As shown in Equations (8-4) and

(8-5), the most common definition of GWPs applies to

pulsed emissions. However, indices have also been developed to evaluated the effect of sustained emissions

(Berntsen et al., 2005; Johnson and Derwent, 1996; Shine

et al., 2005a).

The GWP index has three major advantages over

other indices used to measure the contribution of halocarbons to global warming: transparency, simplicity, and

widespread acceptance. Disadvantages of the GWP index

include: (1) GWPs compare contributions with radiative

forcing and not with the often more relevant factors of

temperature change or economic damage; (2) impacts at

different times in the future are given equal weight (for

times between the time of the pulse and the time of the

pulse plus the time horizon) (Fuglestvedt et al., 2000,

2003; Manne and Richels, 2001; O’Neill, 2000; Shine et

al., 2005a; Smith and Wigley, 2000a, b; Wigley, 1998);

and (3) GWPs are dependent on assumptions regarding

other gas concentrations due to spectral overlaps of absorption bands (e.g., Hurley et al., 2005)). Various alternatives

have been presented to overcome some of these limitations, but discussions of these are beyond the scope of this

chapter.

Direct GWPs are tabulated in Table 8-2. With the

recent publication of the IPCC/TEAP (2005) Special

Report and the soon to be released IPCC Fourth Assessment Report, we will limit additional discussion of

GWPs to updates since the last Ozone Assessment. There

are four reasons that updates have been made: (1) a

change in FCO2 due to an increase in the CO2 atmospheric

mixing ratio from 370 parts per million (ppm) to 378

ppm; (2) a new CO2 response function; (3) updates to two

atmospheric lifetimes; and (4) new radiative forcing

recommendations.

The change in CO2 mixing ratio reflects the continuing increase in the atmospheric concentration of this

radiatively important gas. This increased atmospheric

abundance of CO2 results in a lower radiative efficiency

due to the CO2 bands becoming slightly more saturated.

The formula used to calculate the forcing has also been

changed to 

(8-6)

where c0 is 378 ppm and ∆c is the added pulse. This formula is adopted because it is more consistent with current

estimates of the radiative forcing associated with a doubling of CO2 (IPCC, 2001). The mixing ratio used in

Equation (8-6) in conjunction with a pulse of 1 ppm leads

to a CO2 radiative efficiency of 0.0141 W m-2 ppm-1 compared with a value of 0.0153 W m-2 ppm-1 from the previous Assessment. A new pulse response function for CO2

was calculated with the Bern25CC model (Joos et al.,

2001; Plattner et al., 2001) for a constant CO2 background

mixing ratio of 378 ppm and a pulse size of 40 GtC. The

combination of this slightly changed response function

and the smaller radiative forcing leads to CO2 AGWPs of

0.192, 0.676, and 2.223 W m-2 ppm-1 yr for time horizons

of 20, 100, and 500 years, respectively (F. Joos, personal

communication). These AGWPs are smaller than those

of IPCC (2001) by 7.2%, 2.9%, and 0.8% for these same

respective time horizons. GWPs are calculated relative to

CO2 so these lower CO2 AGWPs result in increased GWP

values by these same percentages in the absence of other

changes.

Lifetimes have been updated for trifluoromethylsulfurpentafluoride (SF5CF3) and methyl chloride

(CH3Cl). A discussion of these updates can be found in

Chapter 1. We also recommend updated radiative efficiencies of six compounds compared with WMO (2003):

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.9

AGWP t F x t dt x x

t

'

'

( ) =  ( )   ∫  0

GWP t

F t dt

F R t dt

x

x x

t

CO

t '

exp

'

' ( ) =

−

( )

( )

∫

∫

τ 0

0

2

5 35 0

0

. ln c c

c

 + ∆










HALOCARBON SCENARIOS, ODPs, AND GWPs

8.10

Table 8-2. Direct Global Warming Potentials for selected gases. 

Industrial Designation Chemical Formula Radiative Lifetime Global Warming Potential for

or Common Name Efficiency 1 (years) Given Time Horizon

(W m-2 ppbv-1) 20 years 100 years 500 years

Carbon dioxide CO2 1.41 × 10-5 2 111

Nitrous oxide N2O 3.03 × 10-3 114 3 289 298 153

Chlorofluorocarbons

CFC-11 CCl3F 0.25 45 6,730 4,750 1,620

CFC-12 CCl2F2 0.32 100 10,990 10,890 5,200

CFC-13 CClF3 0.25 640 10,800 14,420 16,430

CFC-113 CCl2FCClF2 0.30 85 6,540 6,130 2,690

CFC-114 CClF2CClF2 0.31 300 8,040 10,040 8,730

CFC-115 CClF2CF3 0.18 1700 5,310 7,370 9,990

Hydrochlorofluorocarbons

HCFC-21 CHCl2F 0.14 1.7 530 151 46

HCFC-22 CHClF2 0.20 12.0 5,160 1,810 549

HCFC-123 CHCl2CF3 0.14 1.3 273 77 24

HCFC-124 CHClFCF3 0.22 5.8 2,070 609 185

HCFC-141b CH3CCl2F 0.14 9.3 2,250 725 220

HCFC-142b CH3CClF2 0.20 17.9 5,490 2,310 705

HCFC-225ca CHCl2CF2CF3 0.20 1.9 429 122 37

HCFC-225cb CHClFCF2CClF2 0.32 5.8 2,030 595 181

Hydrofluorocarbons

HFC-23 CHF3 0.19 4 270 11,990 14,760 12,230

HFC-32 CH2F2 0.11 4 4.9 2,330 675 205

HFC-41 CH3F 0.02 2.4 323 92 28

HFC-125 CHF2CF3 0.23 29 6,340 3,500 1,100

HFC-134 CHF2CHF2 0.18 9.6 3,400 1,100 335

HFC-134a CH2FCF3 0.16 4 14.0 3,830 1,430 435

HFC-143 CH2FCHF2 0.13 3.5 1,240 353 107

HFC-143a CH3CF3 0.13 52 5,890 4,470 1,590

HFC-152 CH2FCH2F 0.09 0.60 187 53 16

HFC-152a CH3CHF2 0.09 1.4 437 124 38

HFC-227ea CF3CHFCF3 0.26 4 34.2 5,310 3,220 1,040

HFC-236cb CH2FCF2CF3 0.23 13.6 3,630 1,340 407

HFC-236ea CHF2CHFCF3 0.30 10.7 4,090 1,370 418

HFC-236fa CF3CH2CF3 0.28 240 8,100 9,810 7,660

HFC-245ca CH2FCF2CHF2 0.23 6.2 2,340 693 211

HFC-245fa CHF2CH2CF3 0.28 7.6 3,380 1,030 314

HFC-365mfc CH3CF2CH2CF3 0.21 8.6 2,520 794 241

HFC-43-10mee CF3CHFCHFCF2CF3 0.40 15.9 4,140 1,640 499

Chlorocarbons

Methyl chloroform CH3CCl3 0.06 5.0 506 146 45

Carbon tetrachloride CCl4 0.13 26 2,700 1,400 435

Methyl chloride CH3Cl 0.01 1.0 45 13 4
HALOCARBON SCENARIOS, ODPs, AND GWPs

8.11

Table 8-2, continued.

Industrial Designation Chemical Formula Radiative Lifetime Global Warming Potential for

or Common Name Efficiency 1 (years) Given Time Horizon

(W m-2 ppbv-1) 20 years 100 years 500 years

Bromocarbons

Methyl bromide CH3Br 0.01 0.7 17 5 1

Halon-1201 CHBrF2 0.14 5.8 1,380 404 123

Halon-1211 CBrClF2 0.30 16 4,750 1,890 574

Halon-1301 CBrF3 0.32 65 8,480 7,140 2,760

Halon-2402 CBrF2CBrF2 0.33 20 3,680 1,640 503

Fully fluorinated species

Sulfur hexafluoride SF6 0.52 3200 16,260 22,810 32,600

Trifluoromethylsulfur- 650- 13,120- 17,540- 20,060-

pentafluoride SF5CF3 0.57 950 13,180 17,960 22,360

Perfluoromethane CF4 0.10 4 50000 5,210 7,390 11,190

Perfluoroethane C2F6 0.26 10000 8,620 12,200 18,180

Perfluoropropane C3F8 0.26 2600 6,310 8,830 12,450

Perfluorobutane C4F10 0.33 2600 6,330 8,850 12,480

Perfluorocyclobutane c-C4F8 0.32 3200 7,310 10,250 14,660

Perfluoropentane C5F12 0.41 4100 6,510 9,150 13,260

Perfluorohexane C6F14 0.49 3200 6,620 9,290 13,280

Perfluorodecalin C10F18 0.56 5 1000 5,500 7,510 9,440

Halogenated alcohols and ethers

HFE-125 CHF2OCF3 0.44 136 13,790 14,910 8,490

HFE-134 CHF2OCHF2 0.45 26 12,190 6,320 1,960

HFE-143a CH3OCF3 0.27 4.3 2,630 756 230

HCFE-235da2 CHF2OCHClCF3 0.38 2.6 1,230 349 106

HFE-245fa2 CHF2OCH2CF3 0.31 4.9 2,280 659 200

HFE-254cb2 CH3OCF2CHF2 0.28 2.6 1,260 359 109

HFE-7100 (HFE-44-9) CH3OC4F9 0.31 5.0 1,390 404 123

HFE-7200 (HFE-56-9) C2H5OC4F9 0.30 0.77 200 57 17

HFE-245cb2 CH3OCF2CF3 0.32 5.1 2,440 708 215

HFE-347mcc3 CH3OCF2CF2CF3 0.34 5.2 1,980 575 175

HFE-356pcc3 CH3OCF2CF2CHF2 0.33 0.93 386 110 33

HFE-374pc2 CH3CH2OCF2CHF2 0.25 5.0 1,930 557 169

CH3OCF(CF3)2 0.31 3.4 1,200 343 104

HFE-43-10pccc124 a CHF2OCF2OC2F4OCHF2 1.37 6.3 6,320 1,870 569

(CF3)2CHOH 0.28 2.0 764 217 66

HFE-236ca12 CHF2OCF2OCHF2 0.66 12.1 8,040 2,820 859

HFE-338pcc13 CHF2OCF2CF2OCHF2 0.87 6.2 5,070 1,500 456

Species whose lifetimes have a high uncertainty

Nitrogen trifluoride NF3 0.21 4 740 13,370 18,000 21,270

Perfluorocyclopropane c-C3F6 0.42 >1000 >12,700 >17,340 >21,800

HFE-227ea CF3CHFOCF3 0.40 11 4,540 1,540 468

HFE-236ea2 CHF2OCHFCF3 0.44 5.8 3,370 989 301

HFE-236fa CF3CH2OCF3 0.34 3.7 1,710 487 148

HFE-245fa1 CHF2CH2OCF3 0.30 2.2 1,010 286 87
the hydrofluorocarbons HFC-134a (CFH2CF3), HFC-23

(CHF3), HFC-32 (CH2F2), HFC-227ea (CF3CHFCF3), and

CF4 and NF3. The radiative efficiency values used in the

previous Assessment, the currently recommended values,

and the values presented in the pertinent references are

presented in Table 8-3.

Since the last Assessment (WMO, 2003), the

radiative efficiency of HFC-134a has been studied by

Forster et al. (2005). Various laboratory cross sections

and radiative transfer models were used to assess the

primary sensitivities of the radiative efficiency. The

integrated absorption cross sections from six independent laboratory studies of the bands most significant

to the radiative efficiency calculation differed by less

than 5%. While the models showed some difference in

the cloudy-sky forcing for a single atmospheric profile,

the best estimates of the radiative efficiency converged

to a value of 0.16 W m-2 ppb-1. This value is adopted; it

is close to that used in the previous Assessment (0.15 W

m-2 ppb-1) (WMO, 2003).

The radiative efficiency of CF4 has been studied by

Hurley et al. (2005) since the previous ozone Assessment.

There is good agreement (within 4%) in the absorption

cross sections measured in two independent laboratory

studies reported by Hurley et al. (2005) and in the study

of Nemtchinov and Varanasi (2003). The radiative efficiency reported by Hurley et al. (2005), 0.102 W m-2 ppb-1,

is 15% greater than calculated by Myhre et al. (1998).

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.12

Table 8-2, continued.

Industrial Designation Chemical Formula Radiative Lifetime Global Warming Potential for

or Common Name Efficiency 1 (years) Given Time Horizon

(W m-2 ppbv-1) 20 years 100 years 500 years

HFE-329mcc2 CHF2CF2OCF2CF3 0.49 6.8 3,060 919 279

HFE-338mcf2 CF3CH2OCF2CF3 0.43 4.3 1,920 552 168

HFE-347mcf2 CHF2CH2OCF2CF3 0.41 2.8 1,310 374 114

HFE-356mec3 CH3OCF2CHFCF3 0.30 0.94 355 101 31

HFE-356pcf2 CHF2CH2OCF2CHF2 0.37 2.0 931 265 80

HFE-356pcf3 CHF2OCH2CF2CHF2 0.39 3.6 1,760 502 153

CHF2OCH(CF3)2 0.41 3.1 1,330 379 115

-(CF2)4CH(OH)- 0.30 0.85 254 72 22

Note: Values are calculated for a CO2 mixing ratio of 378 ppm, compared with 370 ppm in IPCC/TEAP (2005) and WMO (2003); this leads to slightly

smaller CO2 AGWPs. a Referred to as H-Galden 1040x in previous Assessments. 1 All values from IPCC/TEAP (2005) unless otherwise noted. 2 See Section 8.2.3. 3 This value is an adjustment time that includes feedbacks of emissions on the lifetime. 4 See Table 8-3. 5 From Shine et al. (2005b).

Table 8-3. Radiative efficiency estimates for six compounds whose recommended values have changed

since the previous Assessment. All values are in units of W m-2 ppb-1.

Species WMO Hurley Forster Jain Sihra Gohar This 

(2003) et al. et al. et al. et al. et al. Assessment

(2005) (2005) (2000) (2001) (2004)

HFC-134a 0.15 0.16 0.20 0.159 0.155-0.166 0.16

CF4 0.08 0.102 0.089 1 0.116 0.10

HFC-23 0.16 0.248 0.171 0.181-0.193 0.19

HFC-32 0.09 0.155 0.105 0.110-0.111 0.11

HFC-227ea 0.30 0.322 0.256 0.243-0.271 0.26

NF3 0.13 0.21 2

1 For a constant vertical mixing ratio profile. 2 From Robson et al. (2006).
About a 10% increase is explained by the larger cross section and the remaining 5% is believed to be due to differing radiative transfer codes employed.

Since the last ozone Assessment, radiative efficiencies of HFC-23, HFC-32, and HFC-227ea have been

studied by Gohar et al. (2004), using two independent

sets of radiation codes. There is a greater than 20% difference in the calculated radiative efficiencies of these

gases reported by Sihra et al. (2001) and Jain et al.

(2000), both of which were considered in the previous

Assessment. At the time of the previous Assessment,

the reason for the differences was not clear and averages

of the two datasets were used. While the reason for the

differences between the Jain et al. and Sihra et al. values

are still unknown, the results of Gohar et al. (2004) support the values given by Sihra et al. (2001), leading to

an update of the recommendations given in Table 8-3.

The calculations by Gohar et al. (2004) also illustrate

the sensitivity of the radiative efficiency calculations to

the particular model choice.

The radiative efficiency of NF3 has been re-evaluated recently by Robson et al. (2006). The radiative

efficiency used in previous Assessments (IPCC, 2001;

WMO, 1999, 2003) was calculated from the absorption

cross section data of Molina et al. (1995) by K.P. Shine

using the simple radiative method given in Pinnock et

al. (1995), as no radiative efficiency was provided in the

Molina et al. (1995) work. The Robson et al. (2006)

study suggests that the more intense infrared (IR) features reported by Molina et al. (1995) were saturated,

causing the inferred radiative efficiency to be too small.

Molina et al. (1995) did not report the precise conditions

used to derive their absorption cross section values,

making an unambiguous evaluation of the importance

of saturation impossible. We adopt the radiative efficiency of 0.21 W m-2 ppb-1 from the more comprehensive study by Robson et al. (2006).

The 2σ uncertainty associated with the direct

GWPs shown is estimated to be ±35%. This value has

been adopted from previous ozone and climate Assessments (IPCC, 2001; WMO, 2003; IPCC/TEAP, 2005) and

is primarily due to uncertainties in the radiative efficiencies and lifetimes of the halocarbons and to uncertainties

in our understanding of the carbon cycle (IPCC, 2001).

However, because the uncertainties in the carbon cycle are

thought to be an important part of this uncertainty, the error

in the relative GWP values among halocarbons should be

less than 35% (IPCC, 2001).

The indirect radiative forcing and indirect GWPs

of a species quantify the radiative effects from changes in

abundances of other greenhouse gases resulting from the

addition of the species considered. Because the indirect

forcing and GWPs due to stratospheric ozone loss depend

on the future evolution of stratospheric ozone and thus on

the specific ODS emission scenario, these are presented in

Section 8.5 of this chapter, after the halocarbon scenarios

are discussed.

8.2.4 Degradation Products and Their

Implications for ODPs and GWPs

Degradation products of CFCs, HCFCs, and HFCs

have been discussed in IPCC/TEAP (2005). The main

conclusion was that the intermediate degradation products

of most long-lived CFCs, HCFCs, and HFCs have shorter

lifetimes than the source gases, and therefore have lower

atmospheric concentrations and smaller radiative forcings.

Intermediate products and final products are removed from

the atmosphere via deposition and washout processes and

may accumulate in oceans, lakes, and other aquatic reservoirs. Trifluoroacetic acid is a persistent degradation

product of some HCFCs and HFCs and is removed from

the atmosphere mainly by wet deposition. Its sources (natural and anthropogenic), sinks, and potential environmental effects have been reviewed by Tang et al. (1998),

Solomon et al. (2003), and IPCC/TEAP (2005). The available environmental risk assessment and monitoring data

indicate that the source of trifluoroacetic acid from the

degradation of HCFCs and HFCs will not result in environmental concentrations capable of significant ecosystem damage.

8.3 FUTURE HALOCARBON SOURCE GAS

CONCENTRATIONS

8.3.1 Introduction

Projections of future atmospheric halocarbon

mixing ratios require knowledge of future emissions and

atmospheric/oceanic loss rates in addition to current

atmospheric abundances. In this section, we use estimates

of these quantities in a simple box model to calculate

average future surface mixing ratios, which are assumed

to be related to mean atmospheric mixing ratios by a fixed

factor. These calculated mixing ratios are used to generate

equivalent effective stratospheric chlorine (EESC) estimates, which are used to evaluate the effects of different

halocarbon production/emission scenarios and hypothetical test cases. The lifetimes presented in Table 8-1 (see

also Chapter 1) are estimated from the loss rates of the

ODSs and are used in the box model. Any future changes

in the hydroxyl radical (OH) amount (IPCC, 2001) and/or

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.13
distribution, and any changes in circulation (see, e.g.,

Chapters 5 and 6) that might affect lifetimes, are neglected.

To include more complicated atmospheric interactions

such as these, more sophisticated two-dimensional (2-D)

and three-dimensional (3-D) atmospheric models should

be used.

The box model approach used here has also been

used in previous Assessments (WMO, 1995; 1999; 2003).

Global mean mixing ratios are calculated using the

equation

(8-7)

where ρ is the global mean mixing ratio (in ppt), τ is the

total atmospheric lifetime, E is the emission rate (in

kg/year), and F is the factor that relates the mass emitted

to the global mean mixing ratio, given by

(8-8)

where F is in units of ppt/kg, and M is the molecular

weight (in kg/mole). The i subscripts refer to the species

considered in the calculation. As in the previous ozone

Assessment, the calculated global mean mixing ratios are

multiplied by a factor of 1.07 to represent surface mixing

ratios. This factor is meant to account for the general

decrease of the halocarbon mixing ratios with altitude

above the tropopause. Using a constant factor such as

this neglects the dependence of this factor on the particular species and neglects any change in this factor that

could be caused by changes in circulation or by the variability of the surface emission (and the resulting variability in the atmospheric vertical distribution).

Accurate projections of emissions require an understanding of the amount of halocarbons in equipment and

product banks, the rates of release from these banks, the

quantity of equipment and products using ODSs that will

be put into service, emissive uses, and the future production

of ODSs. Banks here are defined as the quantity of ODSs

produced but not yet emitted to the atmosphere. Due to the

importance of policy decisions, energy cost, technological

advancement, and economic growth rates in estimating

future production and banks, the uncertainty level in future

emissions remains high. However, with each passing year,

as more years of emissions can be calculated from atmospheric abundance observations and fewer years remain for

the legal consumption of ODSs under the Montreal Protocol, future projections should become more constrained.

Scenarios and test cases have been developed to

describe the possible range of future atmospheric abundances of ODSs. These cases are generated using the current understanding of global production, emission, and

banks of the most widely used halocarbons. Scenario A1

is meant to represent the current baseline or “best guess”

scenario, analogous to the Ab scenario from the previous

Assessment. An estimated “maximum” scenario, developed for previous Assessments, is not generated for this

Assessment because developed and developing countries

have produced less CFCs in 2000-2004 than allowed under

the Montreal Protocol and the HCFC production in developing countries is not controlled before 2016. Hence, any

attempt to develop such a scenario would be largely speculative on our part. Indeed, it is this uncertainty in future

HCFC emission that represents a major reason for the

difference that will be discussed later between the current A1 scenario and the comparable Ab scenario of the

previous Assessment.

Consistent with past Assessments, our approach to

relating annual production, emission, and banks sizes

places most confidence on the emission values calculated

from atmospheric observations and global lifetimes. Accurate estimates of yearly averaged ODS emissions from

atmospheric observations are possible when global lifetimes are long and accurately known, and accurate global

mixing ratio observations of ODSs are available. This

approach, when used to estimate current bank sizes from

historic production and emission data, is sometimes

referred to as a top-down approach. In past Assessments,

any inconsistency between mixing ratio observations and

emission estimates based on the best knowledge of ODS

production, sales, and application-specific release functions was eliminated by adjusting the bank size so the

emissions would be consistent with the observations. The

bank that remained after the sum of adjustments for all

years was used in the future projections. Because the bank

is an accumulating difference often between two large

numbers, this method has an uncertainty that is difficult to

quantify and can lead to unreasonable bank sizes; indeed

the estimate of no bank in 2002 for CFC-12 in the previous

ozone Assessment is such an example (IPCC/TEAP, 2005).

One way an unreasonable bank could be attained is in a

case in which annual production numbers are accurately

known, but the atmospheric lifetime assumed is incorrect.

In such a case, a lifetime that is too small (large) will result

in an annual release from the bank that must be too large

(small) for calculated mixing ratios to agree with observations. Such a situation would occur year after year and

could result in a potentially significant error in the bank

estimate today, depending on the compound (Daniel et al.,

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.14

d

dt

FE i

i i

i

i

ρ ρ

τ = −

F

M i

i

= × − 5 68 10 9 .
2006). But without an independent estimate of bank size,

such an error would be difficult to identify.

A different approach for determining bank sizes is

taken in this Assessment because of the new independent

estimates of bank sizes for several ODSs for the years 2002

and 2015 (Clodic and Palandre, 2004; IPCC/TEAP, 2005).

These estimates are independent of atmospheric abundance observations and have been determined from the

number of units of equipment that use a particular ODS

and the amount of ODS in each unit. This is commonly

referred to as a bottom-up approach. An extensive explanation of this methodology can be found in IPCC/TEAP

(2005). Future emissions until 2015 estimated in this

Assessment are calculated by beginning with the 2002

bottom-up bank from IPCC/TEAP (2005), and then by

adding annual production and applying a constant annual

bank emission factor needed to attain the bottom-up bank

estimated in 2015 by IPCC/TEAP (2005). After 2015,

emissions are calculated using the same constant bank

emission factors as immediately before 2015. Thus, when

bottom-up bank size estimates for 2002 and 2015 are available for a particular ODS, the scenarios in this chapter are

consistent with them in most cases.

Bottom-up estimates are also not without problems; indeed bottom-up 2002 emission estimates for CFC11, HCFC-141b, and HCFC-142b (IPCC/TEAP, 2005),

which are dominated by emissions from foams, are smaller

by more than a factor of two when compared with what is

needed to be consistent with mixing ratio observations.

Nevertheless, it is felt that these bottom-up bank estimates

represent an important new constraint to current bank sizes

that warrants a large role in the future projections in this

chapter. In Sections 8.3.2.1 and 8.3.4, more discussion of

the uncertainties involved with bottom-up and top-down

emission estimates and their potential effects on the scenarios is presented.

8.3.2 Baseline Scenario (A1)

In this chapter, estimates of future banks and emissions have been calculated for most ODSs using annual

production reported to the United Nations Environment

Programme (UNEP, 2005), emissions estimated from

atmospheric observations, and bank size estimates based

on the bottom-up calculations of IPCC/TEAP (2005). The

specific information used for each ODS considered is

described in Table 8-4. Due to the uncertainty in the

importance of and likely future trends in very short-lived

(τ < ~0.5 years) organic bromine and chlorine source

gases, they are not considered in any of these scenarios.

However, the bromocarbons considered in this chapter

cannot explain the entire stratospheric inorganic bromine

abundance alone, so the very short-lived gases may prove

to be important. Detailed information about these compounds can be found in Chapter 2.

It should be noted that the assumptions made in calculating the baseline scenario are critical to the interpretation of the reduced production and emission scenarios

based on it. For example, if the production and/or bank

sizes and/or future production rates are underestimated in

the baseline scenario, the “no emission/production” results

presented in the scenarios would be underestimates of the

potential reduction benefit.

8.3.2.1 EMISSIONS

In Figure 8-1, the emission estimates from the 1-

box model for the baseline scenario, A1, are compared

with those of the Ab “best guess” scenario from the previous Assessment, with the IPCC/TEAP (2005) (businessas-usual (BAU) scenario) emissions, with the emissions

of Clodic and Palandre (2004), and with the emissions calculated using the 12-box model discussed in Chapter 1.

The emissions calculated with the 1-box model are calculated directly from mixing ratio observations (those shaded

in Table 8-5) as are the emissions calculated with the 12-

box model. The 12-box model, by having some vertical

resolution, includes some changes in atmospheric lifetimes

due to variations in atmospheric abundance distributions

caused by increasing or decreasing trends; hence, the 12-

box model can presumably better determine the relationship between surface mixing ratios and global averages

and can better estimate changes in global lifetimes that

arise from different atmospheric distributions. The comparisons between the A1 estimated emissions (1-box

model) and the 12-box emissions are good most times, but

do show some systematic differences. Throughout the

measurement periods (shaded regions in the figure), the

12-box model consistently estimates more emission than

the 1-box model, with the cumulative differences ranging

from a 1.2% higher emission for HCFC-141b through the

measurement period as calculated by the 12-box model, to

a 6.3% higher emission for CFC-12. There are also noticeable differences between A1 and Ab emissions, many of

which are due to the additional information acquired from

continued measurements over the past 4 years. Important

differences between A1 and Ab include the larger future

emissions of CFC-11 and CFC-12, due to their larger estimated banks in A1. The CFC-11 increase in emission

past 2010 is also due to a decrease of the bank release

rate from plastic foams. The CFC-11 and CFC-12 bank

size differences as calculated by the bottom-up method

(IPCC/TEAP, 2005) and the top-down method of WMO

(2003) have been shown by Daniel et al. (2006) to be large

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.15
HALOCARBON SCENARIOS, ODPs, AND GWPs

8.16
8.17

HALOCARBON SCENARIOS, ODPs, AND GWPs

Table 8-4. Assumptions made in obtaining production and emission estimates for the baseline A1 scenario.

General Approach for All Species

Production: For the years when production is reported to UNEP, reported values (or best estimates of production values

for cases in which reporting is incomplete or reporting is made by classes of compounds) are used. Before

this, WMO (2003) production values are generally used. In the future, annual figures are determined from

the lesser of the Montreal Protocol limitations and the most recent annual estimates.

Emission: For the years when abundance observations are available, emissions are calculated using the box model

described in Section 8.3.1 with the lifetimes of Table 8-1. Emissions before this are usually consistent with

WMO (2003) but are also forced to yield mixing ratios that meld smoothly into the measurement record.

Future emissions are determined in order to yield banks consistent with IPCC/TEAP (2005).

Bank: The bank assumed to be in place at the start of the measurement record is set at such a value that the IPCC/TEAP

(2005) bank for 2002 is attained; the future annual fractional bank release is adjusted so that the IPCC/TEAP

(2005) bank for 2015 is attained.

Approach for Specific Species

Species Description

CFC-11 Production:

1950-1985: WMO (2003)

1986, 1989-2004: UNEP * 

1987-1988: McCulloch et al. (2001)

2005-2006: Fixed at 2004 levels for Article 5(1) countries, with no other production

2007-2009: Protocol limits on Article 5(1) countries, with no other production

2010-on: No production

Emission:

1980-2004: Emissions calculated from observations

2005-on: Emission is a constant fraction of the bank, determined to give a bank in 2015 consistent with

IPCC/TEAP (2005)

CFC-12 Production:

Same as CFC-11 except:

1987-1988: Interpolated from 1986 and 1989 values

2005-2006: Protocol limits for Article 5(1) countries with no other production

Emission:

Same formalism as for CFC-11

Figure 8-1. Comparison of scenario A1 emissions (black) with those of scenario Ab of the previous

Assessment (green), the emissions of Clodic and Palandre (2004) (red), the IPCC/TEAP (2005) emissions

(crosses), and emissions calculated from atmospheric abundance observations using a 12-box model (small

diamonds), all in units of kilotons per year. Corrected data are used for emissions of halon-1211 according to

TEAP (2005) and hence differ from the data reported in IPCC/TEAP (2005). Also shown for halon-1301 are

the emissions estimates from the 2002 report of the Halons Technical Options Committee (HTOC) (UNEP,

2003) and, for halon-1211, updates to that report (blue curves). Shaded regions indicate years for which A1

emissions are determined with a 1-box model (see text) so that calculated mixing ratios are exactly equal to

the observations. At times, the emissions from the two Assessments are identical, causing the green curve

to obscure the black one.
HALOCARBON SCENARIOS, ODPs, AND GWPs

8.18

CFC-113 Production:

Same formalism as for CFC-11

Emission:

1986-2004: Emissions calculated from observations

CFC-114 Production:

Same formalism as for CFC-11 except:

1987-1988: WMO (2003)

2005-2100: No production

Emission:

1951-1959: WMO (2003)

1960-2100: Assume an annual fractional bank release of 0.4, chosen because it leads to bank results relatively close to those estimated by AFEAS (Alternative Fluorocarbons Environmental Acceptability

Study) in the 1980s

CFC-115 Production:

Same formalism as for CFC-11 except:

1987-1988: WMO (2003)

2005-2100: No production

Emission:

1950-2100: Assume an annual fractional bank release of 0.25, chosen because it leads to bank results

relatively close to those estimated by AFEAS in the 1980s

CCl4 Production:

Not considered due to gaps in our understanding of where much of the global emission originates

Emission:

1951-1979: WMO (2003) (except the mixing ratio at the beginning of 1951 is assumed to be 37.0 rather

than 35.0 (WMO, 2003) to achieve slightly better continuity with the measurements beginning in

1980

1980-2004: Emissions calculated from observations

2005-2015: Linear decrease from 65 ktons in 2005 (compared with 67.4 ktons calculated for 2004) to 0

in 2015

2015-2100: No emission

CH3CCl3 Production:

Not considered

Emission:

1950-1978: WMO (2003)

1979-2004: Emissions calculated from observations

2005-2009: Assume 2004 emissions because this value is lower than the limit in the Protocol (30%

reduction relative to the 1998-2000 emission average; although the Protocol limits production and

consumption and not emission, it is assumed that there is a negligibly small bank for this compound and that the limitations may be applied to emission)

2010-2014: Emission is 30% of the 1998-2000 average

2015-2100: No emission

HCFC-22 Production:

1950-1988: WMO (2003)

1989-2004: UNEP *

2005-2015: Linear interpolation of demand estimates from IPCC/TEAP (2005) for 2002 and 2015

2016-2030: Fixed at 2015 value

Table 8-4, continued.

Approach for Specific Species
HALOCARBON SCENARIOS, ODPs, AND GWPs

8.19

2030-2040: Linear interpolation from 2030 value to 0

Emission:

1950-1992: Emissions calculated to yield mixing ratios consistent with WMO (2003) but moved 1 year

earlier (the 1-year offset is to make the mixing ratios meld smoothly with the observations)

1993-2004: Emissions calculated from observations

2005-2100: Assume a bank release fraction of 0.13 (average of 2002-2004)

Bank:

1993 WMO (2003) bank is reduced by 193 ktons from 1,036 ktons to obtain the IPCC/TEAP (2005)

2002 bank estimate; the bank in 2015 is 2,652 ktons compared with 1,879 ktons in IPCC/TEAP

(2005); a much larger release fraction of 0.19 and increase in emissions in the near future would be

needed to be consistent with the lower bank

HCFC-141b Production:

1989-2004: UNEP *

2005-2015: Linear interpolation of demand estimates from IPCC/TEAP (2005) for 2002 and 2015

2016-2030: Fixed at 2015 value

2030-2040: Linear interpolation from 2030 value to 0

Emission:

1993-2004: Emissions calculated from observations

2005-2100: Assume a bank release fraction of 0.05

Bank:

1993 bank from WMO (2003) is not increased in spite of an apparent need for a larger 1993 bank to

attain the 2002 IPCC/TEAP (2005) bank; the increase is not applied because the increase needed

would lead to a 1993 bank too large to be consistent with estimated global production before 1993;

hence the 2002 bank is 674 ktons, while the IPCC/TEAP (2005) estimate is 836 ktons; the 2015

bank is in agreement with the IPCC/TEAP (2005) bank

HCFC-142b Production:

1992-2004: UNEP * (except for 1993)

1993: AFEAS data used because of unexplained drop in UNEP data for 1993 

2005-2015: Linear interpolation of demand estimates from IPCC/TEAP (2005) for 2002 and 2015

2016-2030: Fixed at 2015 value

2030-2040: Linear interpolation from 2030 value to 0

Emission:

1971-1992: WMO (2003) multiplied by 1.225 to attain a mixing ratio consistent with the start of the

measurements in 1993

1993-2004: Emissions calculated from observations

2005-2100: Assume a bank release fraction of 0.08

Bank:

1993 WMO (2003) bank is increased by 18 ktons to 103 ktons so the bank in 2002 is equal to the

IPCC/TEAP (2005) 2002 bank; the bank release fraction of 0.08, which is consistent with the values

estimated from 2002-2004 (0.085-0.089), leads to a bank in 2015 of 157 ktons, considerably lower

than the 331 ktons estimated by IPCC/TEAP (2005)

Halon-1211 Production:

1989-2004: Corrected values from the Halons Technical Options Committee (HTOC) 2002 report

(UNEP, 2003)

2005-2015: Linear interpolation of demand estimates from IPCC/TEAP (2005) for 2002 and 2015

2016-2030: Fixed at 2015 value

2030-2040: Linear interpolation from 2030 value to 0

Table 8-4, continued.

Approach for Specific Species
HALOCARBON SCENARIOS, ODPs, AND GWPs

8.20

Emission:

1950-1992: WMO (2003) multiplied by 1.02 to attain a mixing ratio in 1993 consistent with measurements

1993-2004: Emissions calculated from observations

2005-2100: Assume a bank release fraction of 0.07

Bank:

2002 bank of 122 ktons is only 3 ktons below the IPCC/TEAP (2005) bank; the bank release fraction of

0.07, which is consistent with the release fraction from 1993-2004, leads to a bank in 2015 of 52

ktons compared with 31 ktons from IPCC/TEAP (2005); the 2002 bank of 125 ktons is larger than

the 72 ktons in WMO (2003)

Note that corrected data are used for banks and emissions of halon-1211 (TEAP, 2005) and hence differ from

the data reported in IPCC/TEAP (2005). The 2002 emission data are changed from 17 to 8 ktons per year.

Halon-1301 Production:

1986-2009: HTOC 2002 report (UNEP, 2003)

2010-2100: No production

Emission:

1950-1995: WMO (2003)

1996-2100: Assume a bank release fraction of 0.05 (leads to mixing ratios between NOAA/ESRL and

AGAGE observations)

Bank:

2002 and 2015 banks agree with IPCC/TEAP (2005)

Halon-1202 Same as WMO (2003)

Halon-2402 Same as WMO (2003)

CH3Br Production:

Natural production/emission assumed to be 146 ktons

2005: Natural plus 10.7 ktons for quarantine/pre-shipment plus 14.1 ktons for critical uses plus 0.5

ktons for Article 5(1) production (same as 2004 reported to UNEP)

2006-2014: Same as 2005 but 13.0 ktons of critical uses

2015-2100: Natural plus 10.7 ktons for quarantine/pre-shipment

Emission:

1950-2004: Emissions calculated from surface observations and South Pole firn observations (using

global lifetime of 0.7 years, Table 8-1)

2005-2100: Emission equal to 0.88 times the anthropogenic production plus natural emissions; this

combination of anthropogenic and natural emissions, constrained by the total emissions derived

from observations of concentrations, leads to an anthropogenic fraction of the total production of

0.30 in 1992, in agreement with Montzka et al. (2003).

Bank:

No bank considered

CH3Cl Emission:

Same as WMO (2003); assumed mostly natural (i.e., no future changes due to anthropogenic activity)

1950-1995: Emissions calculated from firn observations at the South Pole

1996-2100: Emissions held constant at 1995 levels

* Estimated in cases in which reporting is not complete or reporting is made in compound classes rather than individually. The production data per

species per year are obtained from UNEP (UNEP, 2005) and are consistent with the totals per class of species as usually reported by UNEP.

Table 8-4, continued.

Approach for Specific Species
enough to have significant implications concerning the

environmental benefits of reusing or destroying CFCs contained in existing equipment and products. Increases in

the projected HCFC-22 emissions, due to larger reported

(to UNEP) historic production in developing countries in

2000-2004 than projected in the Ab scenario, and expectations of greater future use based on IPCC/TEAP (2005),

are also important.

While the historic emissions calculated with the 1-

box model for this Assessment are completely constrained

by the observations, emissions from Clodic and Palandre

(2004) and IPCC/TEAP (2005) are based exclusively on

estimates of the banks and bank release fractions and can

be substantially different from those estimated from

mixing ratio observations. As previously stated, for some

species that have significant foam applications, i.e., CFC11, HCFC-141b, and HCFC-142b, these emission differences are greater than a factor of two, with the estimated

bottom-up emission underpredicting those derived from

observed mixing ratio changes. In Table 1-7 of Chapter 1,

an uncertainty range for the emissions in 2003 is given

based on an uncertainty range in the lifetimes of the ODSs

and calculations with the 12-box model. The uncertainty

ranges of the inferred 2003 emissions are somewhat larger

than those given in IPCC/TEAP (2005), but the IPCC

ranges were due to model differences and to the variations

in the trends of the different global networks, rather than

different assumed lifetimes. Nevertheless, the top-down

emissions in scenario A1 and bottom-up emissions of

IPCC/TEAP (2005) for CFC-11, HCFC-141b, and HCFC142b cannot be reconciled taking these uncertainties in

emissions into account. These differences illustrate that

there is a need for greater understanding regarding the

possible shortcomings of the bottom-up approach and the

differences between it and past approaches. The discrepancies are large enough to cast some doubt on projected

future emissions that are based on the published bottomup results alone. Section 8.3.4 contains a more extensive

discussion of uncertainties.

8.3.2.2 MIXING RATIOS

The projected ODS mixing ratios for the baseline

scenario are mostly consistent with those of the previous

Assessment before the year 2002 because the previous

Assessment’s past mixing ratios were based on observations, and here observations are used as a direct constraint

(see Figure 8-2). The observations are taken primarily

from the Earth System Research Laboratory / Global

Monitoring Division (ESRL/GMD) (formerly Climate

Monitoring and Diagnostics Laboratory, CMDL)

(Montzka et al., 1999) and Advanced Global Atmospheric

Gases Experiment (AGAGE) networks (Prinn et al., 2005)

(and include the University of East Anglia observations

for halon-1211), and the calibration factors were calculated in a similar manner as in Table 1-15 of WMO (2003).

After correcting for calibration differences, the mean

mixing ratios of the ESRL/GMD and AGAGE networks

are used for temporally overlapping periods for the CFCs,

CH3CCl3, and for CCl4. For periods before any overlap

exists, the ratios of the measurements are used to extend

the record backward in time in a consistent manner. The

average of the December and January observations are

assumed to represent the mixing ratios at the start of the

year, with the mixing ratios for the rest of the year calculated from estimated emissions and lifetimes. Although

the model is run in 0.1-year time steps, yearly observations are used instead of monthly ones in order to avoid

the impact of seasonal variability on the inferred emissions, which are assumed to be constant throughout each

year. Calculated mixing ratios are tabulated in Table 8-5

for each of the considered halocarbons from 1955 through

2100. The yearly observations used are indicated in the

table by the shaded regions.

Time series are shown in Figure 8-2 for the current

A1 scenario and the previous Ab scenario. There are some

differences between mixing ratios of this Assessment

compared with the previous one for the period 2001-2004,

but the greatest differences are found in the future projections. The HCFC-22 projection exhibits the most striking

differences compared with the previous Assessment’s scenario Ab projection with a peak mixing ratio over 150 ppt

higher and peaking more than 20 years later. In contrast,

HCFC-141b, HCFC-142b, and halon-1301 exhibit

decreases in near-term projected mixing ratios compared

with the Ab scenario, due to the decrease in observed

growth rates from 2001 through 2004 and the expectation

of lower future emissions for these species (Figure 8-1).

Due to the increased emissions discussed in Section

8.3.2.1, future CFC-11 and CFC-12 mixing ratios also

exhibit modest, but important, increases compared with

the previous Assessment. Carbon tetrachloride decreases

more slowly in A1 due to greater emission in the coming

decade, with the increased emission based on the uncertain source of much of the global CCl4 emission (see

Chapter 1) and the slow decline of annual emissions over

the past 15 years.

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.21
HALOCARBON SCENARIOS, ODPs, AND GWPs

8.22

Table 8-5. Mixing ratios (ppt) of the ODSs considered in scenario A1. Values are for the beginning of the corresponding year. Potentially

important short-lived gases that may currently contribute 3-8 ppt of stratospheric bromine and ~50 ppt of stratospheric chlorine (see Chapter 2)

are not shown in the table.

CFC-11 CFC-113 CFC-115 CH C3 Cl3 HCFC-141b Halon-1211 Halon-1301 CH B3 r

CFC-12 CFC-114 CCl4 HCFC-22 HCFC-142b Halon-1202 Halon-2402 CH C3 l

1955 3.3 14.3 1.3 2.6 0.0 42.4 0.1 1.0 0.0 0.0 0.00 0.00 0.00 0.00 7.13 492.1

1960 9.5 29.5 1.9 3.9 0.0 52.1 1.5 2.1 0.0 0.0 0.00 0.00 0.00 0.00 7.30 510.31965 23.5 58.8 3.1 5.1 0.0 64.4 4.7 4.9 0.0 0.0 0.00 0.00 0.00 0.00 7.49 528.1

1970 52.8 114.3 5.5 6.5 0.2 75.9 16.3 12.1 0.0 0.0 0.02 0.00 0.00 0.01 7.71 539.9

1975 106.1 203.1 10.4 8.2 0.6 85.5 40.0 23.8 0.0 0.1 0.12 0.00 0.04 0.04 7.97 545.81980 163.7 294.9 20.1 10.3 1.5 93.4 82.0 40.5 0.0 0.3 0.43 0.01 0.24 0.11 8.27 548.31981 173.4 313.6 22.9 10.8 1.8 94.6 89.9 44.4 0.0 0.3 0.53 0.01 0.31 0.12 8.34 548.6

1982 182.3 329.8 26.0 11.3 2.1 95.3 94.5 48.0 0.0 0.3 0.64 0.01 0.39 0.14 8.41 548.91983 189.3 346.5 29.2 11.7 2.4 96.5 98.2 52.0 0.0 0.4 0.76 0.02 0.49 0.16 8.48 549.1

1984 198.2 363.5 32.8 12.2 2.8 98.1 103.3 56.3 0.0 0.4 0.90 0.02 0.61 0.18 8.55 549.3

1985 207.5 381.5 38.1 12.7 3.2 99.6 108.1 60.6 0.0 0.5 1.06 0.02 0.74 0.20 8.62 549.41986 216.7 398.1 41.8 13.2 3.6 100.1 111.1 65.1 0.0 0.5 1.25 0.02 0.89 0.22 8.70 549.51987 228.9 419.2 46.8 13.7 4.0 102.1 113.2 69.7 0.0 0.7 1.46 0.02 1.06 0.24 8.78 549.6

1988 240.9 437.7 54.3 14.3 4.4 102.8 119.4 75.1 0.0 0.9 1.72 0.02 1.25 0.26 8.86 549.7

1989 248.9 458.8 61.6 14.8 4.8 104.1 124.0 81.3 0.0 1.1 2.02 0.03 1.46 0.29 8.94 549.81990 256.6 473.5 67.4 15.3 5.3 104.8 124.9 87.3 0.0 1.4 2.31 0.03 1.66 0.31 9.03 549.8

1991 263.7 489.1 73.2 15.8 5.8 106.4 130.9 93.5 0.0 2.0 2.59 0.03 1.84 0.34 9.12 549.9

1992 266.2 499.7 78.7 16.2 6.3 105.1 130.9 99.9 0.0 2.9 2.84 0.03 1.98 0.36 9.21 549.91993 269.3 509.9 81.7 16.5 6.7 104.7 130.1 104.4 0.0 3.9 3.00 0.03 2.06 0.38 9.30 549.9

1994 269.7 516.3 83.0 16.8 7.2 104.0 121.9 108.8 1.1 5.0 3.18 0.04 2.17 0.40 9.40 549.91995 269.3 522.8 84.0 16.9 7.6 103.1 109.9 112.1 2.7 6.3 3.31 0.04 2.33 0.42 9.50 550.0

1996 268.6 527.3 83.7 17.0 8.0 102.5 98.3 119.6 4.5 7.4 3.47 0.04 2.45 0.43 9.10 550.0

1997 267.8 531.0 83.6 17.1 8.2 101.4 84.2 124.3 6.6 8.5 3.64 0.05 2.52 0.43 9.06 550.01998 266.1 533.8 83.3 17.1 8.5 100.4 71.5 128.4 8.2 9.4 3.77 0.05 2.59 0.42 9.25 550.0

1999 264.3 536.1 82.8 17.1 8.6 99.9 59.7 135.0 10.2 10.5 3.90 0.05 2.65 0.42 9.27 550.02000 262.6 538.1 82.3 17.0 8.7 98.8 49.9 139.8 11.9 11.6 4.01 0.05 2.71 0.41 8.90 550.02001 261.2 539.7 81.9 17.0 8.8 98.0 41.6 144.1 13.5 12.5 4.08 0.04 2.76 0.41 8.48 550.0

2002 259.0 540.2 81.3 17.0 8.9 96.8 34.6 150.8 14.9 13.5 4.14 0.04 2.81 0.40 8.22 550.0
HALOCARBON SCENARIOS, ODPs, AND GWPs

8.23

Table 8-5, continued.

CFC-11 CFC-113 CFC-115 CH C3 Cl3 HCFC-141b Halon-1211 Halon-1301 CH B3 r

CFC-12 CFC-114 CCl4 HCFC-22 HCFC-142b Halon-1202 Halon-2402 CH C3 l

2003 256.9 539.9 80.6 16.9 9.0 95.8 28.9 155.1 16.2 14.1 4.20 0.03 2.85 0.39 8.17 550.0

2004 255.2 539.9 79.7 16.9 9.0 94.9 24.0 159.8 17.0 14.7 4.22 0.03 2.89 0.38 8.01 550.0

2005 253.0 538.9 78.9 16.8 9.1 94.0 20.1 164.7 17.4 15.2 4.24 0.02 2.93 0.37 7.86 550.0

2006 249.6 536.1 78.0 16.8 9.1 92.9 16.9 170.7 17.9 15.5 4.25 0.02 2.96 0.36 7.64 550.0

2007 246.3 533.3 77.2 16.7 9.1 91.7 14.3 177.1 18.5 15.8 4.24 0.02 2.99 0.35 7.55 550.0

2008 242.9 530.3 76.3 16.7 9.1 90.2 12.2 183.9 19.1 16.0 4.21 0.01 3.01 0.34 7.53 550.0

2009 239.6 527.2 75.4 16.6 9.1 88.6 10.4 191.1 19.7 16.2 4.17 0.01 3.04 0.32 7.53 550.0

2010 236.3 523.9 74.6 16.6 9.1 86.8 9.0 198.5 20.4 16.3 4.12 0.01 3.05 0.31 7.52 550.0

2015 219.8 506.0 70.3 16.3 9.1 75.0 4.4 238.7 23.2 16.4 3.73 0.00 3.10 0.26 7.52 550.0

2020 203.8 486.2 66.3 16.0 9.1 61.9 1.6 279.7 25.0 15.3 3.22 0.00 3.08 0.21 6.99 550.0

2025 188.4 465.6 62.5 15.8 9.1 51.1 0.6 314.8 26.2 13.5 2.70 0.00 3.01 0.17 6.99 550.0

2030 173.7 444.9 59.0 15.5 9.0 42.1 0.2 342.1 27.0 11.6 2.22 0.00 2.91 0.13 6.99 550.0

2035 159.7 424.5 55.6 15.3 9.0 34.8 0.1 354.8 27.2 9.8 1.79 0.00 2.79 0.10 6.99 550.0

2040 146.5 404.7 52.4 15.0 9.0 28.7 0.0 330.0 26.0 8.1 1.42 0.00 2.66 0.08 6.99 550.0

2045 134.2 385.4 49.4 14.8 9.0 23.7 0.0 269.6 23.2 6.6 1.12 0.00 2.52 0.06 6.99 550.0

2050 122.7 367.0 46.6 14.5 8.9 19.5 0.0 203.7 19.7 5.3 0.88 0.00 2.38 0.05 6.99 550.0

2055 112.0 349.3 43.9 14.3 8.9 16.1 0.0 147.2 16.3 4.2 0.68 0.00 2.24 0.04 6.99 550.0

2060 102.1 332.4 41.4 14.0 8.9 13.3 0.0 103.5 13.2 3.3 0.53 0.00 2.10 0.03 6.99 550.0

2065 92.9 316.3 39.1 13.8 8.9 11.0 0.0 71.4 10.6 2.6 0.40 0.00 1.97 0.02 6.99 550.0

2070 84.5 300.9 36.8 13.6 8.8 9.0 0.0 48.7 8.4 2.0 0.31 0.00 1.84 0.02 6.99 550.0

2075 76.7 286.3 34.7 13.4 8.8 7.5 0.0 32.9 6.6 1.6 0.23 0.00 1.71 0.01 6.99 550.0

2080 69.6 272.3 32.7 13.1 8.8 6.2 0.0 22.1 5.2 1.2 0.18 0.00 1.60 0.01 6.99 550.0

2085 63.1 259.1 30.9 12.9 8.8 5.1 0.0 14.8 4.1 0.9 0.13 0.00 1.48 0.01 6.99 550.0

2090 57.1 246.5 29.1 12.7 8.7 4.2 0.0 9.8 3.2 0.7 0.10 0.00 1.38 0.01 6.99 550.0

2095 51.7 234.4 27.4 12.5 8.7 3.5 0.0 6.5 2.5 0.6 0.08 0.00 1.28 0.01 6.99 550.0

2100 46.7 223.0 25.9 12.3 8.7 2.9 0.0 4.3 1.9 0.4 0.06 0.00 1.19 0.00 6.99 550.0

Note: Areas are shaded for compounds in years when mixing ratio values are forced to equal global estimates calculated from observations. Global means for

CH B3 r and CH C3 l before 1996 were derived from South Pole firn air observations (Butler et al., 1999) and based on surface observations of CH B3 r after 1996

(Montzka et al., 2003). All other shaded areas are from atmospheric observations from one or more measurement networks (see text). Halon-1301 values

used in this scenario fall between the global estimates made from observations of the NOAA/ESRL and AGAGE networks.
HALOCARBON SCENARIOS, ODPs, AND GWPs

8.24

CFC90 2 00 2010 2020 2030 2040

40

60

80

220

40

60

80

ppt

CFC-12

90 2 0 2010 2020 2030 204

50

550

ppt

CFC- 13

90 2 00 2010 2020 2030 2040

50

70

80

90

ppt

CFC-114

90 2 0 2010 2020 2030 204

4.5

5.0

5.5

6.0

6.5

7.0

7.5

ppt

CCl

90 2 00 2010 2020 2030 2040

20

40

60

80

20

ppt

CH CCl

90 2 0 2010 2020 2030 204

50

50

ppt

HCFC-22

90 2 00 2010 2020 2030 2040

ppt

HCFC-141b

90 2 0 2010 2020 2030 204

10

20

30

40

ppt

HCFC-142b

90 2 00 2010 2020 2030 2040

10

20

30

ppt

Halon 12

90 2 0 2010 2020 2030 204

ppt

Halon 1301

90 2 00 2010 2020 2030 2040

Year

ppt

CH Br

90 2 0 2010 2020 2030 204

Year

10

ppt

Figure 8-2. Projected

mixing ratios of selected halocarbons for

scenarios A1 (solid

black), Ab (green, previous Assessment),

E0 (dashed curve),

and P0 (dotted curve).

Note that for many of

the gases, little or no

future production is included in the A1 scenario, so the P0 curve

is obscured by the A1

curve. For CH3Br, the

E0 and P0 curves are

identical.
8.3.2.3 EQUIVALENT EFFECTIVE STRATOSPHERIC

CHLORINE

The equivalent effective stratospheric chlorine

(EESC) index (see Daniel et al., 1995; WMO, 2003) has

been used in past Assessments and is summarized in Box

8-1. It is used as a measure of the amount of chlorine and

bromine available in the stratosphere to destroy ozone.

Here, we set fCFC-11, the absolute fractional release value for

CFC-11 (Equation 1 of Box 8-1), to be 0.84 as was done in

the previous Assessment, with the ρ values assumed to be

equal to the surface mixing ratios. This factor was incorrectly stated to be 0.8 in footnote “a” of Table 1-4 in the previous Assessment. It should also be noted that the CH3Br

mixing ratio at the tropopause may be less than the surface

mixing ratio by a non-negligible amount due to its relatively

large tropospheric loss (see Chapter 2); this would lead to

an additional reduction in the EESC contribution from

CH3Br that is not considered in this chapter. Newman et al.

(2006) have extended the definition of EESC in order for it

to correspond to the actual amount of inorganic chlorine

and bromine in various stratospheric locations by applying

location-appropriate age spectra, fractional release values,

and “fCFC-11” values. We do not consider the variation of

EESC with stratospheric location, except in accounting for

the older age of air in the Antarctic polar vortex as discussed

below. Instead, we continue using EESC as an index that

is only proportional to overall stratospheric inorganic chlorine and bromine. Very short-lived species are currently not

included in the EESC calculation. In the future, as our

knowledge of the very short-lived bromocarbon source

gases improves (see Chapter 2), it may be warranted to

link EESC directly with spatially varying inorganic chlorine and bromine in order to assess the importance of

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.25

Box 8-1. Equivalent Effective Stratospheric Chlorine (EESC)

Due to the established relationship between stratospheric ozone depletion and inorganic chlorine and bromine

abundances, the temporal evolution of chlorine- and bromine-containing source gases is an important indicator of the

potential effects of anthropogenic activity on the health of stratospheric ozone. Indices have been developed to demonstrate this halogen evolution in a simple manner. They account for the greater per-atom potency of stratospheric

bromine (Br) compared with chlorine (Cl) in its ozone destructiveness with a constant factor, α (Section 8.2.2.3), and

include the varying rates at which Cl and Br will be released in the stratosphere from different source gases.

EESC provides a simple index that relates the time evolution of long-lived surface abundances of ODSs with

the ozone-destructive ability of stratospheric halogens that come from these long-lived source gases (WMO, 1995,

1999, 2003). Contributions of very short-lived chlorine- and bromine-containing sources gases and of tropospheric

inorganic halogens generally have been neglected. EESC is defined as 

(Daniel et al., 1995), where n is the number of chlorine or bromine atoms in the source gas, fi / fCFC-11 represents the

efficiency of the stratospheric halogen release relative to that of CFC-11, denoted by fCFC-11 (Section 8.2.2.2), and

ρi,entry is the tropospheric mixing ratio of source gas i when it entered the stratosphere. Traditionally, ρi,entry is calculated assuming a simple time lag Γ from the surface observations, i.e.,

where ρi(t) is the surface mixing ratio at time t. In WMO (2003), EESC was estimated assuming Γ = 3 years (typical

of the lower, midlatitude stratosphere) to obtain a value appropriate for relating to midlatitude-averaged ozone loss.

Effective equivalent chlorine (EECl) (Montzka et al., 1996) is a quantity similar to EESC, but includes no consideration of the transport lag time.

To retain the simplicity of the EESC index, several assumptions are generally made in its calculation. One such

assumption is that the stratospheric entry mixing ratio for a given time is taken to be the surface mixing ratio at that

time. This could be an overestimate of the EESC contribution for shorter-lived gases, like CH3Br, whose abundance

may be reduced before reaching the tropopause. In Montzka et al. (2003), this issue has been addressed by reducing

the surface mixing ratio by 7% in the calculation of EESC. A second assumption generally made in EESC calculations

EESC t f n f

f n ff CFC i

i

CFC

i entry i

i

CF

( ) = + −

−

11

11

ρ α ,

C

i entry

Br containing

halocarbons

Cl co − − −

∑ 11

ρ ,

ntaining

halocarbons

∑





















(1)

ρ ρ i entry i t t , ( ) = − ( ) Γ (2)

Continued, next page.
HALOCARBON SCENARIOS, ODPs, AND GWPs

8.26

is the neglect of the transport time from the surface to the tropopause. Although the particular location of the surface

emission affects the transport time to the tropopause, this is generally not considered.

Newman et al. (2006) extended the method of calculating EESC to account for the lack of a single transit time

from the surface to a stratospheric location (i.e., stratospheric air is composed of air characterized by a range, or spectrum, of transit times, or ages) and that the fractional release values depend on the age of air. In Newman et al. (2006),

ρi,entry is calculated using 

where G(t) is the age-spectrum, assumed to be an inverse Gaussian function with mean Γ and width ∆ (see Equation 9

of Waugh and Hall, 2002), and fi = fi (Γ). This reduces to the traditional EESC calculation if the mean age Γ = 3 yrs

and the width ∆ = 0, i.e., (3) reduces to (2) when ∆ = 0. Including an age spectrum does not affect the EESC evolution

when the temporal trend is constant, but acts to smooth changes over time when the trend is varying. Specifically,

when an age spectrum is included, the period near when the maximum EESC occurs is characterized by a “flatter” or

less peaked EESC time series.

The figure compares calculations of EESC, normalized by the 1980 value, for midlatitude and polar regions and

illustrates some of the important sensitivities. The dashed lines show the EESC evolution used in Chapter 8 for midlatitudes (black) and inside the polar vortex in the lower stratosphere (red) (i.e., for Γ = 3 yrs and ∆ = 0 yrs for midlatitudes,

and Γ = 6 yrs and ∆ = 0 yrs for the vortex). The solid lines show the same calculations using the fractional release

values from Newman et al. (2006) and the same mean ages as the dashed lines, but with ∆ equal to half the mean age in

each case. The shaded regions show the sensitivity of the calculations using the Newman et al. fractional release values

to a ±1-year mean age and with ∆ always equal to half the mean age (e.g., for Γ = 2 yrs and ∆ = 1 yr, to Γ = 4 yrs and ∆ =

2 yrs for the midlatitude case). Using a larger value of Γ results in a shift in curves and leads to a much later recovery

date, but does not change the EESC shape. Compared with the midlatitude recovery, the polar recovery time is delayed

by much more than the 3-year older age of air because the magnitude of the slope of the EESC curve is projected to be

much smaller in the mid-21st century than it was in 1980s. The inclusion of the age spectrum width in the polar calculation is the reason for the smoother EESC peak period but has little effect during periods when EESC is changing approximately linearly with time, and hence affects the time of return to 1980 levels only slightly.

1980 2000 2020 2040 2060 2080

0.5

1.0

1.5

2.0

EESC (normalized to 1980)

ρ ρ (3) i entry i

t

t t G t t dt , ( ) = ( )' '' ( ) −

−∞

∫

Box 8-1, continued.
these very short-lived species relative to long-lived source

gases for ozone depletion.

As in previous Assessments, the year EESC returns

to its 1980 levels is one metric used to compare different

scenarios. The year 1980 is chosen because this is the

approximate date when midlatitude and Antarctic ozone

depletion has been observed to begin. It generally has been

assumed that if all other atmospheric parameters and

processes remain constant, ozone depletion relates linearly

to EESC above a certain threshold level. An exception to

this relationship is Antarctic ozone depletion. Springtime

depletion became so great around 1990 that there was not

enough ozone left in the lower stratosphere for the column

ozone amount to continue to follow a linear relationship with

EESC. So it is assumed here (in the calculations of indirect

GWPs, Section 8.5) that no additional Antarctic ozone

destruction occurs for EESC values above 1990 levels. The

second metric that has been used to compare scenarios is

the integrated EESC value above the 1980 level, integrated

from 1980 or the current time until EESC returns to the 1980

level. This metric is meant to represent the cumulative ozone

depletion due to ODSs over the specified time frame.

The contributions of the various halocarbon groups

to EESC for scenario A1, as well as for scenario Ab of the

previous Assessment, are shown in Figure 8-3. The CFCs

continue to be the dominant source of EESC, although as

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.27

CFCs

1970 1990 2010 2030 2050

500

1000

1500

ppt

HCFCs

1970 1990 2010 2030 2050

100

200

300

400

CH CCl

1970 1990 2010 2030 2050

Year

100

200

300

400

ppt

CCl

1970 1990 2010 2030 2050

100

200

300

400

CH Br

1970 1990 2010 2030 2050

100

200

300

400

Halons

1970 1990 2010 2030 2050

100

200

300

400

Chlorine Sum

1970 1990 2010 2030 2050

Year

500

1000

1500

2000

2500

ppt

Bromine Sum

1970 1990 2010 2030 2050

Year

300

400

500

600

700

Figure 8-3. Contributions of halocarbon groups to global EESC for this Assessment (A1 scenario, black lines)

and the previous Assessment (Ab scenario, green lines). The “chlorine sum” and “bromine sum” represent the

total EESC from these particular chlorine and bromine long-lived source gases. EESC values for CH3Br, the

halons, and the bromine sum for this Assessment are scaled by 45/60 to eliminate the effect of the change in

the α value, for easier comparison with the previous Assessment values.
HALOCARBON SCENARIOS, ODPs, AND GWPs

8.28

a group their contribution has begun to decrease, driven

mostly by the decline of CFC-11. In terms of EESC, the

most striking differences between the A1 and Ab scenarios are the HCFCs. As discussed in Sections 8.3.2.1

and 8.3.2.2, this increasing importance of the HCFCs is

primarily due to the increase in expected future production and emission of HCFC-22 (Table 8-4). The updates

to future emissions of CFCs and HCFCs lead to slight

increases in the importance of these ODS groups to EESC

in the future. Updates to the fractional halogen release

values of the halons lead to a slightly larger contribution

of these compounds. Estimated CH3Br abundances are

also noticeably different from WMO (2003). The differences from 1996 through 2004 arise because of the availability of new global surface data (Montzka et al., 2003,

and updates). After that period, differences arise because

of the larger assumed anthropogenic fraction of CH3Br

emission compared with WMO (2003) (30% compared

with 20%), also based on Montzka et al. (2003). Overall,

the updates to the future ODS emissions projections and

the increase in the estimated value for α lead to a later

“recovery” of midlatitude-relevant EESC, defined here as

a return to pre-1980 values, by about 5 years.

In past Assessments, quantification of midlatituderelevant EESC was used exclusively. These calculations

assumed a lag of 3 years between surface and stratospheric

mixing ratios and assumed a value for α of 45. However,

it is known that the age of air in the wintertime polar vortex

is greater than 3 years (Andrews et al., 2001; Daniel et al.,

1996; Harnisch et al., 1999), and there are indications that

the value of α should be larger, due to the stronger coupling of bromine monoxide (BrO) with chlorine monoxide

(ClO) (Chipperfield and Pyle, 1998). Both of these points

act to delay the halocarbon recovery of the polar vortex

compared with midlatitudes. Consideration of the spectrum of ages that make up wintertime vortex air can further delay halocarbon recovery (Newman et al., 2006). The

time lag increase is expected to be particularly important

because of the fast increase in halocarbon abundances

in the 1980s and the relatively slow decline expected in

the future due to halocarbon lifetimes of decades. If a

time lag is assumed to be 6 years and a value of α of 65

(Chipperfield and Pyle, 1998), the year of EESC recovery

for Antarctic conditions is 2065, more than 15 years later

compared with the midlatitude EESC recovery; the time

lag increase is responsible for almost 13 of the 15 years,

with the change in α responsible for the rest. Due to the

large sensitivity of the EESC recovery to the age of air,

all scenarios will be performed for midlatitude-relevant

and polar vortex-relevant EESC. Newman et al. (2006)

have also shown that the use of a simple lag can lead to

errors in estimated EESC, particularly near the peak

period. However, because there is little difference in the

recovery date when using a simple lag or a more complicated age spectrum, and because there is no agreement on

a particular globally averaged age spectrum to use, we

continue to use a simple time lag in this Assessment.

8.3.3 Alternative Projections

8.3.3.1 EMISSIONS

The alternative hypothetical cases used to examine

the relative environmental effects of reducing future production and emissions of groups of halocarbons are compared in Table 8-6 and fall into three categories, each of

which is based on the baseline A1 scenario: (1) “no future

emission” cases (E0); (2) “no future production” cases

(P0); and (3) cases of the full capture and destruction of

the 2007 bank (B0). Assumptions of the alternative projections lead to various decreases in future emissions when

compared with the baseline scenario. The hypothetical

elimination of all future emissions (E0) represents the

greatest possible reduction in future abundances. The

hypothetical elimination of future production (P0)

allows the ODSs currently residing in banks to continue

to be released to the atmosphere as they would in the baseline scenario. The projections that consider the full capture and destruction of the 2007 bank (B0) complement

the “no production” cases in that the full 2007 bank is

destroyed in a manner that releases no ODS to the atmosphere; future production and the resulting future buildup

of banks is allowed to continue, however. This expected

future production is small for the CFCs and halons, but

significant for the HCFCs. A case consistent with the mitigation scenario presented in IPCC/TEAP (2005) is

assessed to illustrate the importance of emissions reductions from this carefully developed scenario on the future

evolution of EESC. The mitigation scenario has a significant effect only on the banks of HCFC-22, HCFC-123,

and HFCs in 2015 (the latter two are not considered in the

cases or scenarios of this chapter).

8.3.3.2 MIXING RATIOS

The calculated mixing ratios from 1990 through

2040 are shown in Figure 8-2 for the E0 and P0 cases in

addition to the previously discussed A1 and Ab scenarios.

These calculations suggest that all ODSs considered

here, except for the HCFCs and possibly halon-1301 have

already reached their peak mixing ratios. The HCFCs are

expected to reach peak abundances between 2010 and
HALOCARBON SCENARIOS, ODPs, AND GWPs

8.29

EESC dt

x

 

1980

∫ EESC dt

x

 

2007

∫

Table 8-6. Comparison of scenarios and hypothetical cases a: the year when EESC drops below the

1980 value for both midlatitude and polar vortex cases, and integrated EESC differences (midlatitude

case) relative to the baseline (A1) scenario. Note that the polar recovery times have not been given in previous Assessments; interpretation of any comparison between these numbers and recovery times given in

previous Assessments requires an understanding of the large role played by the different transport times from

the troposphere to the stratospheric midlatitude and polar vortex regions.

Scenario and Cases Percent Difference in Year (x) when EESC is

integrated EESC relative expected to drop below

to baseline scenario for 1980 value

the midlatitude case

Midlatitude Antarctic vortex b

Scenarios

A1: Baseline scenario 2048.9 2065.1

Cases a of zero production from 2007 onward of:

P0: All ODSs −8.0 −17.1 2043.1 2060.3

CFCs −0.1 −0.3 2048.8 2065.0

Halons −0.2 −0.5 2048.8 2065.1

HCFCs −5.5 −11.8 2044.4 2062.2

Anthropogenic CH3Br −2.4 −5.1 2047.9 2063.7

Cases a of zero emissions from 2007 onward of:

E0: All ODSs −19.4 −41.7 2034.0 2049.9

CFCs −5.3 −11.5 2045.0 2060.3

CH3CCl3 −0.1 −0.2 2048.9 2065.1

Halons −6.7 −14.4 2045.6 2061.9

HCFCs −7.3 −15.7 2043.7 2061.8

CCl4 −1.3 −2.9 2048.5 2064.9

Anthropogenic CH3Br −2.4 −5.1 2047.9 2063.7

Cases a of full recovery of the 2007 banks of:

B0: All ODS −12.9 −27.8 2040.8 2056.7

CFCs −5.2 −11.3 2045.1 2060.4

Halons −6.7 −14.3 2045.7 2062.0

HCFCs −1.9 −4.1 2048.4 2064.8

CH3Br sensitivity:

Same as A1, but CH3Br anthropogenic

emissions set to 20% in 1992 c 3.1 6.6 2050.6 2067.7

Same as A1, but zero QPS production

from 2015 onward −1.5 −3.2 2047.9 2063.7

Same as A1, but critical-use exemptions

continued at 2006 level 1.9 4.0-4.7 2050.1 2067.0

a Importance of ozone-depleting substances for future EESC were calculated in the hypothetical “cases” by setting production or emission to zero in

2007 and subsequent years or the bank of the ODS to zero in the year 2007 alone. These cases are not mutually exclusive, and separate effects of

elimination of production, emissions, and banks are not additive. b This metric specifically for Antarctic polar vortex ozone depletion has not been shown in any previous ozone Assessment. c In the baseline scenario, this fraction was assumed to be 30% in 1992, with a corresponding emission fraction of 0.88 of production. In this alternative scenario, an anthropogenic fraction was assumed to be 20%, with an emission fraction of 0.56 of production. In both scenarios, the total historic

emission was derived from atmospheric observations and a lifetime of 0.7 years.
2035, with scenario A1 peak mixing ratios of HCFC-22

and HCFC-141b more than 50% higher than current

values.

A comparison of the A1 curves with the P0 and the

E0 curves illustrates the relative importance of the amount

of an ODS already in existing equipment compared with

what is expected to be produced and emitted in the future.

For CFC-11 and CFC-12, for example, the amounts currently in the atmosphere and in the banks together dominate the future mixing ratio evolution. In contrast, for

HCFC-22, the future production is expected to be far more

important than the current bank. Gases like CH3CCl3,

CCl4, and CFC-113 are believed to have small banks in

applications and are not expected to be produced much in

the future; the future decline of these compounds is thus

controlled almost exclusively by their lifetimes and can

be accelerated little by any policy measures.

8.3.3.3 EQUIVALENT EFFECTIVE STRATOSPHERIC

CHLORINE

Total midlatitude-relevant EESC time series projections are shown in Figure 8-4 for scenarios A1 (baseline);

the zero production P0, zero emission E0, and full 2007

bank recovery B0 cases; and the no-Protocol scenario

(scenario H1 from WMO, 1999). The assumption of a 3-

year time lag implies that any change in emission as early

as 2007 will not affect EESC until 2010. Regions in the

stratosphere characterized by younger (older) ages will

respond more quickly (slowly). The EESC results illustrate that there is still a wide range in EESC projections

before halocarbons recover to pre-1980 conditions. The

calculated changes in EESC due to a cessation in future

production and emission are shown for CFCs, HCFCs,

and halon groups in Figure 8-5 (left column). These time

series show the importance of future HCFC production

(and associated emission) and the potential EESC reduction they can provide. For CFCs and halons, the amount

of ODSs in the banks plays a larger role than does expected

future production. This figure also shows the effect of the

same scenarios on future radiative forcing. A comparison

between the forcing and EESC estimates illustrates the

greater relative importance of the HCFCs and the small

significance of the halons when radiative forcing is considered. This large difference in significance of the halons

for ozone depletion and climate change is due to the fact

that the chemical enhancement of bromine’s importance

relative to chlorine for ozone depletion (α) plays no role

in the direct radiative forcing of bromocarbons.

The times when midlatitude-relevant and Antarcticrelevant EESC values drop below the 1980 levels are shown

in Table 8-6. Table 8-6 also includes the relative change in

integrated midlatitude EESC for the various cases and scenarios when integrating from 1980 to the time of recovery

and from 2007 to the time of recovery. For the midlatitudes,

the quantity integrated from 1980 to the time of recovery

can be thought of as a proxy for integrated ozone loss over

the whole period of loss. The quantity integrated from 2007

can be thought of as a proxy for integrated future ozone loss

(WMO, 1995, 1999, 2003). As stated earlier, such linear

relationships break down in the Antarctic stratosphere due

to the nearly complete loss of ozone over a range of altitudes, and the resulting saturation of halogen-caused ozone

depletion (Chapter 3 of WMO, 2003).

For the baseline scenario, midlatitude EESC is

expected to drop below the 1980 value in 2049. This is

about 5 years later than assessed in WMO (2003), mainly

because of larger future HCFC-22 production and emission and larger CFC-11 and CFC-12 banks and emissions

from these banks compared with WMO (2003).

If all production of anthropogenic ODSs were to

cease in 2007, the year EESC returns to the 1980 value

would be 2043 for the midlatitude case, about 6 years

earlier than for the baseline scenario, and 2060 for the

Antarctic, almost 5 years earlier than for the baseline.

HCFCs are responsible for the majority of this earlier

recovery (Table 8-6). However, because the HCFC consumption in developing countries is not limited until

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.30

Global EESC

980 1990 2000 2010 2020 2030 2040 2050

Year

1500

2000

2500

3000

3500

4000

4500

ppt

A

E

P0

B

H1

Figure 8-4. Midlatitude EESC projection calculated

for 5 scenarios and cases: A1 (baseline), zero production P0, zero emission E0, full 2007 bank

recovery B0, and the no-Protocol scenario H1 (from

WMO, 1999). Any potential contributions of VSLSs

to EESC are not included.
HALOCARBON SCENARIOS, ODPs, AND GWPs

8.31

 

-200

-150

-100

-50

pp

CFCs t

 

-0.08

-0.06

-0.04

-0.02

0.00

W/m

 

-200

-150

-100

-50

pp

HCFCs t

 

-0.08

-0.06

-0.04

-0.02

0.00

W/m

 

-200

-150

-100

-50

pp

Halons t

 

-0.08

-0.06

-0.04

-0.02

0.00

W/m

2000 2020 2040 2060 2080 2100

Year

-200

-150

-100

-50

pp

CH Br t

2000 2020 2040 2060 2080 2100

Year

-0.08

-0.06

-0.04

-0.02

0.00

W/m

EESC

 

1000

2000

3000

000

ppt

EESC

 

1000

2000

3000

000

ppt

Natural CH Cl and CH Br

CFCs

Halons

HCFCs

CH CCl and CCl

Antropogenic CH Br

A1 Scenario

Forcing

 

0.0

0.1

0.2

0.3

.4

W/m

CFCs

HCFCs

CH CCl and CCl

Forcing

 

0.0

0.1

0.2

0.3

.4

W/m

Figure 8-5. Projected EESC and radiative forcing decreases relative to the A1 scenario due to a cessation of

emission (solid curves) and production (dashed curves) in 2007 for CFCs, HCFCs, halons, and anthropogenic

CH3Br. The forcing curves for the halons and CH3Br are nearly indistinguishable from zero due to the much

lower atmospheric abundances of halons and CH3Br compared with CFC and HCFCs.
2015, the assumptions regarding this consumption in the

A1 scenario largely determine the magnitude of this earlier recovery. Smaller contributions come from the future

production of CFCs, halons, and methyl bromide.

In the hypothetical case in which all anthropogenic

emissions of ODSs were to cease in 2007, the future

mixing ratios and EESC are governed entirely by the natural destruction in the atmosphere (and ocean/soil) of the

ODSs. In this case the midlatitude EESC would drop

below the 1980 value in 2034. The largest contributions

to the zero emission cases come from HCFCs (especially

HCFC-22), CFCs, and halons.

In the hypothetical case that the 2007 bank of CFCs,

halons, and HCFCs were fully recovered and prevented

from being emitted into the atmosphere, EESC would drop

below the 1980 value 4, 3, and 0.5 years earlier, respectively, compared with the baseline scenario.

In the IPCC (IPCC/TEAP, 2005) “mitigation” scenario, economic, and technological analyses were performed in an attempt to determine reasonable reductions

in future ODS emissions that could be accomplished

using current best-practice emission reduction techniques. The mitigation scenario shows a decrease from

1879 to 1587 ktons in the bank of HCFC-22 in 2015. This

decrease in the bank leads to a reduced emission over the

period 2007-2050 from 15 to 11 Mt, contributing to an

ODP-weighted decrease of the HCFCs from 1.0 to 0.8 Mt.

The emission reduction in the mitigation scenario corresponds thereby to about 20% of the zero-emissions case

of the HCFCs. Hence, the effects of the mitigation scenario, in terms of the length of time required for the ODSs

to return to 1980 levels and the integrated EESC above

the 1980 level, are also about 20% of the effects of the

zero emissions case of the HCFCs (Table 8-6).

A few additional cases have been assessed for

methyl bromide. There is uncertainty (see Section 1.2.1.7)

regarding the fraction of anthropogenic emissions that

comprise the sources needed to balance the sinks that are

based on the observed surface mixing ratio and a global

lifetime of 0.7 years. Montzka et al. (1999) estimated that

the emission of CH3Br resulting from industrial production is between 10% and 40% of the total emission in 1992.

In the baseline scenario, this fraction is assumed to be 30%,

higher than assumed in WMO (2003), but consistent with

the discussions in Chapter 1. As a sensitivity study, an

anthropogenic fraction of 20% was assumed in an alternative scenario. In our calculations, this 20% fraction is consistent with an emission fraction of 0.56 of the industrial

production, meaning that 56% of the industrially produced

CH3Br is emitted to the atmosphere. This is, in turn, consistent with the range of 0.43-0.87 from agricultural and

related uses as reported by UNEP (1998). The magnitude

of natural emissions for this case is 165 ktons, compared

with 146 ktons in scenario A1, and compared with a total

anthropogenic production of 73 ktons in 1992. In this

alternative scenario, EESC drops below the 1980 value

almost 2 years later than in the baseline scenario, demonstrating the expected effect of a smaller significance of

anthropogenic CH3Br emission reductions in a case in

which natural emission plays a larger role.

A complete phase-out (except for critical and nonregulated uses) of the production and consumption of

methyl bromide in developed countries (non-Article

5(1)) came into effect in 2005. Several countries have

asked and obtained exemptions for critical uses of methyl

bromide. In 2005, these exemptions totaled about 14

ktons and in 2006 about 13 ktons, which is 50-60% of

the reported production in 2003-2004. In the baseline

scenario, it is assumed that the critical-use exemptions

continue at the 2006 level until 2015, when they are no

longer granted and production in Article 5(1) countries

is terminated. The existence of possible stockpiles of

methyl bromide is not taken into account in the scenarios.

If critical-use exemptions continue indefinitely at the

2006 level compared with a cessation of these exemptions in 2010 or 2015, midlatitude integrated EESC

would increase by 4.7% or 4.0%, respectively.

Production of methyl bromide for quarantine and

pre-shipment (QPS) is not controlled by the Montreal

Protocol. In the baseline scenario, the QPS production,

estimated at 10.7 ktons per year, is continued indefinitely. If the QPS production were to cease in 2015, the

year EESC is expected to drop below the 1980 value is

about a year earlier. This effect depends on the assumption of constant QPS emissions in the future in the baseline scenario and would be different if an alternative

future QPS emission assumption were made. The estimated emissions of QPS are close to those of the critical-use exemptions and they have consequently similar

contributions to the integrated EESC.

8.3.4 Uncertainties in ODS Projections

Numerous processes can affect the amount of ozone

loss for a given halogen loading level. Many of these have

been discussed in Chapters 5 and 6, and some will be mentioned in Section 8.4. Here we will focus on the uncertainties in the ODS projections themselves. Unfortunately,

little work has been performed in this area, leading to a

mostly qualitative assessment of the sensitivities of the

projections presented here.

Before future emission projections based on the

bottom-up analyses considered here can be considered fully

reliable, the causes of the discrepancies with the top-down

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.32
estimates need to be better understood. Whether the

missing emission from the bottom-up analysis is due to

faster release from the bank than estimated, additional

past production that has led to a larger bank and a correspondingly higher bank release, additional production and

immediate release from use as solvent or aerosol propellant, or another factor, the particular assumptions made can

lead to important differences in future ODS projections. A

measure of the sensitivity of a particular gas’s mixing ratio

projection to the size of the bank, compared with future

production uncertainties, can be obtained from Figures 8-2

and 8-5. For example, CFC-11 and the CFC group will

have more sensitivity to uncertainties in the size of its bank,

while HCFC-22 and the HCFC group will have more sensitivity to uncertainties in future production. While it is

important to recognize the large emission discrepancies

between these bottom-up analysis and observations for a

few species, it is also important to consider the species that

demonstrate relatively good agreement with emissions

(CFC-12, HCFC-22, and halon-1211). Even for these

species, there are important differences in the sizes of the

estimated bank, depending on the method used to estimate

this quantity. Table 8-7 shows a comparison of bank estimates from this Assessment, the previous Assessment

(WMO, 2003), and IPCC/TEAP (2005). The large discrepancies between the CFC-11 and CFC-12 banks have been

shown by Daniel et al. (2006) to have important implications for estimating the benefit of recycling/destroying the

CFC banks. While the banks are calculated in this

Assessment in a manner to agree best with the IPCC/TEAP

(2005) estimates for many of the considered ODSs, the

cause of this discrepancy has not been resolved.

Another potential explanation for some of the bank

size discrepancy could be due to an error in the assumed

factor of 1.07 that relates the surface mixing ratio to the

globally averaged mixing ratio. A higher (lower) value

for this factor would lead to smaller (larger) emissions

estimated from atmospheric mixing ratio observations,

and thus to larger (smaller) top-down bank sizes. Daniel

et al. (2006) found, for example, that an increase in this

factor to 1.10 from 1.07 would lead to an additional

amount of banked CFC-11 in 2002 of 250 ktons, representing about 20% of the difference between the top-down

estimate of WMO (2003) and the bottom-up estimate of

IPCC/TEAP (2005).

Changes in atmospheric lifetimes, whether due to

changing atmospheric dynamics or chemistry, can also

affect future ODS projections. The recently published suggestion of an important soil sink that could reduce the

carbon tetrachloride lifetime from 26 to 20 years (Happell

and Roche, 2003), and its resulting implications for a

quicker recovery of EESC to 1980 values, is one example

of the importance of lifetime uncertainties. Daniel et al.

(2006) have shown that an increase in the CFC-11 lifetime

by 10% would result in reduced annual emissions required

to be consistent with observed mixing ratios that would

lead to an increase in the 2002 bank by 300 ktons. The

12-box model emission estimates discussed in Section

8.3.2.1 can presumably account for changes in global lifetimes due to changing atmospheric distributions, as well

as changes in the factor relating surface mixing ratio to

the global mixing ratio average, better than the 1-box

model. These differences likely play a role in the higher

emissions estimates of the 12-box model (Section 8.3.2.1);

however, these higher emissions would lead to smaller

bank estimates using the top-down formalism in this

chapter, and thus to even larger bank size discrepancies

with the bottom-up analysis.

The release rate applied to the particular usage

banks can affect the accuracy of future ODS projections.

For example, there is a question as to how quickly ODSs

are released from foam insulation once it is sent to a landfill (IPCC/TEAP, 2005). If foam ODSs were not released

for 40 years, recycling these particular ODSs would have

a much reduced benefit than if release were much sooner.

This results because by the time the ODS would be finally

released in the absence of recycling, EESC would have

almost returned to pre-1980 levels in extrapolar regions of

the lower stratosphere.

Sources of uncertainty in relating tropospheric mixing ratios to EESC include location- and time-dependent

fractional chlorine/bromine release, age of air, and, α values.

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.33

Table 8-7. Comparison of halocarbon banks

(ktons) in 2002 used in this Assessment, the previous Assessment (WMO, 2003), and IPCC/TEAP

(2005).

Species WMO (2003) IPCC/TEAP This

2002 Bank (2005) Assessment

2002 Bank 2002 Bank

CFC-11 594 1,687 1,654

CFC-12 0 711 711

CFC-113 7 0 5

HCFC-22 1,317 1,531 1,531

HCFC-141b 753 836 674

HCFC-142b 210 224 224

Halon-1211 72 125 125

Halon-1301 58 42 42

CH3CCl3 530 *

* Production and bank were not estimated; we assume emissions are

equal to production, implying a zero bank. Protocol limitations are

placed directly on emissions.
We have given examples of the potential significance of age

of air and α variations, but there have been no studies examining the importance of the fractional halogen release variations with location. To the extent that much of the ozone

depletion occurs in the lower stratosphere, where the release

values remain relatively constant with location (Solomon et

al., 1992), it has generally been assumed that any variability

would represent a minor effect. The fractional chlorine

release values of HCFC-141b and HCFC-142b represent

two cases in which there is significant uncertainty regarding the correct value. The previous Assessment discussed

the large differences between the previously assumed

model-derived values and the values estimated from observations (Schauffler et al., 2003). The model-derived values

were retained because of the large importance of the ageof-air correction in interpreting observations of gases

with such a large growth rate. However, it is unlikely that

the age-of-air correction could be responsible for such a

large discrepancy. Unfortunately, no work has been performed since the previous Assessment to resolve these

questions. Therefore, we continue to use the same values

that were used in the previous Assessment, but caution that

these values could be too large. While the ODPs of HCFC141b and -142b are much larger than the ODP of HCFC22, the effect of this fractional chlorine release uncertainty

is limited by the lower projected mixing ratios of these

two HCFCs compared with HCFC-22 (Figure 8-2). Nevertheless, uncertainties in the fractional release values for

many of the ODSs do represent an important uncertainty in

EESC calculations. These release values affect the relative

importance of certain ODSs to total EESC and can thus

affect future EESC evolution. If, for example, the Schauffler

et al. release values (2- to 4.5-year means) were used instead

of the ones assumed in this chapter, the EESC return to 1980

levels would occur about 3 years earlier.

It is important to recognize that errors in the accurate modeling of the baseline scenario can translate

directly into errors in interpreting the alternative scenarios

and test cases. For example, an error in the value of α

will lead directly to an error in the assessment of the zero

emission and production test cases for reducing EESC,

particularly for the bromocarbons. As another example,

any unreasonable assumption in the baseline scenario for

future production of any species will necessarily affect

the conclusions derived from both the “no production”

and “no emission” test cases.

8.4 OTHER PROCESSES RELEVANT TO

FUTURE OZONE EVOLUTION

To enable policy-relevant discussion of future scenarios, this chapter has made several simplifying assumptions. This chapter characterizes the complexities of

ozone depletion using simple parameterizations based on

ODPs. But perhaps more important, the scenarios are

assumed to evolve within an unchanging background

state. Given projections of climate change (e.g., IPCC,

2001), and our understanding of ozone and climate interactions (WMO, 2003; IPCC/TEAP, 2005; Chapter 5 and

Chapter 6), this assumption is known to be inaccurate.

This section qualitatively discusses a few possible consequences of a changing background state on the relationship between ODS emissions, EESC, and stratospheric

ozone depletion. More details regarding the effects of

climate change on the future ozone evolution can be found

in Chapters 5 and 6.

Both natural and anthropogenic forcings of climate

change have the potential to alter important chemistry

related to future stratospheric ozone depletion for a given

future evolution of ODSs and EESC. Two primary natural forcing mechanisms include solar changes and volcanic eruptions. Currently, solar changes are not expected

to have a long-term effect on ozone recovery. However, a

volcanic eruption that injects a large amount of sulfate

aerosols into the stratosphere in the next few decades may

enhance halogen ozone destruction. While this should

perturb ozone chemistry for only a few years, the perturbations during that time could be substantial.

Future anthropogenic emissions of gases like

carbon dioxide (CO2), nitrous oxide (N2O), and methane

(CH4) will also affect the chemical response of ozone

depletion to ODS emissions and EESC. The expected

stratospheric cooling induced by increasing concentrations, primarily of CO2, is expected to slow gas-phase

ozone depletion reactions and increase global ozone;

conversely, polar springtime ozone could be reduced by

cooler temperatures and the resulting increase in heterogeneous ozone loss. The expected future increase in

N2O will increase stratospheric NOx, which is expected

to exacerbate globally averaged ozone depletion. However, increases in CH4 abundances could accelerate

ozone recovery. Hence, the globally averaged net effect

of increases in these two important trace species depends

on the specific future emission scenario. Emissions of

hydrocarbons and NOx are also expected to affect the

lifetime and concentrations of stratospheric source

gases, such as HCFCs, CH3Br, CH3Cl, and CH4, and

thereby affect ozone.

Climate change by any cause (human, natural, or

variability) can also be expected to affect atmospheric

dynamics, which could lead to changes in ozone. These

changes could be due, for example, to changes in the

transport of ozone itself, or to changes in the lifetimes of

the source gases.

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.34
So while this chapter examines the changes in

EESC as a proxy for stratospheric ozone loss, it should be

noted that neither total column amounts nor especially the

latitudinal and vertical distribution of ozone are expected

to return to their pre-1980 state at the same time that EESC

returns to pre-1980 levels.

8.5 INDIRECT GWPS

ODSs can affect climate through their direct radiative forcing (Section 8.2) in a similar manner to other

greenhouse gases. They also can uniquely affect climate

by their destruction of lower stratospheric ozone, itself a

greenhouse gas. This ozone destruction leads to a negative radiative forcing, and in the global mean acts in

opposition to the ODSs’ greenhouse warming. Potential

implications of this offset have been discussed in Solomon

and Daniel (1996) and include the possibility that the

ozone destruction by the ODSs over the last 25 years may

have offset part of the globally averaged surface warming

due to the ODSs that would otherwise have been experienced. In past ozone Assessments, this indirect forcing

offset has been combined with the direct Global Warming

Potentials, to produce a “net” GWP. However, recent work

suggests that the surface temperature response to recovering stratospheric ozone may be larger than suggested by

radiative forcing alone (Joshi et al., 2003). As a potential

remedy, it has been suggested that GWPs could be modified to account for the efficacy of any given forcing at

modifying surface temperature (Berntsen et al., 2005;

Fuglestvedt et al., 2003). Studies also have shown that

although it may be appropriate to combine these direct and

indirect effects when evaluating long-term global mean

surface temperature response, it may not be appropriate to

simply offset these effects when evaluating other aspects

of the climate response. Many physical characteristics of

the direct and indirect climate change effects, especially

in the stratosphere, do not act to offset each other (Forster

and Joshi, 2005; IPCC/TEAP, 2005). Furthermore, GWP

calculations for gases in the troposphere show that indirect GWPs vary with emission location (Berntsen et al.,

2005). Due to these complications, in this report we follow

the IPCC/TEAP (2005) formalism and do not present net

GWPs, but keep direct and indirect GWPs separate.

The updated indirect GWP values for selected

halocarbons are shown in Table 8-8. These values are

calculated using a formalism similar to what was used in

past Assessments and are based on discussions in Daniel

et al. (1995). Specifically, a pulse of an ODS can lead to

additional ozone depletion, which will be associated with

a negative radiative forcing. For indirect GWP calculations, it is assumed that the ozone depletion occurs at

midlatitudes when the total EESC amount is above the

1980 level. Below this level, it is assumed that no additional ozone loss occurs due to the pulse emission. In

the stratosphere inside the polar vortex, additional loss

from the pulse is assumed to occur when the EESC value

is above the 1980 level and below the 1990 level. Above

the 1990 level, it is assumed that due to ozone loss saturation over a large altitude range, no additional loss

occurs due to the pulse emission. In past Assessments

and here, all EESC values used in the indirect calculations have been generated with a 3-year time lag. If a 6-

year time lag were used for the polar EESC calculations

along with a value of 65 for α, the indirect GWPs would

change by less than 20%.

For the purpose of calculating these indirect GWPs,

the ozone radiative forcing assumed to be due to changes

in chlorine and bromine from 1979 to 1997 is the same as

discussed in IPCC (2001) and IPCC/TEAP (2005), which

is −0.15 ± 0.10 W m-2. Hansen et al. (2005) calculate a

value of −0.06 W m-2, but suggest that this may be a lower

limit for the forcing due to chlorine- and bromine-induced

ozone depletion because they do not consider tropospheric ozone trends outside the polar regions that may

HALOCARBON SCENARIOS, ODPs, AND GWPs

8.35

Table 8-8. Direct and indirect Global Warming

Potentials of selected halocarbons calculated for

a 100-year time horizon.

Gas Direct GWP 1 Indirect GWP 2

CFC-11 4,750 ± 1,660 −3,790 ± 2,620

CFC-12 10,890 ± 3,810 −2,160 ± 1,520

CFC-113 6,130 ± 2,150 −2,530 ± 1,780

HCFC-22 1,810 ± 630 −286 ± 192

HCFC-123 77 ± 27 −83 ± 55

HCFC-124 609 ± 213 −120 ± 81

HCFC-141b 725 ± 254 −667 ± 447

HCFC-142b 2,310 ± 810 −362 ± 244

CH3CCl3 146 ± 51 −643 ± 431

CCl4 1,400 ± 490 −3,630 ± 2,470

CH3Br 5 ± 2 −2,150 ± 1,440

Halon-1211 1,890 ± 660 −40,280 ± 27,120

Halon-1301 7,140 ± 2,500 −49,090 ± 34,280

Halon-2402 1,640 ± 570 −62,000 ± 41,930

HCFC-225ca 122 ± 43 −93 ± 50

HCFC-225cb 595 ± 208 −156 ± 85

1 Direct GWP uncertainties represent 35% of the direct value. 2 Indirect GWP uncertainties represent 1σ ranges, including a ±10-year

1σ ozone recovery uncertainty and a 67% 1σ ozone forcing uncertainty. The ozone forcing uncertainty is responsible for more than

90% of the total stated uncertainty. 
have arisen due to chlorine and bromine changes. A recent

study that compares model calculations of ozone changes

since preindustrial times suggests a range of −0.123 to

+0.066 W m-2 for the forcing due to stratospheric ozone

changes through the year 2000, when climate changes are

neglected (Gauss et al., 2006). However, we have chosen

to retain the previous forcing estimate because the Gauss

et al. (2006) study does not quantify the mitigating forcing effects of tropospheric ozone and ozone precursor

increases on stratospheric ozone, particularly in the lower

stratosphere, where ozone changes are so important to

radiative forcing calculations. It would be inappropriate

to include the effect of tropospheric ozone chemistry not

due to halogens in the indirect GWP estimates of chlorocarbons and bromocarbons.

The differences between the indirect GWP estimates in Table 8-8 and those in WMO (2003) (net GWPs

minus direct GWPs of Table 1-8) are attributable to: a

new ODS scenario that leads to a return to pre-1980 levels

of EESC later; an updated CO2 mixing ratio; a new CO2

pulse response function; a change in the year of emission from 2002 to 2007; the updated value of α; and different fractional bromine release values that are now

consistent with Table 8-1 (and Table 1-4 in WMO, 2003)

rather than the values used for CH3Br, halon-1211, and

halon-1301 of 1.08, 1.1, and 0.8, respectively, used in

Table 1-8 of WMO (2003). Differences compared with

IPCC/TEAP (2005) are due to the updated CO2 mixing

ratio and response function, to the new ODS scenario, and

to the different value of α.

The direct GWP uncertainties represent 35% of the

direct GWP values, consistent with Section 8.2.3. The

indirect uncertainties represent a 1σ uncertainty due to a

±10-year 1σ error in the year of EESC recovery to 1980

levels and a 67% 1σ uncertainty in the ozone radiative

forcing; this ozone forcing uncertainty accounts for nearly

all of the quoted uncertainty total. It should be noted that

these errors associated with GWPs are in addition to the

potentially large difference in climate sensitivity (i.e., the

amount of temperature change per change in radiative

forcing, dT/dF) between stratospheric ozone loss and the

well-mixed greenhouse gases (Joshi et al., 2003). While

the climate sensitivity does not affect the GWP values, it

does have relevance to the comparability of GWPs in

assessing the impact of greenhouse gases on climate.

There remains a large disparity in estimates of the variations of this climate sensitivity of ozone depletion in the

upper troposphere and lower stratosphere; if this uncertainty can be reduced in the future, it may be advantageous to include this in assessing the relative climate

impact of halocarbons (e.g., Berntsen et al., 2005;

Fuglestvedt et al., 2003).

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8.39
TWENTY QUESTIONS AND ANSWERS

ABOUT THE OZONE LAYER: 2006 UPDATE

Lead Author:

D.W. Fahey

The update of this component of the Assessment was discussed by the 77 scientists who attended the Panel Review

Meeting for the 2006 Ozone Assessment (Les Diablerets, Switzerland, 19-23 June 2006). In addition, subsequent contributions, reviews, or comments were provided by the following individuals: S.A. Montzka (special recognition), R.J.

Salawitch (special recognition), D.L. Albritton, S.O. Andersen, P.J. Aucamp, M.P. Baldwin, A.F. Bias, G. Bodeker, J.F.

Bornman, G.O. Braathen, J.P. Burrows, M.-L. Chanin, C. Clerbaux, M. Dameris, J.S. Daniel, S.B. Diaz, E.G. Dutton,

C.A. Ennis, V. Eyring, V.E. Fioletov, N.P. Gillet, N.R.P. Harris, M.K.W. Ko, L. Kuijpers, G.L. Manney, R.L. McKenzie,

R. Müller, E.R. Nash, P.A. Newman, T. Peter, A.R. Ravishankara, A. Robock, M.L. Santee, U. Schmidt, G. Seckmeyer,

T.G. Shepherd, R.S. Stolarski, W.T. Sturges, J.C. van der Leun, G.J.M. Velders, D.W. Waugh, C.S. Zerefos.

TWENTY QUESTIONS AND ANSWERSABOUT THE OZONE LAYER: 2006 UPDATE

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.9

Q7. What emissions from human activities lead to ozone depletion? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.11

Q8. What are the reactive halogen gases that destroy stratospheric ozone? . . . . . . . . . . . . . . . . . . . . . . . . . Q.14

Q9. What are the chlorine and bromine reactions that destroy stratospheric ozone? . . . . . . . . . . . . . . . . . . Q.17

Q10. Why has an “ozone hole” appeared over Antarctica when ozone-depleting gases are present

throughout the stratosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.19

III. STRATOSPHERIC OZONE DEPLETION

Q11. How severe is the depletion of the Antarctic ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.22

Q12. Is there depletion of the Arctic ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.26

Q13. How large is the depletion of the global ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.28

Q14. Do changes in the Sun and volcanic eruptions affect the ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . Q.30

IV. CONTROLLING OZONE-DEPLETING GASES

Q15. Are there regulations on the production of ozone-depleting gases? . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.32

Q16. Has the Montreal Protocol been successful in reducing ozone-depleting gases

in the atmosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.34

V. IMPLICATIONS OF OZONE DEPLETION

Q17. Does depletion of the ozone layer increase ground-level ultraviolet radiation? . . . . . . . . . . . . . . . . . . Q.37

Q18. Is depletion of the ozone layer the principal cause of climate change? . . . . . . . . . . . . . . . . . . . . . . . . . Q.40

VI. STRATOSPHERIC OZONE IN THE FUTURE

Q19. How will recovery of the ozone layer be identified? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.43

Q20. When is the ozone layer expected to recover? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.45

ADDITIONAL TOPICS

• Global Ozone Dobson Network . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.8

• Understanding Stratospheric Ozone Depletion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.10

• Heavier-Than-Air CFCs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.13

• Replacing the Loss of Ozone in the Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.16

• The Discovery of the Antarctic Ozone Hole . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.21

• The Anomalous 2002 Antarctic Ozone Hole . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.25

TWENTY QUESTIONS: 2006 UPDATE

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

radiation 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 may 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 predict how much ozone depletion was

occurring and how much more could occur in the future.

Observations of the ozone layer showed that depletion was indeed occurring. The most severe and most surprising ozone loss was discovered to be recurring in

springtime over Antarctica. The loss in this region 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, we know that if

ozone-depleting gases were to continue to accumulate in

the atmosphere, the result would be more depletion of the

ozone layer.

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

this global issue. As a result of the broad compliance with

the Protocol and its Amendments and Adjustments and,

of great significance, industry’s development of “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 by the late

21st century. The 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, actions, and verification. It is a

story written by many: scientists, technologists, economists, legal experts, and policymakers. And, dialogue has

been a key ingredient.

To help foster continued interaction, this component

of the Scientific Assessment of Ozone Depletion: 2006

presents 20 questions and answers about the often-complex science of ozone depletion. The answers are updates

of those first presented in the previous ozone Assessment,

Scientific Assessment of Ozone Depletion: 2002. 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

2006 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 The update of this component of the Assessment was discussed by the 77 scientists who attended the Panel Review Meeting for the

2006 Ozone Assessment (Les Diablerets, Switzerland, 19-23 June 2006). In addition, subsequent contributions, reviews, or comments were provided by the following individuals: S.A. Montzka (special recognition), R.J. Salawitch (special recognition), D.L.

Albritton, S.O. Andersen, P.J. Aucamp, M.P. Baldwin, A.F. Bias, G. Bodeker, J.F. Bornman, G.O. Braathen, J.P. Burrows, M.-L.

Chanin, C. Clerbaux, M. Dameris, J.S. Daniel, S.B. Diaz, E.G. Dutton, C.A. Ennis, V. Eyring, V.E. Fioletov, N.P. Gillet, N.R.P.

Harris, M.K.W. Ko, L. Kuijpers, G.L. Manney, R.L. McKenzie, R. Müller, E.R. Nash, P.A. Newman, T. Peter, A.R. Ravishankara, A.

Robock, M.L. Santee, U. Schmidt, G. Seckmeyer, T.G. Shepherd, R.S. Stolarski, W.T. Sturges, J.C. van der Leun, G.J.M. Velders,

D.W. Waugh, C.S. Zerefos.

TWENTY QUESTIONS: 2006 UPDATE

Q.3

Ozone is a gas that is naturally present in our atmosphere. Ozone has the chemical formula O3 because an

ozone molecule contains three oxygen atoms (see Figure

Q1-1). 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 óζειν (ozein in Latin), meaning “to smell.” Ozone

has a pungent odor that allows it to be detected even at

very low amounts. Ozone will rapidly react with many

chemical compounds and is explosive in concentrated

amounts. Electrical discharges are generally used to produce ozone for industrial processes such as air and water

purification and bleaching 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). Most ozone resides in the lower 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 = 1,000 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 imagine 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). 

Figure Q1-2. Atmospheric ozone. Ozone is present

throughout the lower atmosphere (troposphere and

stratosphere). 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.

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.”

Ozone and Oxygen

Ozone in the Atmosphere
TWENTY QUESTIONS: 2006 UPDATE

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 in comparison,

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 ozone

and processes that destroy ozone. The balance is determined by the amounts of reacting gases and 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-producing reactions in a certain location,

ozone abundances will increase. Similarly, if conditions

change to favor reactions that destroy ozone, 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).

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, each atom then undergoes a binding

collision with another oxygen molecule to form an ozone

molecule. In the overall process, three oxygen molecules plus sunlight 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 an oxygen molecule (O2 ) being broken apart 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.

Stratospheric Ozone Production
Natural ozone. In the absence of human activities

on Earth’s surface, ozone would still be present near the

surface and throughout the troposphere and stratosphere

because ozone is a natural component of the clean atmosphere. All ozone molecules are chemically identical, with

each containing three oxygen atoms. However, ozone in

the stratosphere (good ozone) has very different environmental consequences for humans and other life forms than

excess ozone in the troposphere near Earth’s surface (bad

ozone). Natural ozone in the troposphere is also considered “good” because it initiates the chemical removal of

many pollutants, such as carbon monoxide and nitrogen

oxides, as well as greenhouse gases such as methane.

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, increased exposure to UV-B increases the risk of

skin cancer (see Q17), cataracts, and a suppressed immune

system. 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 halogen source gases released in human

activities 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 regulation of ozonedepleting gases (see Q15 and Q16).

Bad ozone. Excess ozone formed near Earth’s surface in reactions caused by the presence of human-made

pollutant gases is considered bad ozone. Increased ozone

amounts are harmful to humans, plants, and other living

systems because ozone reacts strongly to destroy or alter

many other molecules. Excessive ozone exposure reduces

crop yields and forest growth. In humans, exposure to

high levels of ozone can reduce lung capacity; cause chest

pains, throat irritation, and coughing; and worsen preexisting 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 excess ozone production

near Earth’s surface.

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 of UV-B

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, excess ozone at Earth’s surface that

is formed from pollutants is considered “bad” ozone because it can be harmful to humans, plants, and animals.

The ozone that occurs naturally near the surface and in the lower atmosphere is also beneficial because ozone

helps remove pollutants from the atmosphere.

TWENTY QUESTIONS: 2006 UPDATE

Q.5

UV Protection by the Ozone Layer
TWENTY QUESTIONS: 2006 UPDATE

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. Since

about 1980, regions of low total ozone have occurred 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 troposphere extends to

a higher altitude in the tropics, and consequently, the

thickness of the ozone layer is smallest there.

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 different

locations over the globe. Reductions in ultraviolet radiation from the sun in its 11-year cycle, for example, reduce

the production of ozone.

Scientists have a good understanding of how chemistry and air motions 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 and longer. 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).

Q.6

Global Satellite Maps of Total Ozone
TWENTY QUESTIONS: 2006 UPDATE

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: local 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.

Local measurements. Local 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. The balloons ascend far enough in the

atmosphere to measure ozone in the stratospheric ozone

layer. Ozonesondes are launched regularly at many locations around the world. Local 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 nearly the entire globe on a

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 locally 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 regularly 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 on

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.

Measuring Ozone in the Atmosphere
TWENTY QUESTIONS: 2006 UPDATE

Q.8

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, now called a Dobson spectrophotometer, 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 location of the instrument. 

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).
TWENTY QUESTIONS: 2006 UPDATE

Q.9

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, at Earth’s surface, of source

gases containing the halogens chlorine and bromine (see

Q7). The halogen source gases include manufactured

chemicals released to the atmosphere by a variety of human

activities. Chlorofluorocarbons (CFCs) are an important

example of chlorine-containing gases. Emitted 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. However, small amounts

of these gases dissolve or are taken up 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). The rate of conversion is related

to the atmospheric lifetime of a gas (see Q7). Source

gases with lifetimes greater than a few years may circulate between the troposphere and stratosphere multiple

times before full conversion occurs. 

The reactive gases formed in the eventual conversion

of the halogen source gases react chemically to destroy

ozone in the stratosphere (see Q9). The average depletion

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, at Earth’s surface,

of ozone-depleting gases containing chlorine and bromine. Most of these gases accumulate in the lower

atmosphere because they are unreactive and do not dissolve readily in rain or snow. Eventually, these emitted source 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 natura processes.

Halogen source gases accumulate in the

atmosphere and are distributed throughou he

lower atmosphere by winds and other air motions.

Halogen source gases are transported to the

stratosphere by ai motions.

Most halogen source gases are converted in the

stratosphere o reactive halogen gases in chemical

eactions involving ultraviolet radiation from the Sun.

Air containing reactive halogen gases returns

o the troposphere and these gases are removed

rom the air by moisture in clouds and rain.

Reactive halogen gases cause chemical

depletion of stratospheric total ozone ove h e

globe except at tropica l atitudes.

Emissions

Accumulation

Transport

Chemical reaction

Conversion

Removal

Polar stratospheric clouds increase ozone

depletion by reactive halogen gases

causing severe ozone loss n polar regions

n winter and spring.

Principal Steps in the Depletion

of Stratospheric Ozone
TWENTY QUESTIONS: 2006 UPDATE

Q.10

of total ozone attributed to reactive gases is smallest in

the tropics and largest at high 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 substantial ozone

destruction in polar regions in winter and 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 have conducted observations to find out which gases are present in various regions of the

stratosphere and at what concentrations. They have monitored the change in these abundances over time

periods spanning a daily cycle to decades. 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 (remotely) using

optical and microwave signals. High-altitude aircraft and balloon instruments detect both gases locally in

the stratosphere (see Q5). For example, these observations show that ClO is present at elevated amounts

in the Antarctic and Arctic stratospheres in the late winter/early spring season, when the most severe ozone

depletion occurs (see Q8).
TWENTY QUESTIONS: 2006 UPDATE

Q.11

Principal 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), historically have been the most important chlorine-containing gases that are emitted by human

activities and destroy stratospheric ozone (see Figure Q7-

1). These and other 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 the principal 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).

Other human sources of chlorine and bromine.

Other chlorine- and bromine-containing gases are released

regularly in human activities. Common examples are the

use of chlorine gases to disinfect swimming pools and

wastewater, fossil fuel burning, and various industrial

processes. These activities do not contribute significantly

to stratospheric amounts of chlorine and bromine because

either the global source is small or the emitted gases are

short-lived (very reactive or highly soluble) and, therefore, are removed from the atmosphere before they reach

the stratosphere. 

Natural sources of chlorine and bromine. There

are a few halogen source gases present in the stratosphere

that have large natural sources. These include 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 17%

of the chlorine currently in the stratosphere and about 30%

of the bromine (see Figure Q7-1). Very short-lived source

gases containing bromine, such as bromoform (CHBr3),

are also released to the atmosphere primarily from the

oceans. Only a small fraction of these emissions reaches

the stratosphere, because these gases are rapidly removed

in the lower atmosphere. The contribution of these very

short-lived gases to stratospheric bromine is estimated to

be about 24%, but this has a large uncertainty. The contribution to stratospheric chlorine of short-lived chlorinated gases from natural and human sources is much

smaller (< 3%) and is included in the “Other gases” category in Figure Q7-1. 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 naturally 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 brominecontaining 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 substantially

destroyed in the troposphere, and therefore only a fraction of each 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

Q7: What emissions from human activities lead to ozone depletion?

Certain industrial processes and consumer products result in the emission of “halogen source gases” to the

atmosphere. These gases bring chlorine and bromine to the stratosphere, which cause depletion of the ozone

layer. For example, chlorofluorocarbons (CFCs), 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 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.
TWENTY QUESTIONS: 2006 UPDATE

Q.12

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 (see Q18). 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, because bromine is much more effective overall (about 60 times) on a per-atom basis than chlorine in chemical reactions that destroy ozone in the stratosphere. The gases with small ODP values generally have

short atmospheric lifetimes or fewer chlorine and bromine

atoms. 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. Many 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 cause ozone

depletion. Iodine is a component of several gases that are

naturally emitted from the oceans. Although iodine can

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 by human activities. These partitioned

columns show how the principal chlorine and bromine source gases contribute to the respective total amounts of chlorine

and bromine as measured in 2004. Note the large difference in the vertical scales: total chlorine in the stratosphere is

160 times more abundant than total bromine. 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. 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 “Other gases” category

includes minor CFCs and short-lived gases. 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 total bromine than of total chlorine. (The unit “parts per trillion” is used here as a

measure of the relative abundance of a gas in air: 1 part per trillion indicates the presence of one molecule of a gas per

trillion other air molecules.)

Primary Sources of Chlorine and Bromine for the Stratosphere in 2004
TWENTY QUESTIONS: 2006 UPDATE

Q.13

participate in ozone destruction reactions, these iodinecontaining source gases generally have very short lifetimes and, as a result, most are removed in the troposphere

before they 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 in the stratosphere to form water vapor and

reactive hydrogen, and nitrogen oxides, respectively.

These reactive products also participate in the production and loss balance of stratospheric ozone (see Q2).

The overall effect of increases in these other gases on

ozone is much smaller than that caused by increases in

chlorine- and bromine-containing gases from human

activities (see Q18).

Heavier-Than-Air CFCs

CFCs and other halogen source gases reach the stratosphere despite the fact that they are “heavier

than air.” All the principal source gases are emitted and accumulate in the lower atmosphere (troposphere).

The distributions of gases in the troposphere and stratosphere are not controlled by the molecular weight of

the gases because air is in continual motion in these regions as a result of winds and convection. Air

motions ensure that most source gases become 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 industries and governments.

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, does the influence of winds and

convection diminish to the point where heavy gases begin to separate from lighter gases as a result of

gravity.

Table Q7-1. Atmospheric lifetimes, emissions,

and Ozone Depletion Potentials of halogen

source gases. a

Halogen Atmospheric Global Ozone

Source Lifetime Emissions Depletion

Gas (years) in 2003 b Potential

(ODP) d

Chlorine

CFC-12 100 101-144 1

CFC-113 85 1-15 1

CFC-11 45 60-126 1

Carbon tetrachloride

(CCl4) 26 58-131 0.73

HCFCs 1-26 312-403 0.02-0.12

Methyl chloroform

(CH3CCl3) 5 ~20 0.12

Methyl chloride 1.0 1700-13600 0.02

Bromine

Halon-1301 65 ~3 16

Halon-1211 16 7-10 7.1

Methyl bromide

(CH3Br) 0.7 160-200 0.51

Very short-lived gases

(e.g., CHBr3) < 0.5 c c

a Includes both human activities and natural sources.

b Emission in gigagrams per year (1 gigagram = 109 grams =

1000 metric tons). c Estimates are uncertain for most species. d Values are calculated for emissions of equal mass for each gas.
TWENTY QUESTIONS: 2006 UPDATE

Q.14

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 in the stratosphere, the

halogen source gases chemically convert at different rates

to form the reactive halogen gases. The conversion occurs

in the stratosphere instead of the troposphere because solar

UV radiation is more intense in the stratosphere.

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 to form ClO (see Q10).

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 local and remote measurement techniques. The

measurements from space at middle latitudes displayed in

Q8: What are the reactive halogen gases that destroy stratospheric ozone?

Emissions from human activities and natural processes include large sources of chlorine- and bromine-containing gases that eventually reach 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 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. The reactive gases separate into reservoir gases, which do not destroy ozone, and reactive gases, which participate in ozone destruction cycles (see Q9).
TWENTY QUESTIONS: 2006 UPDATE

Q.15

Figure Q8-2 are representative of how the amounts of

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 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 at midlatitudes, 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 in the lower

stratosphere than for reactive chlorine, in part because of

the lower abundance of bromine. The most widely

observed bromine gas is bromine monoxide (BrO).

Recent observations have shown that measured BrO abundances in the stratosphere are larger than expected from

the conversion of the halons and methyl bromide to BrO,

suggesting a significant contribution from the very shortlived bromine-containing gases.

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.

Total

available

chlorine

Figure Q8-2. Reactive chlorine gas observations.

The abundances of chlorine source gases and reactive chlorine gases as measured from space are displayed with altitude for a midlatitude location. In the

troposphere (below about 10 kilometers), all chlorine is

contained in the source gases. In the stratosphere,

reactive chlorine gases increase with altitude as chlorine source gases decrease. This is a consequence of

chemical reactions involving ultraviolet sunlight (see

Figure Q8-1). The principal reactive gases formed are

HCl, ClONO2, and ClO. Summing the source gases

with the reactive gases gives total 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.)

Measurements of Reactive Chlorine from Space
TWENTY QUESTIONS: 2006 UPDATE

Q.16

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 this case,

the quantities are very small in comparison with other tropospheric sources. 

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)

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 (see Q10). 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

in late winter/early spring. Ozone values measured

simultaneously within the ozone layer show very

depleted values.

Replacing the Loss of Ozone in the Stratosphere

The idea is sometimes put forth that humans could replace the loss of global stratospheric ozone by

making ozone and transporting it to the stratosphere. Ozone amounts in the stratosphere reflect a balance

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 stratospheric 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 3,000 megatons (1

megaton = 1 billion kilograms) with most residing in the stratosphere. The replacement of the average

global ozone loss of about 4% would require 120 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. Concerns for a global delivery system would include further significant energy use and

unforeseen environmental consequences.

Satellite Observations in the Lower Stratosphere
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), 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 happens to react with another gas, breaking 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/early 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 + O Cl + O2

C + O 3 ClO + O2

Net: O + O3 2O2

Oxygen molecule (O2

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. Certain ozone destruction reactions become most

effective in polar regions because the reactive gas chlorine monoxide reaches very high levels there in the late

winter/early spring season.

TWENTY QUESTIONS: 2006 UPDATE

Q.17

Ozone Destruction Cycle 1
TWENTY QUESTIONS: 2006 UPDATE

Q.18

season (see Q11 and Q12). At high ClO abundances, the

rate of ozone destruction can reach 2 to 3% per day in

late winter/early spring.

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 ultraviolet

sunlight is most intense.

Cycles 2 and 3 require visible sunlight 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/early 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 midwinter in the polar stratospheres. The

visible sunlight needed in Cycles 2 and 3 is not sufficient

to form ozone because this process requires ultraviolet

sunlight. In the stratosphere in the late winter/early spring

period, ultraviolet sunlight is weak because Sun angles

are low. As a result, ozone is destroyed by Cycles 2 and 3

in the sunlit winter stratosphere but is not produced in significant amounts.

Other reactions. Global ozone abundances are

controlled by many 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

(ClO)2 + sunlight 

ClOO

2(Cl + O3

Net: 2O3

ClO + BrO Cl + Br + O2

BrCl + sunlight Cl + Br

Net: 2O3 3O2

ClO + BrO BrCl + O2

Cl + O3 ClO + O2

Br + O3 BrO + O2

Cycle 2 Cycle 3

Ozone Destruction Cycles

or ) (ClO)2

ClOO + Cl

C + O2

ClO + O2

3O2

(

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 this 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 gases react

and are reformed in each cycle. Sunlight is required to complete each cycle and to help form and maintain ClO abundances.

Ozone Destruction Cycles
TWENTY QUESTIONS: 2006 UPDATE

Q.19

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 upper 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 temperaQ10: 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” occurs because of the special weather conditions that exist there and nowhere else on the globe.

The very low 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.

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. Average minimum values over

Antarctica are as low as –90°C in July and

August in a typical year. Over the Arctic,

average minimum values are near –80°C in

January and February. Polar stratospheric

clouds (PSCs) are formed when winter minimum temperatures fall below the formation

temperature (about –78°C). This occurs on

average for 1 to 2 months over the Arctic and

5 to 6 months over Antarctica (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 (see Q9). The range of winter 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 present for many months, and

severe ozone depletion now occurs in each

winter season.

Minimum Air Temperatures in the

Polar Lower Stratosphere
TWENTY QUESTIONS: 2006 UPDATE

Q.20

tures 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.

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

meteorological differences between the hemispheres,

resulting from the differences 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 than the Arctic. Once chemical

changes occur in the cold 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 relative 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 (see Figure

Q10-1). As a result, PSCs are often found over large areas

of the winter polar regions and over a significant altitude

range. At low polar temperatures, nitric acid (HNO3) and

water condense on preexisting sulfur-containing particles

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 increase by early 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

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 and nitric acid (HNO3). 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 (see Q9).

Arctic Polar Stratospheric Clouds
TWENTY QUESTIONS: 2006 UPDATE

Q.21

highly reactive chlorine gas ClO remains chemically

active for a longer period, thereby increasing chemical

ozone destruction. Denitrification occurs each winter in

the Antarctic and in some, but not all, Arctic winters,

because PSC formation temperatures are required over an

extensive time period.

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 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 total ozone were observed in the early 1980s over research stations

located on the Antarctic continent. The measurements were made with ground-based Dobson spectrophotometers (see box in Q5). 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 three scientists from the British Antarctic Survey published them in the

journal Nature in 1985. Soon after, satellite measurements confirmed the spring ozone depletion and further showed that in each late winter/early 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

polar stratospheric cloud (PSC) formation; (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 satellite observations of total ozone,

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 is defined here as the area contained within

the 220-Dobson unit (DU) contour in total ozone maps

(light blue color in Figure Q11-1). The maximum area

has reached 25 million square kilometers (about 10 million square miles) in recent years, which is nearly twice

the area of the Antarctic continent (see Figure Q11-2).

Minimum values of total ozone inside the ozone hole averaged in late September have reached below 100 DU,

which is well below normal springtime values of about

200 DU (see Figure Q11-2). 

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. Balloon

TWENTY QUESTIONS: 2006 UPDATE

Q.22

III. STRATOSPHERIC OZONE DEPLETION

Q11: How severe is the depletion of the Antarctic ozone layer?

100 300 400 500

Total Ozone (Dobson units)

4 October 2001

200

Antarctic Ozone Hole

Figure Q11-1. Antarctic “ozone hole.” Total ozone

values are shown for high southern latitudes as measured by a satellite instrument. The dark blue and purple

regions over the Antarctic continent show the severe

ozone depletion or “ozone hole” now found during 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 200 DU (see Q4). In

late spring or early summer (November-December) the

ozone hole disappears in satellite images as ozonedepleted air is displaced and mixed with ozone-rich air

transported poleward 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 early spring (August-November). Peak depletion occurs in

early 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 far exceeds the

size of the Antarctic continent.

Antarctic Ozone Hole
TWENTY QUESTIONS: 2006 UPDATE

Q.23

measurements shown in Figure Q11-3 give an example of

such depletion 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 chlorine monoxide (ClO)

abundances. The average South Pole ozone profiles for

the decades 1962-1971 and 1992-2001 (see Figure Q11-

3) show how reactive halogen gases have dramatically

altered the ozone layer. For the 1960s, the ozone layer is

clearly evident in the October average profile and has a

peak near 16 kilometers. For 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, but significant spring ozone depletion has been observed every year only 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 how ozone depletion has changed with time in

both polar regions for the last three decades. Changes in

ozone hole areas and minimum Antarctic ozone amounts

are shown in Figure Q11-2. Depletion has increased since

1980 to become fairly stable in the 1990s and early 2000s,

with the exception of 2002 (see Q11-box). Total ozone

averaged over the Antarctic region in late winter/early

spring shows similar features (Figure Q12-1). Average

values decreased steadily through the 1980s and 1990s,

reaching minimum values that were 37% less than in preozone-hole years (1970-1982). The year-to-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 (see Figure Q11-box).

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/early

spring season. In spring, temperatures in the lower polar

stratosphere eventually warm, thereby ending PSC formation as well as 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 mixed 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.

Figure Q11-2. Antarctic ozone hole

features. Values are shown for key

parameters of the Antarctic ozone hole:

the area enclosed by the 220-DU total

ozone contour and the minimum total

ozone amount, as determined from

space-based observations. The values

are averaged for each year near the

peak of ozone depletion, as defined by

the dates shown in each panel. The

ozone hole areas are contrasted to the

areas of continents in the upper panel.

The intensity of ozone depletion gradually increased beginning in 1980. In the

1990s, the depletion reached fairly

steady values, except for the anomalously low depletion in 2002 (see Figure

Q11-box). The intensity of Antarctic

ozone depletion will decrease as part of

the ozone recovery process (see Q19

and Q20).

Antarctic Ozone Depletion
TWENTY QUESTIONS: 2006 UPDATE

Q.24

0 5 10 15 0 5 10 15

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) 10

15

Polar Ozone Depletion

South Pole (90°S) Finland (67°N)

Figure Q11-3. 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

with balloonborne instruments 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 “milli-Pascals”

(mPa), which is a measure of absolute pressure (100 million mPa = atmospheric sea-level pressure).

Polar Ozone Depletion
TWENTY QUESTIONS: 2006 UPDATE

Q.25

The Anomalous 2002 Antarctic Ozone Hole

The 2002 Antarctic ozone hole showed features that surprised scientists. They considered it anomalous at the time because the hole had much less area as viewed from space and much less ozone depletion as measured by minimum column ozone amounts when compared with values in several preceding

years (see Figure Q11-box). The 2002 ozone hole area and minimum ozone values stand out clearly in displays of the year-to-year changes in these quantities (see Figure Q11-2). The smaller area was unexpected because the conditions required to deplete ozone, namely low temperatures and available reactive

halogen gases, are not expected to have large year-to-year variations. Ozone was being depleted in August

and early September 2002, but the hole broke apart into two separate depleted regions during the last

week of September. The depletion in these two regions was significantly less than was observed inside

either the 2001 or 2003 ozone holes, but still substantially greater than was observed in the early 1980s.

The anomalous behavior in 2002 occurred because of specific atmospheric air motions that sometimes

occur in polar regions, not large decreases in reactive chlorine and bromine amounts in the Antarctic stratosphere. The Antarctic stratosphere was warmed by very strong, large-scale weather systems in 2002 that

originated in the lower atmosphere (troposphere) at midlatitudes in late September. In late September,

Antarctic temperatures are generally very low (see Q10) and ozone destruction rates are near their peak

values. These tropospheric systems traveled poleward and upward into the stratosphere, upsetting the circumpolar wind flow and warming the lower stratosphere where ozone depletion was ongoing. The higherthan-normal impact of these weather disturbances during the critical time period for ozone loss reduced the

total loss of ozone in 2002.

The warming in 2002 was unprecedented in Antarctic meteorological observations. Warming events

are difficult to predict because of their complex formation conditions. 

Large Antarctic ozone depletion returned in 2003 through 2005, in a manner similar to that observed

from the mid-1990s to 2001 (see Figures Q11-box and Q11-2). The high ozone depletion found since the

mid-1990s, with the exception of 2002, is expected to be typical of coming years. A significant, sustained

reduction of Antarctic ozone depletion, defined as ozone recovery, requires the removal of halogen source

gases from the stratosphere (see Q19 and Q20).

Figure Q11-Box. Anomalous 2002 ozone hole. Views from space of the Antarctic ozone hole as observed on 24

September in each of three consecutive years. The hole split and elongated in 2002, reducing the total depletion of

ozone observed that year in comparison with 2001 and 2003. The anomalous depletion in 2002 is attributable to an

early warming of the polar stratosphere caused by air disturbances originating in midlatitudes, rather than to large

changes in the amounts of reactive chlorine and bromine in the Antarctic stratosphere.

Antarctic Ozone Hole
TWENTY QUESTIONS: 2006 UPDATE

Q.26

Significant ozone depletion in the Arctic stratosphere

occurs in cold winters because of reactive halogen gases.

The depletion, however, is much less than the depletion

that now 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 a variety of instruments (see Q5), as for the

Antarctic (see Q11). These 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 balloonborne measurements 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 late

winter/early 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 and

ground-based 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 Antarctic

abundances (see Figure Q12-1). Decreases from the preozone-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 about 30% but have remained

smaller than those found in the Antarctic since the mid1980s. The year-to-year changes in the 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.

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/early spring

period (January-April). However, the maximum depletion is 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

well-calibrated satellite instruments. Shown

here is a comparison of average springtime

total ozone values found between 1970 and

1982 (solid and dashed red lines) 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 since 1990. The

ozone changes are the combination of

chemical destruction and natural variations.

Variations in meteorological conditions influence the year-to-year 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 naturally larger at the beginning of each winter season because more ozone is transported poleward each season in

the Northern Hemisphere than in the Southern Hemisphere.

Average Total Ozone in Polar Regions
TWENTY QUESTIONS: 2006 UPDATE

Q.27

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, which slows the conversion of reactive chlorine gases to form ClO and, as a consequence, reduces 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/early spring season. In

spring, temperatures in the lower stratosphere eventually

warm, thereby ending PSC formation as well as 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, total

ozone values in the Arctic are considerably higher than in

the Antarctic at the beginning of each winter season (see

Figure Q12-1).
TWENTY QUESTIONS: 2006 UPDATE

Q.28

Stratospheric ozone has decreased over the globe

since the 1980s. The depletion, which in the period

1997-2005 averaged about 4% (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 late winter/early

spring period. The next largest losses are observed in the

Northern Hemisphere, caused in part by late winter/early

spring losses over the Arctic. Ozone-depleted air over both

polar regions is dispersed 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 the presence of lower

amounts of reactive halogen gases (see Q8).

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

Figure Q13-1. Global total ozone changes. Satellite

observations show a decrease in global total ozone

values over more than two decades. The top panel

compares global ozone values (annual averages) with

the average from the period 1964 to 1980. Seasonal

and solar effects have been removed from the data. On

average, global ozone decreased each year between

1980 and the early 1990s. The decrease worsened

during the few years when volcanic aerosol from the

Mt. Pinatubo eruption in 1991 remained in the stratosphere. Now global ozone is about 4% below the 1964-

to-1980 average. The bottom panel compares ozone

changes between 1980 and 2004 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 Antarctic ozone

hole. Long-term changes in the tropics are much

smaller because reactive halogen gases are less 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 4% lower over the

globe. The average depletion exceeds the natural variability of the ozone layer. The ozone loss is very small

near the equator and increases with latitude toward the poles. The larger polar depletion is primarily a result

of the late winter/early spring ozone destruction that occurs there each year.

Global Total Ozone Change
TWENTY QUESTIONS: 2006 UPDATE

Q.29

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 very small. In contrast, stratospheric

air in polar regions has been in the stratosphere for an

average of 4 to 7 years; therefore, the abundance of reactive halogen gases is much larger. 

Seasonal changes. The magnitude of global ozone

depletion also depends on the season of the year. In comparison with the 1964-1980 averages, total ozone averaged for 2002-2005 is about 3% lower in northern middle

latitudes (35°N-60°N) and about 6% lower at southern

middle latitudes (35°S-60°S). The seasonality of these

changes is also somewhat different in the two hemispheres. In the summer/autumn periods, the decline in

total ozone is about 2% in the Northern Hemisphere and

5% in the Southern Hemisphere. In winter/spring, the

decline is about 5-6% in both hemispheres.
TWENTY QUESTIONS: 2006 UPDATE

Q.30

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). The

Sun’s output has gone through maximum values around

1969, 1980, 1991, and 2002. In 2006, the solar output

was decreasing towards a minimum.

Figure Q14-1. Solar changes and volcanoes. Total

ozone values have decreased beginning in the early

1980s (see middle panel). The ozone values shown are

3-month averages corrected for seasonal effects.

Incoming solar radiation, which produces ozone in the

stratosphere, changes on a well-recognized 11-year

cycle. The amount of solar radiation at a wavelength of

10.7-cm is often used to document the 11-year cycle (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 2005 period.

The largest recent eruptions are El Chichón (1982) and

Mt. Pinatubo (1991) (see red arrows in bottom panel).

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 only temporarily

because they 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.

The Solar Cycle, Global Ozone,

and Volcanic Eruptions
TWENTY QUESTIONS: 2006 UPDATE

Q.31

Over the last two decades, average total ozone has

decreased over the globe. Average values in recent years

show about 4% depletion from pre-1980 values (see

Figure Q14-1). The ozone values shown are 3-month

averages corrected for seasonal effects but not for solar

effects. 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 examinations of long-term ozone

changes presented in this and previous international scientific assessments quantitatively account for the influence of the 11-year solar cycle.

Past volcanoes. Large volcanic eruptions inject

sulfur gases directly into the stratosphere, causing new

sulfate particles to be formed. The particles initially form

in the stratosphere above and downwind of the volcano

location and then often spread throughout the hemisphere

or globally as air is transported by stratospheric winds.

The presence of volcanic particles in the stratosphere is

shown in observations of solar transmission through the

atmosphere. When large amounts of particles are present

in the stratosphere, transmission of solar radiation is significantly reduced. The large eruptions of El Chichón

(1982) and Mt. Pinatubo (1991) are recent examples of

events that temporarily reduced solar 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). The amount

of ClO produced is proportional to the total abundance

of reactive chlorine in the stratosphere (see Figure Q16-

1). 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 increased global ozone

depletion for a few years (see Figure Q14-1). After a few

years, however, the effect of volcanic particles on ozone

is diminished by their gradual removal from the stratosphere by natural air circulation. Because of particle

removal, the two 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 large number of particles

produced by the Mt. Pinatubo eruption, combined with

the relatively large amounts of reactive 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 abundances of reactive halogen gases in the stratosphere were smaller. In

the early decades of the 21st century, the abundance of

halogen source 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, ozone losses could be larger than previously observed and persist longer. Only later in the 21st

century when halogen gas abundances have declined close

to pre-1980 values will the effect of volcanic eruptions on

ozone be lessened. 
TWENTY QUESTIONS: 2006 UPDATE

Q.32

Montreal Protocol. In 1985, a treaty called the

Vienna 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, following country ratification, entered

into force 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 to

replace the principal halogen source gases, the Montreal

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 190 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 phased out before the

middle of the 21st century. 

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. Projections of the future abundance of effective

stratospheric chlorine (see Q16) 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 the

Protocol, stratospheric halogen gases are projected to

increase significantly in the 21st century. The “zero emissions” line shows a hypothetical case of stratospheric

abundances if all emissions were reduced to zero beginning in 2007. The lower panel shows how excess skin

cancer cases (see Q17) might increase with no regulation and how they might be reduced under the Protocol

provisions. (The unit “parts per trillion” is defined in the

caption of Figure Q7-1.)

Predicted Abundance (parts per trillion) 15000

10000

5000

0

Cases per million people per year

500

400

300

0

200

100

1980 2000 2020 2040 2060 2080 2100 Year

Effect of the Montreal Protocol
TWENTY QUESTIONS: 2006 UPDATE

Q.33

Protocol was strengthened with Amendments and

Adjustments. These revisions put additional substances

under regulation, accelerated existing control measures,

and prescribed phaseout dates for the production and consumption 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 and consumption 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. Future stratospheric

abundances of effective stratospheric chlorine (see Q16)

can be calculated based on the provisions of the Montreal

Protocol. The concept of effective stratospheric chlorine

accounts for the combined effect on ozone of chlorineand 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 and consumption as allowed

by the Protocol’s original provisions agreed upon

in Montreal in 1987.

• Restricted production and consumption as outlined

in the subsequent Amendments and Adjustments

as decided in London in 1990, Copenhagen in 1992,

and Beijing in 1999. 

• Zero emissions of ozone-depleting gases starting

in 2007. 

In each case, production of a gas is assumed to result

in its eventual emission to the atmosphere. Without the

Montreal Protocol and with continued production and use

of CFCs and other ozone-depleting gases, effective stratospheric chlorine is projected to have increased 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 also increased substantially at Earth’s surface, causing a rise in excess skin

cancer cases (see Q17 and lower panel of Figure Q15-1). 

The 1987 provisions of the Montreal Protocol alone

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 humanproduced ozone-depleting gases will ultimately be phased

out and effective stratospheric chlorine will slowly decay,

reaching pre-1980 values in the mid-21st

century (see Q16).

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. Examples of longterm storage are CFCs in refrigeration equipment and

foams, and halons in fire-fighting equipment. Some continued production and consumption of ozone-depleting

gases is allowed, particularly in developing nations, under

the 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 ozone-depleting gases

are set to zero beginning in 2007. The reductions in effective stratospheric chlorine below the values expected with

the 1999 Beijing agreement would be relatively small. 

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 88 to 98% less effective than CFC-12 in

depleting stratospheric ozone because they are chemically

removed primarily in the troposphere (see Q18). 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 chlorine abundance in the stratosphere, the

Montreal Protocol requires a gradual phaseout of HCFC

consumption in developed and developing nations that

will be complete in 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

(see Q18). As a consequence, the Montreal Protocol does

not regulate the HFCs. However, HFCs (as well as all

halogen source gases) are radiatively active gases that

contribute to human-induced 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 (UNFCCC).
TWENTY QUESTIONS: 2006 UPDATE

Q.34

Effective stratospheric chlorine. The Montreal

Protocol has been successful in slowing and reversing the

increase of ozone-depleting gases (halogen source gases)

in the atmosphere. An important measure of its success is

the change in the value of effective stratospheric chlorine.

Effective stratospheric chlorine values are a measure of

the potential for ozone depletion in the stratosphere,

obtained by summing over adjusted amounts of all chlorine and bromine gases. The adjustments account for the

different rates of decomposition of the gases and the

greater per-atom effectiveness of bromine in depleting

ozone (see Q7). Although chlorine is much more abundant in the stratosphere than bromine (160 times) (see

Figure Q7-1), bromine atoms are about 60 times more

effective than chlorine atoms in chemically destroying

ozone molecules. Increases in effective stratospheric

chlorine in the past decades have caused ozone depletion.

Accordingly, ozone is expected to recover in the future as

effective stratospheric chlorine values decrease.

Effective stratospheric chlorine changes. In the

latter half of the 20th century up until the 1990s, effective

stratospheric chlorine values steadily increased (see

Figure Q16-1). Values are derived from individual

halogen source gas abundances obtained from measurements, historical estimates of abundance, and projections

of future abundance. As a result of the Montreal Protocol

regulations, the long-term increase in effective stratospheric chlorine slowed, reached a peak, and began to

decrease in the 1990s. This initial 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. The decrease will continue

because, as emissions become small, natural destruction

processes gradually remove halogen-containing gases

from the global atmosphere. Reduction of effective stratospheric chlorine amounts to 1980 values or lower will

require many decades because the lifetimes of halogen

source gas molecules in the atmosphere range up to 100

years (see Figure Q16-1 and Table Q7-1). 

Individual halogen source gas reductions. The

reduction in the atmospheric abundance of a gas in

response to regulation depends on a number of factors that

include (1) how rapidly gas reserves are used and released

to the atmosphere, (2) the lifetime for the removal of the

gas from the atmosphere, and (3) the total amount of the

gas that has already accumulated in the atmosphere.

The regulation of human-produced 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 with 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.

Methyl chloroform and CFCs. The largest reduction in the abundance of a halogen source gas has

occurred for methyl chloroform (CH3CCl3) (see Figure

Q16-1). The implementation of the Montreal Protocol

caused global production of methyl chloroform to be

reduced to near zero. Atmospheric abundances subsequently dropped rapidly because methyl chloroform has

a short atmospheric lifetime (about 5 years). Methyl chloroform is used mainly as a solvent and has no significant

long-term storage following production. 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 amounts of these gases have all peaked, and CFC11 and CFC-113 abundances have decreased slightly (see

Figure Q16-1). As emissions of CFCs are reduced, their

atmospheric abundances will decrease more slowly than

methyl chloroform because of longer CFC atmospheric

lifetimes (see Table Q7-1) and because CFCs escape very

slowly to the atmosphere from their use in refrigeration

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. Around midcentury, the effective

abundance of ozone-depleting gases should fall to values that were present before the Antarctic “ozone

hole” began to form in the early 1980s.
TWENTY QUESTIONS: 2006 UPDATE

Q.35

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 left panel). Effective stratospheric chlorine values are a measure of the

potential for ozone depletion in the stratosphere, obtained by summing over adjusted amounts of all chlorine and bromine

gases. Effective stratospheric chlorine levels as shown here for midlatitudes will return to 1980 values around 2050. The

return to 1980 values will occur around 2065 in polar regions. In 1980, ozone was not significantly depleted by the chlorine and bromine then present in the stratosphere. A decrease in effective stratospheric chlorine abundance follows

reductions in emissions of individual halogen source gases. Overall emissions and atmospheric concentrations have

decreased and will continue to decrease given international compliance with the Montreal Protocol provisions (see Q15).

The changes in the atmospheric abundance of individual gases at Earth’s surface shown in the panels were obtained

using a combination of direct atmospheric measurements, estimates of historical abundance, and future projections of

abundance. The past increases of CFCs, along with those of CCl4 and CH3CCl3, have slowed significantly and most have

reversed in the last decade. HCFCs, which are used as CFC substitutes, will continue to increase in the coming decades.

Some halon abundances will also continue to grow in the future while current halon reserves are depleted. Smaller relative decreases are expected for CH3Br in response to production and use 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.)

Past and Expected Future Abundances of Atmospheric Halogen Source Gases
TWENTY QUESTIONS: 2006 UPDATE

Q.36

and foam products.

HCFC substitute gases. The Montreal Protocol

allows for the use of hydrochlorofluorocarbons (HCFCs)

as short-term 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 pose a lesser threat to the ozone layer than CFCs

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 consumption will

reach zero in developed nations by 2030 and in developing nations by 2040 (see Q15). 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.

Halons. The atmospheric abundances of halon-1211

and halon-1301 account for a significant fraction of

bromine from all 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 fire-extinguishing equipment and are

gradually being released, and production and consumption are still allowed in developing nations. Atmospheric

halon abundances can be expected to remain high well

into the 21st century because of their long lifetimes and

continued release.

Methyl chloride and methyl bromide. Both

methyl chloride (CH3Cl) and methyl bromide (CH3Br)

are distinct among principal halogen source gases

because a substantial fraction of their emissions is associated with natural processes (see Q7). The average

atmospheric abundance of methyl chloride, which is not

regulated under the Montreal Protocol, will remain fairly

constant throughout this century if natural sources remain

unchanged. 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

bromide, which is regulated under the Protocol, has

already decreased in recent years and is projected to

decrease further as a result of production phaseouts in

developed and developing countries. For the later decades

of the century, methyl bromide abundances are shown as

nearly constant in Figure Q16-1. However, these abundances are uncertain because the amounts of exempted

uses of methyl bromide under the Montreal Protocol are

not known for future years. 
TWENTY QUESTIONS: 2006 UPDATE

Q.37

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, other life forms, and materials (see Q3). The

amount of UV that produces an “erythemal” or sunburning

response in humans is often separately evaluated. Longterm changes in sunburning UV at a particular location

have been estimated from the changes in total ozone at

that location. The results show that average erythemal

UV has increased due to ozone reduction by up to a few

percent per decade between 1979 and 1998 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 erythemal

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 UV Index is essentially the

amount of erythemal irradiance as measured on a horizontal surface. The daily maximum UV Index varies with

location and season, as shown for three locations in

Figure Q17-1. Changes in surface UV radiation.

Ultraviolet (UV) radiation at Earth’s surface has increased

over much of the globe since 1979. Also known as “erythemal radiation,” sunburning UV is harmful to humans

and other life forms. The increases shown here for 1979-

1998 are estimated from observed decreases in ozone

and the relationship between ozone and surface UV

established at some surface locations. The estimates

are based on the assumption that all other factors that

influence the amount of UV radiation reaching the Earth’s

surface, such as aerosol abundances and cloudiness,

are unchanged. The estimated changes in ultraviolet

radiation in the tropics are the smallest because observed

ozone changes are the smallest there.

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.

Changes in Surface Ultraviolet Radiation
TWENTY QUESTIONS: 2006 UPDATE

Q.38

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 higher

latitude. At a given latitude, UV Index values increase in

mountainous regions. The UV Index becomes zero in

periods of continuous darkness found during winter 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, Antarctica (64°S),

sometimes equal or exceed even the peak summer

values measured in San Diego, California (32°N).

Other causes of long-term UV changes. The surface UV values 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. At some

ground sites, measurements indicate that long-term

changes in UV have resulted from changes in one or more

of these factors. The impact of some of the changes can

be complex. For example, an increase in cloud cover

usually results in a reduction of UV radiation below the

clouds, but can increase radiation above the clouds (in

mountainous regions).

UV changes and skin cancer. Skin cancer cases in

humans are expected to increase with the amount of 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 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 stratospheric chloWinter 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,

and with the Sun’s elevation in the local

sky. 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. The lowest average UV Index

values occur at high latitudes. As an

example, the figure compares the seasonal UV Index at three locations. The

UV Index is higher throughout the year

in San Diego, a low-latitude location, than

in Barrow, a high-latitude location. Index

values are zero at high latitudes in winter

when darkness is continuous. The effect

of Antarctic ozone depletion 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 decade 1991-2001, Antarctic ozone depletion has increased 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 and even in the summer in San Diego, which is located at much lower latitude.

Seasonal Changes in the UV Index
TWENTY QUESTIONS: 2006 UPDATE

Q.39

rine based on the 1992 and earlier Montreal Protocol

provisions and assuming that other factors (besides

ozone) affecting surface UV are unchanged in the future.

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),

with the assumption that the population’s sun exposure

remains unchanged. The case estimates include the fact

that skin cancer in humans occurs long after the exposure to sunburning UV. The results illustrate 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 have been expected to increase

substantially throughout the century.
TWENTY QUESTIONS: 2006 UPDATE

Q.40

Radiative forcing of climate change. Human activities and natural processes 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 in Earth’s atmosphere changes the balance 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. 

A summary of radiative forcings resulting from the

increases in long-lived greenhouse gases in the industrial

era is shown in Figure Q18-1. All forcings shown relate

to human activities. Positive forcings generally lead to

warming and negative forcings lead to cooling of Earth’s

surface. The accumulation of carbon dioxide represents

the largest forcing term. Carbon dioxide concentrations

are increasing in the atmosphere primarily as the result of

burning coal, oil, and natural gas for energy and transportation; and from cement manufacturing. The atmospheric abundance of carbon dioxide is currently about

35% above what it was 250 years ago, in preindustrial

times. In other international assessments, much of the

observed surface warming over the last 50 years has been

linked to increases in carbon dioxide and other greenhouse

gas concentrations caused by human activities. 

Stratospheric and tropospheric ozone. Stratospheric and tropospheric ozone both absorb infrared radiation emitted by Earth’s surface, effectively trapping heat

in the atmosphere. Stratospheric ozone also significantly

absorbs solar radiation. As a result, increases or decreases

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 absorbs solar

radiation and 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 both ozone-depleting gases and substitute gases contribute to climate change.

Figure Q18-1. Radiative forcing of climate

change from atmospheric gas changes.

Human activities since the start of the

Industrial Era (around 1750) have caused

increases in the abundances of several longlived 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.

Other international assessments have

shown that the largest radiative forcings

come from carbon dioxide, followed by

methane, tropospheric ozone, the halogencontaining 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.

Radiative Forcing

(watts per square meter)

warming cooling

Radiative Forcing of Climate Change from

Atmospheric Gas Changes (1750-2000)
TWENTY QUESTIONS: 2006 UPDATE

Q.41

in stratospheric or tropospheric ozone cause radiative forcings and represent direct links of ozone to climate change.

In recent decades, stratospheric ozone has decreased due

to rising chlorine and bromine amounts in the atmosphere,

while troposphere ozone in the industrial era has increased

due to pollution from human activities (see Q3).

Stratospheric ozone depletion causes a negative radiative

forcing, while increases in tropospheric ozone cause a

positive radiative forcing (see Figure Q18-1). The radiative forcing due to tropospheric ozone increases is currently larger than that associated with stratospheric ozone

depletion. The negative forcing from ozone depletion represents an offset to the positive forcing from the halogen

source gases, which cause ozone depletion. 

Halogen source gases and HFCs. An important link

between ozone depletion and climate change is the radiative forcing from halogen source gases and hydrofluorocarbons (HFCs). Halogen source gases are the cause of

ozone depletion (see Q7) and HFCs are substitute gases

(see Q15). Both groups of gases cause radiative forcing

in the atmosphere, but with a wide range of effectiveness.

The principal gases in each group are intercompared in

Figure Q18-2 (top panel) using their “ozone depletion

potentials” (ODPs) and “global warming potentials”

(GWPs), which indicate the effectiveness of each gas in

causing ozone depletion and climate change, respectively.

The ODPs of CFC-11 and CFC-12, and the GWP of CO2

all are assigned a value of 1.0. For ozone depletion, the

halons are the most effective gases (for equal mass

amounts) and HFCs cause no ozone depletion (see Q7).

For climate change, all gases make a contribution, with

CFC-12 and HFC-23 having the largest effect (for equal

mass amounts). Montreal Protocol actions (see Q15) that

have led to reductions in CFC concentrations and

increases in HCFC and HFC concentrations have also

reduced the total radiative forcing from these gases. It is

important to note that, despite a GWP that is small in comparison to many other greenhouse gases, CO2 is the most

important greenhouse gas related to human activities

because its atmospheric abundance is so much greater than

the abundance of other gases. 

The relative importance of total emissions of halogen

source gases and HFCs to ozone depletion and climate

change is illustrated for a single year (2004) of emissions

in the bottom panel of Figure Q18-2. The values displayed

are proportional to the product of 2004 annual global

emissions and the ODP or GWP. The results in the lower

panel are shown relative to CFC-11, because it is often

used as a reference gas. The comparison shows that the

importance of CFC emissions in 2004 to future ozone

depletion exceeds that of the halons, despite the higher

halon ODP values, because CFC emissions are larger.

Similarly, the contributions of CFC and HCFC-22 emissions in 2004 to climate change are currently larger than

the halon or HFC contributions. These 2004 results represent only incremental contributions of these gases to

either ozone depletion or climate change. The overall

contribution of a gas depends on its total accumulation

in the atmosphere, which in turn depends on its longterm emission history and atmospheric lifetime (see Q7

and Q16). In the case of ozone depletion, the relative

contributions of the halogen source gases can be compared through their respective contributions to effective

stratospheric chlorine (see Q16).

As a group, the principal halogen source gases represent a positive direct radiative forcing in the Industrial

Era that is comparable to the forcing from methane, the

second most important greenhouse gas. In the coming

decades, the abundances of these ozone-depleting gases

and their associated positive radiative forcings are

expected to decrease (see Q16). Future growth in HFC

emissions, while uncertain, will contribute a positive

forcing that will counter the decrease from ozonedepleting gases. Finally, reductions in ozone-depleting

gases will be followed by reductions in stratospheric

ozone depletion and its associated negative radiative

forcing. 

Impact of climate change on ozone. 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, lower temperatures and stronger polar winds

could both affect the extent and severity of winter polar

ozone depletion. While the Earth’s surface is expected to

warm in response to the net positive radiative forcing from

greenhouse gas increases, the stratosphere is expected to

cool. 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. In the upper stratosphere at altitudes

above PSC formation regions, a cooler stratosphere is

expected to increase ozone amounts and, hence, hasten

recovery, because lower temperatures favor ozone production over loss (see Q2). Similarly, changes in atmospheric composition that lead to a warmer climate may also

alter ozone amounts (see Q20).
TWENTY QUESTIONS: 2006 UPDATE

Q.42

Figure Q18-2. Evaluation of ozone-depleting gases and their substitutes. Ozone-depleting gases (halogen source

gases) and their substitutes can be compared via their ozone depletion potentials (ODPs) and global warming potentials

(GWPs). The GWPs are evaluated for a 100-yr time interval after emission. The CFCs, halons, and HCFCs are ozonedepleting gases (see Q7) and HFCs, used as substitute or replacement gases, do not destroy ozone. The ODPs of CFC11 and CFC-12, and the GWP of CO2 have values of 1.0 by definition. Larger ODPs or GWPs indicate greater potential

for ozone depletion or climate change, respectively. The top panel compares ODPs and GWPs for emissions of equal

mass amounts of each gas. The ODPs of the halons far exceed those of the CFCs. HFCs have zero ODPs. All gases

have non-zero GWPs that span a wide range of values. The bottom panel compares the contributions of the 2004

emissions of each gas, using CFC-11 as the reference gas. Each bar represents the product of a global emission

value and the respective ODP or GWP factor. The comparison shows that 2004 emissions of ozone-depleting gases

currently contribute more than substitute gas emissions to both ozone depletion and climate change. Future projections guided by Montreal Protocol provisions suggest that the contributions of ozone-depleting gases to climate change

will decrease, while those of the substitute gases will increase.

Evaluation of Selected Ozone-Depleting Substances and Substitute Gases
TWENTY QUESTIONS: 2006 UPDATE

Q.43

Recovery process. Identifying the recovery of the

ozone layer from depletion associated with halogen gases

will rely on comparisons of the latest ozone values with

values measured in the past. Because of its importance,

ozone will likely be measured continuously in the future

using a variety of techniques and measurement platforms

(see Q5). Atmospheric computer models will be used to

predict future abundances of ozone and attribute observed

changes to ozone-depleting gases and other factors. 

The recovery process is schematically shown for

global ozone in Figure Q19-1. Ozone has declined from

pre-1980 amounts due to past increases in halogen gases

in the stratosphere (see Q16). In the future, as the overall

decline in these gases continues in response to Montreal

Protocol provisions, global ozone is expected to recover,

approaching or exceeding pre-1980 values (see Q20).

Ozone recovery attributable to decreases in ozonedepleting gases can be described, in general, as a process

VI. STRATOSPHERIC OZONE IN THE FUTURE

Q19: How will recovery of the ozone layer be identified?

Scientists expect to identify the recovery of the ozone layer with detailed ozone measurements in the atmosphere and with global models of ozone amounts. Increases in global ozone and reductions 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. 

Figure Q19-1. Recovery stages of global ozone. Significant ozone depletion from the release of ozone-depleting

gases in human activities first became recognized in the 1980s. The Montreal Protocol provisions are expected to further

reduce and eliminate these gases in the atmosphere in the coming decades, thereby leading to the return of ozone

amounts to near pre-1980 values. The timeline of the recovery process is schematically illustrated with three stages

identified. The large uncertainty range illustrates natural ozone variability in the past and potential uncertainties in global

model projections of future ozone amounts. When ozone reaches the full recovery stage, global ozone values may be

above or below pre-1980 values, depending on other changes in the atmosphere (see Q20).

Recovery Stages of Global Ozone
TWENTY QUESTIONS: 2006 UPDATE

Q.44

involving three stages:

(1) The initial slowing of ozone decline, identified as

the occurrence of a statistically significant reduction in the rate of decline in ozone. 

(2) The onset of ozone increases (turnaround), identified as the occurrence of statistically significant

increases in ozone above previous minimum

values.

(3) The full recovery of ozone from ozone-depleting

gases, identified as when ozone is no longer significantly affected by ozone-depleting gases from

human activities. 

Each recovery stage is noted in Figure Q19-1. The red

line and shaded region in the figure indicate the expected

average value and the uncertainty range, respectively, in

global ozone amounts. The large uncertainty range illustrates natural ozone variability in the past and potential

uncertainties in global model projections of future ozone

amounts. 

In the full recovery of global ozone, the milestone of

the return of ozone to pre-1980 levels is considered important because prior to 1980 ozone was not significantly

affected by human activities. As a consequence, this milestone is useful, for example, to gauge when the adverse

impacts of enhanced surface ultraviolet (UV) radiation on

human health and ecosystems caused by ozone-depleting

substances are likely to become negligible. The uncertainty range in model results indicates that ozone amounts

may be below or above pre-1980 values when ozone has

fully recovered from the effects of ozone-depleting gases

from human activities (see Q20). The wide range of

uncertainty for global ozone in the final stage of recovery

represents, in part, the difficulty in accurately forecasting

the effects of future changes in climate and atmospheric

composition on the abundance of ozone (see Q20). 

Natural factors. Stratospheric ozone is influenced

by two important natural factors, namely, changes in the

output of the Sun and volcanic eruptions (see Q14).

Evaluations of ozone recovery 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. The uncertainty range in Figure Q19-1

includes solar changes. Volcanic eruptions are particularly important because they enhance 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 also

limits how easily small improvements in ozone abundances can be detected.
TWENTY QUESTIONS: 2006 UPDATE

Q.45

Q20: When is the ozone layer expected to recover?

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 in effective stratospheric

chlorine will follow the reductions in emissions that are

expected to continue 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 increases

in atmospheric abundances and the initial decline of several halogen gases have already been observed (see Figure

Q16-1). One gas, methyl chloroform, has already

decreased by about 90% from its peak value. Natural

chemical and transport processes limit the rate at which

halogen gases are removed from the stratosphere. The

atmospheric lifetimes of the halogen source gases range

up to 100 years (see Table Q7-1). Chlorofluorocarbon-12

(CFC-12), with its 100-year lifetime, will require about

200 to 300 years before it is removed (less than 5%

remaining) from the atmosphere (see Figure Q16-1). At

midlatitudes, effective stratospheric chlorine is not

expected to reach pre-1980 values until about 2050.

Ozone projections. Computer models of the atmosphere are used to assess past changes in the global ozone

distribution and to project 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 model projections

indicate that for 60°N-60°S total ozone, the first two

stages of recovery (slowing of the decline and turnaround

Figure Q20-1. Global ozone recovery predictions.

Observed values of midlatitude total ozone (top panel) and

September-October minimum total ozone values, over

Antarctica (bottom panel) have decreased beginning in the

early 1980s. As halogen source gas emissions decrease

in the 21st century, ozone values are expected to recover

by increasing toward pre-1980 values. Atmospheric computer models that account for changes in halogen gases

and other atmospheric parameters are used to predict how

ozone amounts will increase. These model results show

that full recovery is expected in midlatitudes by 2050, or

perhaps earlier. Recovery in the Antarctic will occur somewhat later. The range of model projections comes from the

use of several different models of the future atmosphere.

Substantial recovery of the ozone layer is expected near the middle of the 21st century, assuming global compliance with the Montreal Protocol. Recovery will occur as chlorine- and bromine-containing gases that

cause ozone depletion decrease in the coming decades under the provisions of the Protocol. However, the

influence of changes in climate and other atmospheric parameters could accelerate or delay ozone recovery,

and volcanic eruptions in the next decades could temporarily reduce ozone amounts for several years.

Recovery of Global Ozone
TWENTY QUESTIONS: 2006 UPDATE

Q.46

(see Q19)) will be reached before 2020. Full recovery,

with ozone reaching or exceeding pre-1980 values, is

expected to occur by the middle of the 21st century. The

range of projections comes from several computer models

of the atmosphere. Some of these models indicate that

recovery of 60°N-60°S total ozone may come well before

midcentury.

Models predict that Antarctic ozone depletion will

also reach the first two stages of recovery by 2020, but

somewhat more slowly than 60°N-60°S total ozone. Full

recovery could occur by mid-century but some models

show later recovery, between 2060 and 2070. Declines in

effective stratospheric chlorine amounts will occur later

over the Antarctic than at lower latitudes because air in

the Antarctic stratosphere is older than air found at lower

latitudes. As a result, reductions in halogen loading to

pre-1980 values will occur 10-15 years later in the

Antarctic stratosphere than in the midlatitude stratosphere.

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 significant ozone depletion (see Figure Q16-1).

However, climate and other atmospheric factors will not

be the same in 2050 as in 1980, and this could cause ozone

abundances in 2050 to be somewhat different from those

observed in 1980. Stratospheric ozone abundances are

affected by a number of natural and human-caused factors in addition to the atmospheric abundance of halogen

gases. Important examples are stratospheric temperatures

and air motions, volcanic eruptions, solar activity, and

changes in atmospheric composition. Separating the

effects of these factors is challenging because of the complexity of atmospheric processes affecting ozone. 

The ozone recovery projections in Figures Q19-1 and

Q20-1 attempt to take these various factors into account.

For example, since 1980 human activities have increased

the atmospheric abundance of important greenhouse

gases, including carbon dioxide, methane, and nitrous

oxide. Other international assessments have shown that

the accumulation of these gases is linked to the warmer

surface temperatures and lower stratospheric temperatures observed within recent decades. Warmer surface

temperatures could change the emission rates of naturally

occurring halogen source gases. Lower temperatures in

the upper stratosphere (at about 40 kilometers (25 miles)

altitude) accelerate ozone recovery because ozone

destruction reactions proceed at a slower rate. In contrast,

reduced temperatures in the polar lower stratosphere

during winter might increase the occurrence of polar

stratospheric clouds (PSCs) and, therefore, enhance chemical ozone destruction (see Q10). Further increases of

stratospheric water vapor, such as those that have occurred

over the last two decades, could also increase PSC occurrences and associated ozone destruction. Therefore, a

cooler, wetter polar stratosphere could delay polar ozone

recovery beyond what would be predicted for the 1980

atmosphere. Increased abundances of methane and

nitrous oxide due to human activities also cause some

change in the overall balance of the chemical production

and destruction of global stratospheric ozone. Finally,

one outcome that cannot be included precisely in models

is the occurrence of one or more large volcanic eruptions

in the coming decades. Large eruptions would increase

stratospheric sulfate particles for several years, temporarily reducing global ozone amounts (see Q14). 

As a consequence of these potential changes, the

return of effective stratospheric chlorine and ozone to pre1980 levels may not occur at the same time. In some

regions of the stratosphere, ozone may remain below pre1980 values after effective chlorine has declined to pre1980 levels.
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 (retired) USA

Robert T. Watson World Bank USA

SCIENTIFIC STEERING COMMITTEE

Marie-Lise Chanin Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

Susana B. Diaz Centro Austral de Investigaciones Cientificas Argentina

John A. Pyle University of Cambridge UK

A.R. Ravishankara NOAA Earth System Research Laboratory USA

Theodore G. Shepherd University of Toronto Canada

AUTHORS AND CONTRIBUTORS

CHAPTER 1

LONG-LIVED COMPOUNDS

Chapter Lead Authors

Cathy Clerbaux Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

Derek M. Cunnold Georgia Institute of Technology USA

Coauthors

John Anderson Hampton University USA

Andreas Engel Universität Frankfurt Germany

Paul J. Fraser CSIRO Division of Marine and Atmospheric Research Australia

Emmanuel Mahieu Université de Liège Belgium

Alistair Manning Climate Research Division, Met Office UK

John Miller NOAA Earth System Research Laboratory USA

Stephen A. Montzka NOAA Earth System Research Laboratory USA

Ray Nassar University of Waterloo Canada

Ronald Prinn Massachusetts Institute of Technology USA

Stefan Reimann Swiss Federal Laboratories for Materials Testing and Research Switzerland

A.1

Final Release: February 2007

From Scientific Assessment of Ozone Depletion: 2006
Curtis P. Rinsland NASA Langley Research Center USA

Peter Simmonds University of Bristol UK

Daniel P. Verdonik Hughes Associates, Inc. USA

Ray F. Weiss Scripps Institution of Oceanography USA

Donald J. Wuebbles University of Illinois USA

Yoko Yokouchi National Institute for Environmental Studies Japan

Contributors

Peter Bernath University of Waterloo Canada

Donald R. Blake University of California at Irvine USA

Thomas Blumenstock Forschungszentrum Karlsruhe Germany

James H. Butler NOAA Earth System Research Laboratory USA

Pierre-Francois Coheur Chimie Quantique et Photophysique, Université Libre de Bruxelles Belgium

Geoffrey S. Dutton CIRES/NOAA Earth System Research Laboratory USA

David Etheridge CSIRO Division of Marine and Atmospheric Research Australia

Lucien Froidevaux NASA Jet Propulsion Laboratory USA

Paul B. Krummel CSIRO Division of Marine and Atmospheric Research Australia

Jens Mühle Scripps Institution of Oceanography USA

Simon O’Doherty University of Bristol UK

David E. Oram University of East Anglia UK

Ellis E. Remsberg NASA Langley Research Center USA

Robert Rhew University of California at Berkeley USA

James M. Russell III Hampton University USA

Cathy Trudinger CSIRO Division of Marine and Atmospheric Research Australia

Darryn W. Waugh The Johns Hopkins University USA

Rudolphe Zander Université de Liège Belgium

CHAPTER 2

HALOGENATED VERY SHORT-LIVED SUBSTANCES

Chapter Lead Authors

Katherine S. Law Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

William T. Sturges University of East Anglia UK

Coauthors

Donald R. Blake University of California at Irvine USA

Nicola J. Blake University of California at Irvine USA

James B. Burkholder NOAA Earth System Research Laboratory USA

James H. Butler NOAA Earth System Research Laboratory USA

R. Anthony Cox University of Cambridge UK

Peter H. Haynes University of Cambridge UK

Malcolm K.W. Ko NASA Langley Research Center USA

Karin Kreher National Institute of Water and Atmospheric Research New Zealand

Céline Mari Laboratoire d’Aérologie, CNRS/Université Paul Sabatier France

Klaus Pfeilsticker Universität Heidelberg Germany

John M.C. Plane University of Leeds UK

Ross J. Salawitch Jet Propulsion Laboratory/California Institute of Technology USA

Cornelius Schiller Forschungszentrum Jülich Germany

Björn-Martin Sinnhuber University of Bremen Germany

AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.2
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.3

Roland von Glasow Universität Heidelberg Germany

Nicola J. Warwick University of Cambridge UK

Donald J. Wuebbles University of Illinois USA

Shari A. Yvon-Lewis Texas A&M University USA

Contributors

André Butz Universität Heidelberg Germany

David B. Considine NASA Langley Research Center USA

Marcel Dorf Universität Heidelberg Germany

Lucien Froidevaux NASA Jet Propulsion Laboratory USA

Laurie J. Kovalenko NASA Jet Propulsion Laboratory USA

Nathaniel J. Livesey NASA Jet Propulsion Laboratory USA

Ray Nassar University of Waterloo Canada

Christopher E. Sioris Harvard-Smithsonian Center for Astrophysics USA

Debra K. Weisenstein Atmospheric and Environmental Research, Inc. USA

CHAPTER 3

GLOBAL OZONE: PAST AND PRESENT

Chapter Lead Authors

Martyn P. Chipperfield University of Leeds UK

Vitali E. Fioletov Environment Canada Canada

Coauthors

Bram Bregman KNMI, now with TNO Business Unit Built and Geosciences The Netherlands

John P. Burrows University of Bremen Germany

Brian J. Connor National Institute of Water and Atmospheric Research New Zealand

Joanna D. Haigh Imperial College London UK

Neil R.P. Harris European Ozone Research Coordinating Unit/University of Cambridge UK

Alain Hauchecorne Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

Lon Hood University of Arizona USA

S. Randy Kawa NASA Goddard Space Flight Center USA

Janusz W. Krzy ́scin Polish Academy of Sciences Poland

Jennifer Logan Harvard University USA

Nzioka John Muthama University of Nairobi Kenya

Lorenzo M. Polvani Columbia University USA

William Randel National Center for Atmospheric Research USA

Toru Sasaki Japan Meteorological Agency Japan

Johannes Staehelin Swiss Federal Institute of Technology – Zurich Switzerland

Richard S. Stolarski NASA Goddard Space Flight Center USA

Larry W. Thomason NASA Langley Research Center USA

Joseph M. Zawodny NASA Langley Research Center USA

Contributors

Greg Bodeker National Institute of Water and Atmospheric Research New Zealand

Philippe Demoulin Université de Liège Belgium

Wuhu Feng University of Leeds UK

Lawrence E. Flynn NOAA National Environmental Satellite, Data, and Information Service USA

Stacey Hollandsworth Frith Science Systems and Applications, Inc. USA
Serge Guillas Georgia Institute of Technology USA

Mohammad Ilyas University College of Engineering North Malaysia Malaysia

Bjørn Knudsen Danish Meteorological Institute Denmark

J. Ben Liley National Institute of Water and Atmospheric Research New Zealand

Richard D. McPeters NASA Goddard Space Flight Center USA

Alvin J. Miller NOAA NWS Climate Prediction Center USA

Rolf Müller Forschungszentrum Jülich Germany

Samuel Oltmans NOAA Earth System Research Laboratory USA

Yvan J. Orsolini Norwegian Institute for Air Research Norway

Irina Petropavlovskikh CIRES/NOAA Earth System Research Laboratory USA

Wolfgang Steinbrecht German Weather Service, Hohenpeissenberg Germany

Hamish Struthers National Institute of Water and Atmospheric Research New Zealand

David W. Tarasick Environment Canada Canada

Yukio Terao Harvard University USA

CHAPTER 4

POLAR OZONE: PAST AND PRESENT

Chapter Lead Authors

Paul A. Newman NASA Goddard Space Flight Center USA

Markus Rex Alfred Wegener Institute for Polar and Marine Research, Potsdam Germany

Coauthors

Pablo O. Canziani Pontificia Universidad Católica Argentina/CONICET Argentina

Kenneth S. Carslaw University of Leeds UK

Katja Drdla NASA Ames Research Center USA

Sophie Godin-Beekmann Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

David M. Golden Stanford University and SRI International USA

Charles H. Jackman NASA Goddard Space Flight Center USA

Karin Kreher National Institute of Water and Atmospheric Research New Zealand

Ulrike Langematz Freie Universität Berlin Germany

Rolf Müller Forschungszentrum Jülich Germany

Hideaki Nakane National Institute for Environmental Studies Japan

Yvan J. Orsolini Norwegian Institute for Air Research Norway

Ross J. Salawitch Jet Propulsion Laboratory/California Institute of Technology USA

Michelle L. Santee Jet Propulsion Laboratory USA

Marc von Hobe Forschungszentrum Jülich Germany

Shigeo Yoden Kyoto University Japan

Contributors

Greg Bodeker National Institute of Water and Atmospheric Research New Zealand

Katja Frieler Alfred Wegener Institute for Polar and Marine Research, Potsdam Germany

Florence Goutail Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

Manuel López-Puertas Instituto de Astrofisica de Andalucia Spain

Gloria L. Manney NASA JPL and New Mexico Institute of Mining and Technology USA

Eric R. Nash Science Systems and Applications, Inc. USA

Cora E. Randall CU Laboratory for Atmospheric and Space Physics USA

Howard J. Singer NOAA National Weather Service USA

Cynthia Shaw Singleton CU Laboratory for Atmospheric and Space Physics USA

AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.4
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.5

Richard M. Stimpfle Harvard University USA

Simone Tilmes National Center for Atmospheric Research USA

Mark Weber University of Bremen Germany

CHAPTER 5

CLIMATE-OZONE CONNECTIONS

Chapter Lead Authors

Mark P. Baldwin NorthWest Research Associates, Inc. USA

Martin Dameris DLR Institut für Physik der Atmosphäre Germany

Coauthors

John Austin NOAA Geophysical Fluid Dynamics Laboratory USA

Slimane Bekki Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

Bram Bregman KNMI, now with TNO Business Unit Built and Geosciences The Netherlands

Neal Butchart Climate Research Division, Met Office UK

Eugene Cordero San Jose State University USA

Nathan P. Gillett University of East Anglia UK

Hans-F. Graf University of Cambridge UK

Claire Granier NOAA ESRL and Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) USA/France

Douglas E. Kinnison National Center for Atmospheric Research USA

Shyam Lal Physical Research Laboratory India

Thomas Peter Swiss Federal Institute of Technology – Zurich Switzerland

William Randel National Center for Atmospheric Research USA

John F. Scinocca Canadian Centre for Climate Modelling and Analysis, Univ. Victoria Canada

Drew Shindell NASA Goddard Institute for Space Studies USA

Hamish Struthers National Institute of Water and Atmospheric Research New Zealand

Masaaki Takahashi University of Tokyo Japan

David W.J. Thompson Colorado State University USA

Contributors

David Battisti University of Washington USA

Peter Braesicke University of Cambridge UK

Rolando R. Garcia National Center for Atmospheric Research USA

Peter H. Haynes University of Cambridge UK

Elisa Manzini Istituto Nazionale di Geofisica e Vulcanologia Italy

Katja Matthes National Center for Atmospheric Research USA

Giovanni Pitari Università degli Studi dell’Aquila Italy

V. Ramaswamy NOAA Geophysical Fluid Dynamics Laboratory USA

Karen Rosenlof NOAA Earth System Research Laboratory USA

Benjamin D. Santer Lawrence Livermore National Laboratory USA

Robert B. Scott University of Texas USA

Andrea Stenke DLR Institut für Physik der Atmosphäre Germany

Claudisa Timmreck Max-Planck-Institut für Meteorologie – Hamburg Germany
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.6

CHAPTER 6

THE OZONE LAYER IN THE 21ST CENTURY

Chapter Lead Authors

Greg Bodeker National Institute of Water and Atmospheric Research New Zealand

Darryn W. Waugh The Johns Hopkins University USA

Coauthors

Hideharu Akiyoshi National Institute for Environmental Studies Japan

Peter Braesicke University of Cambridge UK

Veronika Eyring DLR Institut für Physik der Atmosphäre Germany

David W. Fahey NOAA Earth System Research Laboratory USA

Elisa Manzini Istituto Nazionale di Geofisica e Vulcanologia Italy

Michael Newchurch University of Alabama in Huntsville USA

Robert W. Portmann NOAA Earth System Research Laboratory USA

Alan Robock Rutgers University USA

Keith P. Shine The University of Reading UK

Wolfgang Steinbrecht German Weather Service, Hohenpeissenberg Germany

Elizabeth C. Weatherhead CIRES/NOAA Earth System Research Laboratory USA

Contributors

John Austin NOAA Geophysical Fluid Dynamics Laboratory USA

Slimane Bekki Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

Christoph Brühl Max-Planck-Institut für Chemie – Mainz Germany

Neal Butchart Climate Research Division, Met Office UK

Martyn P. Chipperfield University of Leeds UK

Martin Dameris DLR Institut für Physik der Atmosphäre Germany

Tatiana Egorova World Radiation Center/Institute for Atmospheric and Switzerland

Climate Science ETH

Vitali E. Fioletov Environment Canada Canada

Andrew Gettelman National Center for Atmospheric Research USA

Marco A. Giorgetta Max-Planck-Institut für Meteorologie – Hamburg Germany

Douglas E. Kinnison National Center for Atmospheric Research USA

Eva Mancini Università degli Studi dell’Aquila Italy

Marion Marchand Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

Paul A. Newman NASA Goddard Space Flight Center USA

Steven Pawson NASA Goddard Space Flight Center USA

Giovanni Pitari Università degli Studi dell’Aquila Italy

David Plummer Environment Canada Canada

Bjørg Rognerud University of Oslo Norway

Eugene Rozanov World Radiation Center/Institute for Atmospheric and Switzerland

Climate Science ETH

Ross J. Salawitch Jet Propulsion Laboratory/California Institute of Technology USA

Theodore G. Shepherd University of Toronto Canada

Kiyotaka Shibata Meteorological Research Institute Japan

Björn-Martin Sinnhuber University of Bremen Germany

Miriam Sinnhuber University of Bremen Germany

Sergei P. Smyshlyaev Russian State Hydrometeorological University Russia

Richard S. Stolarski NASA Goddard Space Flight Center USA
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.7

Hamish Struthers National Institute of Water and Atmospheric Research New Zealand

Wenshou Tian University of Leeds UK

Guus J.M. Velders Netherlands Environmental Assessment Agency The Netherlands

Debra K. Weisenstein Atmospheric and Environmental Research, Inc. USA

Eun-Su Yang Georgia Institute of Technology USA

CHAPTER 7

SURFACE ULTRAVIOLET RADIATION: PAST, PRESENT, AND FUTURE

Chapter Lead Authors

Alkiviadis F. Bais Aristotle University of Thessaloniki Greece

Dan Lubin Scripps Institution of Oceanography USA

Coauthors

Antti Arola Finnish Meteorological Institute Finland

Germar Bernhard Biospherical Instruments Inc. USA

Mario Blumthaler Medical University of Innsbruck Austria

Natalia E. Chubarova Moscow State University Russia

Carynelisa Erlick The Hebrew University Israel

H. Peter Gies Australian Radiation Protection and Nuclear Safety Agency Australia

Nickolay A. Krotkov NASA Goddard Space Flight Center USA

Kathleen O. Lantz CIRES/NOAA Earth System Research Laboratory USA

Bernhard Mayer Deutsches Zentrum für Luft- und Raumfahrt Germany

Richard L. McKenzie National Institute of Water and Atmospheric Research New Zealand

Rubén D. Piacentini National University of Rosario Argentina

Gunther Seckmeyer University of Hannover Germany

James R. Slusser Colorado State University USA

Christos S. Zerefos National and Kapodestrian University of Athens Greece

Contributors

Uwe Feister Deutscher Wetterdienst Germany

Vitali E. Fioletov Environment Canada Canada

Julian Gröbner Physikalisch-Meteorologisches Observatorium Davos Switzerland

Esko Kyrö Finnish Meteorological Institute Finland

Harry Slaper National Institute for Public Health and the Environment The Netherlands

CHAPTER 8

HALOCARBON SCENARIOS, OZONE DEPLETION POTENTIALS,

AND GLOBAL WARMING POTENTIALS

Chapter Lead Authors

John S. Daniel NOAA Earth System Research Laboratory USA

Guus J.M. Velders Netherlands Environmental Assessment Agency The Netherlands

Coauthors

Anne R. Douglass NASA Goddard Space Flight Center USA

Piers M.D. Forster University of Leeds UK

Didier A. Hauglustaine CNRS – Laboratoire des Sciences du Climat et de l’Environnement France

Ivar Isaksen University of Oslo Norway
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.8

Lambert Kuijpers Technical University Eindhoven The Netherlands

Archie McCulloch University of Bristol UK

Timothy J. Wallington Ford Motor Company USA

Contributors

Paul Ashford Caleb Management Services Ltd. UK

Stephen A. Montzka NOAA Earth System Research Laboratory USA

Paul A. Newman NASA Goddard Space Flight Center USA

Darryn W. Waugh The Johns Hopkins University USA

TWENTY QUESTIONS AND ANSWERS ABOUT THE OZONE LAYER: 2006 UPDATE

Lead Author

David W. Fahey NOAA Earth System Research Laboratory USA

CHAPTER EDITORIAL CONTRIBUTORS

CHAPTER 1: LONG-LIVED COMPOUNDS

Maya George Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

CHAPTER 4: POLAR OZONE: PAST AND PRESENT

Rose M. Kendall Computer Sciences Corporation USA

Kathy A. Thompson Computer Sciences Corporation USA

CHAPTER 5: CLIMATE-OZONE CONNECTIONS

Amy Guenther NorthWest Research Associates USA

CHAPTER 6: THE OZONE LAYER IN THE 21ST CENTURY

Hisako Shiona National Institute of Water and Atmospheric Research New Zealand

TWENTY QUESTIONS AND ANSWERS ABOUT THE OZONE LAYER: 2006 UPDATE

Brenda S. Irish NOAA Earth System Research Laboratory USA

Debra Dailey-Fisher NOAA Earth System Research Laboratory USA

REVIEWERS

Ayité-Lô Nohende Ajavon Université de Lomé Togo

Daniel L. Albritton NOAA Earth System Research Laboratory (Retired) USA

Stephen O. Andersen U.S. Environmental Protection Agency USA

Gustavo A. Argüello Universidad Nacional de Cordoba Argentina

Elliot Atlas University of Miami USA

Pieter J. Aucamp Ptersa Environmental Consultants South Africa

John Austin NOAA Geophysical Fluid Dynamics Laboratory USA
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.9

Gnon Baba Université de Lomé Togo

Alkiviadis F. Bais Aristotle University of Thessaloniki Greece

Mark P. Baldwin NorthWest Research Associates, Inc. USA

Leonard A. Barrie World Meteorological Organization Switzerland

Gufran Beig Indian Institute of Tropical Meteorology India

Slimane Bekki Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

Lars Olof Björn Lund University Sweden

Greg Bodeker National Institute of Water and Atmospheric Research New Zealand

Rumen D. Bojkov Dresden Technical University Germany

Janet Bornman University of Waikato New Zealand

Geir O. Braathen World Meteorological Organization Switzerland

Claus Brüning European Commission Belgium

John P. Burrows University of Bremen Germany

Neal Butchart Climate Research Division, Met Office UK

James H. Butler NOAA Earth System Research Laboratory USA

James M. Calm Engineering Consultant USA

Pablo O. Canziani Pontificia Universidad Católica Argentina/CONICET Argentina

Daniel Cariolle Météo-France France

Lucy Carpenter University of York UK

Marie-Lise Chanin Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

Martyn P. Chipperfield University of Leeds UK

Hans Claude German Meteorological Service, Hohenpeissenberg Germany

Cathy Clerbaux Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

William J. Collins Hadley Centre for Climate Prediction and Research, Met Office UK

David B. Considine NASA Langley Research Center USA

John Crowley Max-Planck-Institut für Chemie – Mainz Germany

Paul Crutzen Max-Planck-Institut für Chemie – Mainz Germany

Derek M. Cunnold Georgia Institute of Technology USA

Martin Dameris DLR Institut für Physik der Atmosphäre Germany

John S. Daniel NOAA Earth System Research Laboratory USA

Michael Danilin The Boeing Company USA

Susana B. Diaz Centro Austral de Investigaciones Cientificas Argentina

Anne R. Douglass NASA Goddard Space Flight Center USA

Ellsworth G. Dutton NOAA Earth System Research Laboratory USA

Kalju Eerme Tartu Observatory Estonia

Christine A. Ennis CIRES/NOAA Earth System Research Laboratory USA

Wayne Evans NorthWest Research Associates, Inc. USA

Veronika Eyring DLR Institut für Physik der Atmosphäre Germany

David W. Fahey NOAA Earth System Research Laboratory USA

Vitali E. Fioletov Environment Canada Canada

Horst Fischer Max Planck Institut für Chemie – Mainz Germany

Lawrence E. Flynn NOAA National Environmental Satellite, Data, and Information Service USA

Ian Folkins Dalhousie University Canada

John Frederick University of Chicago USA

Lucien Froidevaux NASA Jet Propulsion Laboratory USA

Jan Fuglestvedt University of Oslo Norway

Marvin A. Geller State University of New York, Stony Brook USA

Andrew Gettelman National Center for Atmospheric Research USA

Nathan P. Gillett University of East Anglia UK

Sophie Godin-Beekmann Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.10

Marco González United Nations Environment Programme Kenya

Florence Goutail Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

Hans-F. Graf University of Cambridge UK

Claire Granier NOAA ESRL and Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) USA/France

Neil R.P. Harris European Ozone Research Coordinating Unit/University of Cambridge UK

Dennis Hartmann University of Washington USA

Sachiko Hayashida Naru Women’s University Japan

Peter H. Haynes University of Cambridge UK

Jay Herman NASA Goddard Space Flight Center USA

David J. Hofmann NOAA Earth System Research Laboratory USA

Dale Hurst CIRES/NOAA Earth System Research Laboratory USA

Abdelmoneim A.R. Ibrahim Egyptian Meteorological Authority Egypt

Mohammad Ilyas University College of Engineering North Malaysia Malaysia

Takashi Imamura National Institute for Environmental Studies Japan

Ivar Isaksen University of Oslo Norway

Elena Jimenez Universidad de Castilla – La Mancha Spain

Paul Johnston National Institute of Water and Atmospheric Research New Zealand

Roderic L. Jones University of Cambridge UK

Hiroshi Kanzawa Nagoya University Japan

David Karoly University of Oklahoma USA

Jussi Kaurola Finnish Meteorological Institute Finland

Jack A. Kaye NASA Headquarters USA

M.A.K. Khalil Portland State University USA

Bjørn Knudsen Danish Meteorological Institute Denmark

Malcolm K.W. Ko NASA Langley Research Center USA

Chester Koblinsky NOAA Climate Program Office USA

Peter Koepke University of Munich Germany

Janusz W. Krzy ́scin Polish Academy of Sciences Poland

Lambert Kuijpers Technical University Eindhoven The Netherlands

Michael J. Kurylo NASA Headquarters USA

Katherine S. Law Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

Mark G. Lawrence Max-Planck-Institut für Chemie – Mainz Germany

Zenobia Litynska Institute of Meteorology and Water Management Poland

Jennifer Logan Harvard University USA

Dan Lubin Scripps Institution of Oceanography USA

Gloria L. Manney NASA JPL / New Mexico Institute of Mining and Technology USA

Elisa Manzini Istituto Nazionale di Geofisica e Vulcanologia Italy

Mack McFarland DuPont Fluoroproducts USA

Richard L. McKenzie National Institute of Water and Atmospheric Research New Zealand

Richard D. McPeters NASA Goddard Space Flight Center USA

Davit Melkonyan Armenian State Hydrometeorological and Monitoring Service Armenia

Pauline Midgley M&D Consulting Germany

Alvin J. Miller NOAA NWS Climate Prediction Center USA

Kenneth R. Minschwaner New Mexico Institute of Mining and Technology USA

Stephen A. Montzka NOAA Earth System Research Laboratory USA

Rolf Müller Forschungszentrum Jülich Germany

Nzioka John Muthama University of Nairobi Kenya

Hideaki Nakane National Institute for Environmental Studies Japan

Eric R. Nash Science Systems and Applications, Inc. USA
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.11

Paul A. Newman NASA Goddard Space Flight Center USA

Ole John Nielsen University of Copenhagen Denmark

Alan O’Neill University of Reading UK

Samuel Oltmans NOAA Earth System Research Laboratory USA

Eduardo Palenque Universidad Mayor de San Andrés Bolivia

Panos Papagiannakopoulos University of Crete Greece

Thomas Peter Swiss Federal Institute of Technology – Zurich Switzerland

Ulrich Platt Universität Heidelberg Germany

R. Alan Plumb Massachusetts Institute of Technology USA

Ian Plumb CSIRO Industrial Physics Australia

Lamont R. Poole Science Applications International Corporation USA

Robert W. Portmann NOAA Earth System Research Laboratory USA

Gilles Poulet Université d’Orléans France

Michael Prather University of California at Irvine USA

John A. Pyle University of Cambridge UK

S. Ramachandran Physical Research Laboratory India

V. Ramaswamy NOAA Geophysical Fluid Dynamics Laboratory USA

William Randel National Center for Atmospheric Research USA

A.R. Ravishankara NOAA Earth System Research Laboratory USA

Thomas Reddmann Forschungszentrum Karlsruhe Germany

Claire Reeves University of East Anglia UK

Stefan Reimann Swiss Federal Laboratories for Materials Testing and Switzerland

Research, EMPA

Markus Rex Alfred Wegener Institute for Polar and Marine Research, Potsdam Germany

Alan Robock Rutgers University USA

José M. Rodríguez NASA Goddard Space Flight Center USA

Geert-Jan Roelofs University of Utrecht The Netherlands

Howard Roscoe British Antarctic Survey UK

Karen Rosenlof NOAA Earth System Research Laboratory USA

Ross J. Salawitch Jet Propulsion Laboratory/California Institute of Technology USA

Michelle L. Santee Jet Propulsion Laboratory USA

Benjamin D. Santer Lawrence Livermore National Laboratory USA

Yasuhiro Sasano National Institute for Environmental Studies Japan

Robert Sausen DLR Institut für Physik der Atmosphäre Germany

Ulrich Schmidt Johann Wolfgang Goethe Universität Germany

Robyn Schofield CIRES/NOAA Earth System Research Laboratory USA

Ulrich Schumann DLR Institut für Physik der Atmosphäre Germany

Gunther Seckmeyer University of Hannover Germany

Megumi Seki United Nations Environment Programme Kenya

Theodore G. Shepherd University of Toronto Canada

Kiyotaka Shibata Meteorological Research Institute Japan

Keith P. Shine The University of Reading UK

Emily Shuckburgh University of Cambridge UK

Isobel Simpson University of California at Irvine USA

Pavla Skrivankova Czech Hydrometeorological Institute Czech Republic

Susan Solomon NOAA Earth System Research Laboratory USA

Johannes Staehelin Swiss Federal Institute of Technology – Zurich Switzerland

Knut Stamnes Stevens Institute of Technology USA

Richard M. Stimpfle Harvard University USA

Richard S. Stolarski NASA Goddard Space Flight Center USA
A.12

Frode Stordal University of Oslo Norway

Susan Strahan NASA Goddard Space Flight Center USA

Hamish Struthers National Institute of Water and Atmospheric Research New Zealand

William T. Sturges University of East Anglia UK

Azadeh Tabazadeh Stanford University USA

Darin Toohey University of Colorado USA

Jan C. van der Leun Ecofys The Netherlands

Costas Varotsos University of Athens Greece

Guus J.M. Velders Netherlands Environmental Assessment Agency The Netherlands

Mario Visca Ausimont, S.p.A. Italy

John Michael Wallace University of Washington USA

Darryn W. Waugh The Johns Hopkins University USA

Elizabeth C. Weatherhead CIRES/NOAA Earth System Research Laboratory USA

Ann Webb University of Manchester UK

Philipp Weihs Universität für Bodenkultur Wein Austria

Ray F. Weiss Scripps Institution of Oceanography USA

Paul Wennberg California Institute of Technology USA

Paul Wine Georgia Institute of Technology USA

Stephen Wood National Institute of Water and Atmospheric Research New Zealand

Donald J. Wuebbles University of Illinois USA

Eun-Su Yang Georgia Institute of Technology USA

Margarita Yela Instituto Nacional de Técnica Aeroespacial Spain

Shigeo Yoden Kyoto University Japan

Vladimir Yushkov Central Aerological Observatory Russia

Rudolphe Zander Université de Liège Belgium

Francesco Zaratti Universidad Mayor de San Andrés Bolivia

Christos S. Zerefos National and Kapodestrian University of Athens Greece

Tong Zhu Peking University China

OZONE PEER-REVIEW MEETING

Les Diablerets, Switzerland

19-23 June 2006

Ayité-Lô Nohende Ajavon Université de Lomé Togo

Stephen O. Andersen U.S. Environmental Protection Agency USA

John Austin NOAA Geophysical Fluid Dynamics Laboratory USA

Alkiviadis F. Bais Aristotle University of Thessaloniki Greece

Mark P. Baldwin NorthWest Research Associates, Inc. USA

Gufran Beig Indian Institute of Tropical Meteorology India

Lars Olof Björn Lund University Sweden

Greg Bodeker National Institute of Water and Atmospheric Research New Zealand

Geir O. Braathen World Meteorological Organization Switzerland

John P. Burrows University of Bremen Germany

Marie-Lise Chanin Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

Martyn P. Chipperfield University of Leeds UK

Cathy Clerbaux Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

AUTHORS, CONTRIBUTORS, AND REVIEWERS
A.13

Derek M. Cunnold Georgia Institute of Technology USA

Martin Dameris DLR Institut für Physik der Atmosphäre Germany

John S. Daniel NOAA Earth System Research Laboratory USA

Susana B. Diaz Centro Austral de Investigaciones Cientificas Argentina

Kalju Eerme Tartu Observatory Estonia

Christine A. Ennis CIRES/NOAA Earth System Research Laboratory USA

David W. Fahey NOAA Earth System Research Laboratory USA

Vitali E. Fioletov Environment Canada Canada

Ian Folkins Dalhousie University Canada

Marvin A. Geller State University of New York, Stony Brook USA

Nathan P. Gillett University of East Anglia UK

Marco González United Nations Environment Programme Kenya

Florence Goutail Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

Neil R.P. Harris European Ozone Research Coordinating Unit/University of Cambridge UK

Peter H. Haynes University of Cambridge UK

David J. Hofmann NOAA Earth System Research Laboratory USA

Abdelmoneim A.R. Ibrahim Egyptian Meteorological Authority Egypt

M.A.K. Khalil Portland State University USA

Malcolm K.W. Ko NASA Langley Research Center USA

Chester Koblinsky NOAA Climate Program Office USA

Lambert Kuijpers Technical University Eindhoven The Netherlands

Michael J. Kurylo NASA Headquarters USA

Katherine S. Law Service d’Aéronomie / IPSL (CNRS-UPMC-UVSQ) France

Elisa Manzini Istituto Nazionale di Geofisica e Vulcanologia Italy

Mack McFarland DuPont Fluoroproducts USA

Richard L. McKenzie National Institute of Water and Atmospheric Research New Zealand

Stephen A. Montzka NOAA Earth System Research Laboratory USA

Rolf Müller Forschungszentrum Jülich Germany

Nzioka John Muthama University of Nairobi Kenya

Paul A. Newman NASA Goddard Space Flight Center USA

Alan O’Neill University of Reading UK

Thomas Peter Swiss Federal Institute of Technology – Zurich Switzerland

R. Alan Plumb Massachusetts Institute of Technology USA

Lamont R. Poole Science Applications International Corporation USA

Robert W. Portmann NOAA Earth System Research Laboratory USA

John A. Pyle University of Cambridge UK

S. Ramachandran Physical Research Laboratory India

V. Ramaswamy NOAA Geophysical Fluid Dynamics Laboratory USA

A.R. Ravishankara NOAA Earth System Research Laboratory USA

Stefan Reimann Swiss Federal Laboratories for Materials Testing and Switzerland

Research, EMPA

Markus Rex Alfred Wegener Institute for Polar and Marine Research, Potsdam Germany

Alan Robock Rutgers University USA

José M. Rodríguez NASA Goddard Space Flight Center USA

Ross J. Salawitch Jet Propulsion Laboratory/California Institute of Technology USA

Ulrich Schmidt Johann Wolfgang Goethe Universität Germany

Gunther Seckmeyer University of Hannover Germany

Megumi Seki United Nations Environment Programme Kenya

Theodore G. Shepherd University of Toronto Canada

Kiyotaka Shibata Meteorological Research Institute Japan
AUTHORS, CONTRIBUTORS, AND REVIEWERS

A.14

Susan Solomon NOAA Earth System Research Laboratory USA

Johannes Staehelin Swiss Federal Institute of Technology – Zurich Switzerland

Knut Stamnes Stevens Institute of Technology USA

Richard S. Stolarski NASA Goddard Space Flight Center USA

William T. Sturges University of East Anglia UK

Jan C. van der Leun Ecofys The Netherlands

Guus J.M. Velders Netherlands Environmental Assessment Agency The Netherlands

Darryn W. Waugh The Johns Hopkins University USA

Elizabeth C. Weatherhead CIRES/NOAA Earth System Research Laboratory USA

Ray F. Weiss Scripps Institution of Oceanography USA

Donald J. Wuebbles University of Illinois USA

Vladimir Yushkov Central Aerological Observatory Russia

Francesco Zaratti Universidad Mayor de San Andrés Bolivia

Christos S. Zerefos National and Kapodestrian University of Athens Greece

Tong Zhu Peking University China
AUTHORS, CONTRIBUTORS, AND REVIEWERS

Sponsoring Organizations Liaisons

Geir O. Braathen World Meteorological Organization Switzerland

Marco González United Nations Environment Programme Kenya

Daniel L. Albritton National Oceanic and Atmospheric Administration (retired) USA

Michael J. Kurylo National Aeronautics and Space Administration USA

Claus Brüning European Commission Belgium

Coordinating Editor

Christine A. Ennis CIRES/NOAA Earth System Research Laboratory USA

Technical Editing

Christine A. Ennis CIRES/NOAA Earth System Research Laboratory USA

Reference Research and Editing

Heather L. McCullough NOAA Boulder Laboratories Library USA

Publication/Graphics Design and Layout

Debra Dailey-Fisher NOAA Earth System Research Laboratory USA

Dennis Dickerson (Graphics Design, “Twenty Questions”) Respond Grafiks USA

Editorial Assistance

Karen M. Layman NOAA Earth System Research Laboratory USA

Debra Dailey-Fisher NOAA Earth System Research Laboratory USA

Conference Coordination and Documentation

Geir O. Braathen World Meteorological Organization Switzerland

Chantal Renaudot World Meteorological Organization Switzerland

Catherine Michaut CNRS/Institut Pierre-Simon Laplace France

Christine A. Ennis CIRES/NOAA Earth System Research Laboratory USA

Jeanne S. Waters NOAA Earth System Research Laboratory USA

Kathy A. Thompson Computer Sciences Corporation USA

Conference Support

Jeanne S. Waters NOAA Earth System Research Laboratory USA

Kathy A. Thompson Computer Sciences Corporation USA

Catherine A. Burgdorf CIRES/NOAA Earth System Research Laboratory USA

Debra Dailey-Fisher NOAA Earth System Research Laboratory USA

Marie-Claude Mermillod World Meteorological Organization Switzerland

Rose M. Kendall Computer Sciences Corporation USA

Laurent Bourdette Institut Pierre-Simon Laplace France

Karen M. Layman NOAA Earth System Research Laboratory USA

Melinda Tignor IPCC Working Group I Technical Support Unit USA

Kyle Terran University Corporation for Atmospheric Research USA

Computing and Networking Support

Catherine A. Burgdorf CIRES/NOAA Earth System Research Laboratory USA

Richard J. Tisinai CIRES/NOAA Earth System Research Laboratory USA

Vladislava Toukalek World Meteorological Organization Switzerland

Jalil Housni World Meteorological Organization Switzerland

Document Distribution

Jeanne S. Waters NOAA Earth System Research Laboratory USA

A.15

B.1

APPENDIX B

MAJOR ACRONYMS AND ABBREVIATIONS

A1 baseline halocarbon scenario (Chapter 8)

AAO Antarctic oscillation

ACE-FTS Atmospheric Chemistry Experiment Fourier Transform Spectrometer

AER Atmospheric and Environmental Research, Inc. (United States)

AERONET Aerosol Robotic Network

AFEAS Alternative Fluorocarbons Environmental Acceptability Study

AFGL Air Force Geophysics Laboratory (United States)

AGAGE Advanced Global Atmospheric Gases Experiment

AGCM Atmospheric General Circulation Model

AGWP Absolute Global Warming Potential

AIDA Aerosol Interactions and Dynamics in the Atmosphere

amsl above mean sea level

AO Arctic Oscillation

AOD aerosol optical depth

AOGCM coupled ocean-atmosphere general circulation model

ARPANSA Australian Radiation Protection and Nuclear Safety Agency (Australia)

ASAP Assessment of Stratospheric Aerosol Properties (SPARC)

ATLAS Atmospheric Laboratory for Applications and Science

atm atmosphere (unit of pressure)

ATMOS Atmospheric Trace Molecule Spectroscopy

AVHRR Advanced Very High Resolution Radiometer

B0 full capture and destruction of the 2007 bank cases (Chapter 8)

BAU business as usual (scenario)

BDC Brewer-Dobson circulation

BfS Bundesamt fur Strahlenschutz (Germany)

BL boundary layer

BOM Bureau of Meteorology (Australia)

BUV Backscatter (or Backscattered) Ultraviolet (spectrometer)

CCD charge-coupled device

CCM Chemistry-Climate Model

CCMVal CCM Validation Activity

CCN cloud condensation nuclei

CCSP Climate Change Science Program (United States)

CCSR Center for Climate System Research (University of Tokyo)

CDOM chromophoric dissolved organic matter

CERES Clouds and the Earth’s Radiant Energy System

CFC chlorofluorocarbon

CIE Commission Internationale de l’Éclairage (France)

CIRES Cooperative Institute for Research in Environmental Sciences (United States)

CLaMS Chemical Lagrangian Model of the Stratosphere

cm centimeters (unit of length)

CMAM Canadian Middle Atmosphere Model
ACRONYMS

B.2

CMDL Climate Monitoring and Diagnostics Laboratory (NOAA)

CMF cloud modification factor

CNRM Centre National de Recherches Météorologiques (France)

CNRS Centre National de la Recherche Scientifique (France)

CONICET Consejo de Investigaciones Cientificas y Técnicas (Argentina)

CPT cold point tropopause

CSIRO Commonwealth Scientific and Industrial Research Organisation (Australia)

CTM chemical transport model

CU University of Colorado (United States)

CUSUM cumulative sum of residuals

DHM dihalomethane

DIAL Differential Absorption Lidar

DJF December-January-February

DLR Deutschen Zentrum für Luft- und Raumfahrt (Germany)

DMS dimethyl sulfide

DNA deoxyribonucleic acid

DOAS Differential Optical Absorption Spectroscopy

DOC dissolved organic carbon

DOM dissolved organic matter

DU Dobson unit

DWD Deutscher Wetterdienst (Germany)

E0 “no future emission” case (Chapter 8)

EC European Commission

ECMWF European Centre for Medium-Range Weather Forecasts (United Kingdom)

EDUCE European Database for UV Climatology and Evaluation

EECl effective equivalent chlorine

EESC equivalent effective stratospheric chlorine

ENSO El Niño-Southern Oscillation

ENVISAT Environmental Satellite

EOS Earth Observing System

EP Eliassen-Palm

EPA Environmental Protection Agency (United States)

ERA ECMWF Re-Analysis

ER-2 Earth Resources-2 (aircraft)

ERS-2 European Remote Sensing-2 (satellite)

ESA European Space Agency

ESRL Earth System Research Laboratory (NOAA)

ETH Eidgenössische Technische Hochschule (Swiss Federal Institute of Technology) (Switzerland)

EUVDB European Ultraviolet Database

ExTL extratropical tropopause layer

F1 future control simulation (Chapter 6)

FASTRT Fast and Easy Radiative Transfer

FC fluorocarbon

FMI Finnish Meteorological Institute (Finland)

FTIR Fourier transform infrared

FTS Fourier Transform Spectrometer

FUB Freie Universität Berlin (Germany)
ACRONYMS

g gram (unit of mass)

GC gas chromatograph

GCM general circulation model

GCR galactic cosmic ray

GEOS Goddard Earth Observing System

GEWEX Global Energy and Water Cycle Experiment (WCRP)

GFDL Geophysical Fluid Dynamics Laboratory (NOAA)

Gg Gigagrams (109 grams) (unit of mass)

GHG greenhouse gas

GHz Gigahertz (109 cycles per second) (frequency unit)

GMAO Global Monitoring and Assimilation Office (NASA)

GMD Global Monitoring Division (NOAA/ESRL)

GMI Global Modeling Initiative (CTM)

GOME Global Ozone Monitoring Experiment

GSFC Goddard Space Flight Center (NASA)

GtC gigatons carbon

GWP Global Warming Potential

H1 no-Protocol halocarbon scenario (Chapter 8)

HALOE Halogen Occultation Experiment

HBFC hydrobromofluorocarbon

HCFC hydrochlorofluorocarbon

HFC hydrofluorocarbon

HFE hydrofluorinated ether or hydrofluoroether

hPa hectoPascal (102 Pascal) (unit of pressure)

HTOC Halons Technical Options Committee (TEAP)

IASB Institut d’Aéronomie Spatiale de Belgique (Belgium)

IFU Institute for Atmospheric Environmental Research (Germany)

ILAS Improved Limb Atmospheric Spectrometer

IMG Interferometric Monitor for Greenhouse Gases

INSPECTRO Influence of Clouds on the Spectral Actinic Flux in the Lower Troposphere

IPCC Intergovernmental Panel on Climate Change

IPSL Institut Pierre-Simon Laplace (France)

IR infrared

ISCCP International Satellite Cloud Climatology Project

IUPAC International Union of Pure and Applied Chemistry

JJA June-July-August

JMA Japan Meteorological Agency (Japan)

JPL Jet Propulsion Laboratory (NASA)

JRC Joint Research Centre (Italy)

K Kelvin (unit of temperature)

kg kilogram (103 grams) (unit of mass)

km kilometer (103 meters) (unit of length)

KNMI Koninklijk Nederlands Meteorologisch Instituut (The Netherlands)

kton kilotons (103 tons) (unit of mass)

l liter (unit of volume)

LaRC Langley Research Center (NASA)

B.3
ACRONYMS

LMS lowermost stratosphere

LOPES long path extinction spectrometer 

LPDM Lagrangian particle dispersion model

LPMA Limb Profile Monitor of the Atmosphere

LS lower stratosphere

LT Lagrangian transport

m meter (unit of length)

M Molar (moles per liter) (unit of concentration)

MAESTRO Measurement of Aerosol Extinction in the Stratosphere and Troposphere

Retrieved by Occultation

MAM March-April-May

MBL marine boundary layer

MBTOC Methyl Bromide Technical Options Committee

mg milligram (10-3 grams) (unit of mass)

MIPAS Michelson Interferometer for Passive Atmospheric Sounding 

MLS Microwave Limb Sounder

mm millimeters (10-3 meters) (unit of length)

MNP Netherlands Environmental Assessment Agency (The Netherlands)

mol mole (unit, amount of substance)

molec molecule

MOZAIC Measurement of Ozone and Water Vapor by Airbus In-Service Aircraft

MRI Meteorological Research Institute (Japan)

MSC Meteorological Service of Canada (Canada)

MSG METEOSAT Second Generation

MSU Microwave Sounding Unit

MVIRI METEOSAT Visible and Infrared Instrument

NAM Nothern Hemisphere 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)

NCC no climate change simulation (Chapter 6)

NCEP National Centers for Environmental Prediction (United States)

NDACC Network for the Detection of Atmospheric Composition Change

NDSC Network for the Detection of Stratospheric Change (now NDACC)

NH Northern Hemisphere

NIES National Institute for Environmental Studies (Japan)

NILU Norwegian Institute for Air Research (Norway)

NIST National Institute of Standards and Technology (formerly NBS, United States)

NIWA National Institute of Water and Atmospheric Research (New Zealand)

nm nanometers (10-9 meters) (unit of length)

NMI National Metrology Institutes

NOAA National Oceanic and Atmospheric Administration (United States)

n-PB n-propyl bromide

NSF National Science Foundation (United States)

NWP numerical weather prediction

NWS National Weather Service (NOAA)

B.4
ACRONYMS

ODP Ozone Depletion Potential

ODS ozone-depleting substance

OHA ozone hole area

OMD ozone mass deficit

OMI Ozone Monitoring Instrument

OPAC Optical Properties of Aerosols and Clouds

P0 “no future production” case (Chapter 8)

P1 past trends simulation (Chapter 6)

PAR photosynthetically active radiation

PFC perfluorocarbon

PG product gas

PGI product gas injection

pmol picomole (10-12 mole) (unit, amount of substance)

POAM Polar Ozone and Aerosol Measurement

POC particulate organic carbon

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

PSS photochemical steady state

PTFE polytetrafluoroethylene

PV potential vorticity

PVU potential vorticity unit

QA/QC quality assurance/quality control

QASUME Quality Assurance of Spectral Solar UV Measurements in Europe

QBO quasi-biennial oscillation

QPS quarantine and pre-shipment

RAMI Radiation Transfer Model Intercomparison

RATPAC Radiosonde Atmospheric Temperature Products for Assessing Climate

REF reference simulation (Chapter 6)

REPROBUS Reactive Processes Ruling the Ozone Budget in the Stratosphere

RF radiative forcing

RIVM Rijksinstituut voor Volksgezondheid en Milieu (National Institute of Public Health

and the Environment) (The Netherlands)

RT radiative transfer

SAD surface area density

SAGE Stratospheric Aerosol and Gas Experiment

SAM Southern Hemisphere annular mode

SAOZ Système d’Analyse par Observation Zénithale

SBDART Santa Barbara Discrete-ordinate Hemispheric Radiative Transfer

SBUV/SBUV2 Solar Backscatter (or Backscattered) Ultraviolet (spectrometer)

SCIAMACHY Scanning Imaging Absorption Spectrometer for Atmospheric Chartography

SCOUT-O3 Stratospheric-Climate Links with Emphasis on the Upper Troposphere and

Lower Stratosphere

B.5
ACRONYMS

SCTM Stratospheric Chemistry Transport Model

SeaWiFS Sea-viewing Wide Field-of-view Sensor

SG source gas

SGI source gas injection

SH Southern Hemisphere

SHADOZ Southern Hemisphere Additional Ozonesondes

SIDES Surface Irradiation Derived from a Range of Satellite-Based Sensors

SLIMCAT Single-Layer Isentropic Model of Chemistry and Transport

SMR Submillimeter Radiometer or Submillimeter Receiver

SOGE System for Observation of Halogenated Greenhouse Gases in Europe

SOLSTICE Solar Stellar Irradiance Comparison Experiment

SON September-October-November

SPARC Stratospheric Processes and Their Role in Climate (WCRP)

SPE solar proton event

SPURT Spurenstofftransport in der Tropopausenregion

SRES Special Report on Emissions Scenarios (IPCC)

SSA stratospheric sulfate aerosol

SST sea surface temperature

SSU Stratospheric Sounding Unit

STACCATO Influence of Stratosphee-Troposphere Exchange in a Changing Climate on Atmosheric 

Transport and Oxidation Capacity

STAR System for Transfer of Atmospheric Radiation

STREAM Stratosphere-Troposphere Experiments by Aircraft Measurements

STT stratosphere-to-troposphere transport

SUNY State University of New York (United States)

SUSIM Solar Ultraviolet Spectral Irradiance Monitor

SYNOP Surface Synoptic Observations

SZA solar zenith angle

2-D two-dimensional

3-D three-dimensional

TEAP Technology and Economic Assessment Panel (UNEP)

TEMIS Tropospheric Emission Monitoring Internet Service

THM trihalomethane

TOMS Total Ozone Mapping Spectrometer

TOVS TIROS Operational Vertical Sounder

TRACE-P Transport and Chemical Evolution over the Pacific 

TROICA Trans-Siberian Observations Into the Chemistry of the Atmosphere

TST troposphere-to-stratosphere transport

TTL tropical tropopause layer

TUV Tropospheric Ultraviolet Visible

UARS Upper Atmosphere Research Satellite

UEA University of East Anglia (United Kingdom)

UH University of Heidelberg (Germany)

UKMO United Kingdom Met Office (United Kingdom)

UM University of Manchester (United Kingdom)

UNEP United Nations Environment Programme

UPMC Université Pierre et Marie Curie (France)

UT upper troposphere

B.6
ACRONYMS

UTLS upper troposphere/lower stratosphere

UV ultraviolet

UVMFRSR ultraviolet multifilter rotating shadowband radiometer

UVSQ Université Versailles Saint-Quentin (France)

VAR variational analysis

VSL very short-lived

VSLS very short-lived substances

W Watts (unit of energy)

WCRP World Climate Research Programme

WHO World Health Organization

WMGHG well-mixed greenhouse gas

WMO World Meteorological Organization

WOUDC World Ozone and Ultraviolet Radiation Data Centre (Canada)

B.7

APPENDIX C

MAJOR CHEMICAL FORMULAE AND NOMENCLATURE

FROM THIS ASSESSMENT

HALOGEN-CONTAINING SPECIES

Cl atomic chlorine Br atomic bromine

Clx inorganic chlorine oxides/radicals Brx inorganic bromine oxides/radicals

Cly total inorganic chlorine Bry total inorganic bromine

CCly organic chlorine CBry organic bromine

Cl2 molecular chlorine Br2 molecular bromine

ClO chlorine monoxide BrO bromine monoxide

Cl2O dichlorine monoxide Br2O dibromine monoxide

ClOx chlorine radicals BrOx bromine radicals

OClO chlorine dioxide OBrO bromine dioxide

ClOO chloroperoxy radical

Cl2O2, ClOOCl dichlorine peroxide (ClO dimer)

ClONO2, ClNO3 chlorine nitrate BrONO2, BrNO3 bromine nitrate

HCl hydrogen chloride (hydrochloric acid) HBr hydrogen bromide

HOCl hypochlorous acid HOBr hypobromous acid

BrCl bromine chloride

F atomic fluorine I atomic iodine

F2 molecular fluorine I2 molecular iodine

Ix inorganic iodine oxides/radicals

Fy total inorganic fluorine Iy total inorganic iodine

IO iodine monoxide

IOx iodine radicals

OIO iodine dioxide

IOIO iodine oxide dimer

IONO2 iodine nitrate

HF hydrogen fluoride (hydrofluoric acid) HI hydrogen iodide

HOI hypoiodous acid

SF6 sulfur hexafluoride

SF5CF3 trifluoromethylsulfurpentafluoride

NF3 nitrogen trifluoride

C.1
CHEMICAL FORMULAE

C.2

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 CH3CF2CCl2F

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 (BFC-12B2) CBr2F2 Halon-1301 (BFC-13B1) CBrF3

Halon-1211 (BCFC-12B1) CBrClF2 Halon-2402 (BFC-114B2) CBrF2CBrF2

CHLOROCARBONS BROMOCARBONS

CH3Cl (HCC-40) methyl chloride, chloromethane CH3Br (HBC-40B1) methyl bromide, bromomethane

CH2Cl2 (HCC-30) methylene chloride, dichloromethane CH2Br2 methylene bromide, dibromomethane

CHCl3 (HCC-20) chloroform, trichloromethane CHBr3 bromoform, tribromomethane

CCl4 (CC-10) carbon tetrachloride C2H5Br ethyl bromide, bromoethane

C2HCl3 trichloroethene, trichloroethylene CH2BrCH2Br 1,2 dibromoethane

C2Cl4 tetrachloroethene, perchloroethene CH3CH2CH2Br, 1-bromopropane,

CH3CH2Cl chloroethane, ethyl chloride (n-C3H7Br) n-propyl bromide, n-PB

CH2ClCH2Cl 1,2 dichloroethane COHBr formyl bromide

CH3CCl3 (HCC-140a) methyl chloroform COBr2, Br2C(O) carbonyl dibromide

C2H2Cl4 tetrachloroethane

CH3CH2CH2Cl n-propyl chloride

COCl2, Cl2C(O) phosgene, carbonyl chloride
C.3

IODOCARBONS FLUOROCARBONS

CH3I iodomethane, methyl iodide CF4 (FC-14) perfluoromethane, 

CH2I2 diiodomethane carbon tetrafluoride

CH3CH2I, C2H5I iodoethane, ethyl iodide C2F6, CF3CF3 (FC-116) perfluoroethane

CH3CHICH3 2-iodopropane, isopropyl iodide C3F8, CF3CF2CF3 (FC-218) perfluoropropane

(i-C3H7I) c-C3F6 (FC-C216) perfluorocyclopropane

CH3CH2CH2I 1-iodopropane, n-propyl iodide C4F10 (FC-31-10) perfluorobutane

(n-C3H7I) c-C4F8 (FC-C318) perfluorocyclobutane

C5F12 (FC-41-12) perfluoropentane

C6F14 (FC-51-14) perfluorohexane

C10F18 perfluorodecalin

COF2 carbonyl fluoride

OTHERS

CHBr2Cl dibromochloromethane CF3I (FIC-1311) trifluoroiodomethane, 

CH2BrCl bromochloromethane trifluoromethyl iodide

CHBrCl2 bromodichloromethane COClF chlorofluorocarbonyl

CH2BrI bromoiodomethane SF6 sulfur hexafluoride

CHBrF2 bromodifluoromethane SF5CF3 trifluoromethylsulfurpentafluoride

CH2ClI chloroiodomethane

OTHER CHEMICAL 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(1

D), O3) or H2O2 hydrogen peroxide

oxidant (O3 + NO2) HOx odd hydrogen (H, OH, HO2, H2O2)

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 RONO2 alkyl nitrates

NO3 nitrogen trioxide, nitrate radical PAN peroxyacetylnitrate (CH3C(O)OONO2)

N2O5 dinitrogen pentoxide CH3CN acetonitrile

NOx nitrogen oxides (NO + NO2)

NOy total reactive nitrogen 

(usually includes NO, NO2, NO3,

N2O5, ClONO2, HNO4, HNO3)

CHEMICAL FORMULAE
CHEMICAL FORMULAE

C.4

S atomic sulfur H2S hydrogen sulfide

SO2 sulfur dioxide CS2 carbon disulfide

H2SO4 sulfuric acid COS, OCS carbonyl sulfide

CH3SCH3 DMS, dimethyl sulfide

C carbon atom

CO carbon monoxide

CO2 carbon dioxide

NMHC nonmethane hydrocarbon CH2O, HCHO formaldehyde

CH4 methane CH3CHO acetaldehyde

C2H6 ethane CH3OH methyl alcohol, methanol

C3H8 propane CH3O2 methyl peroxy radical

C2H4 ethylene, ethene RO alkoxy radicals

C2H2 acetylene, ethyne RO2 organic peroxy radical