World Meteorological Organization

Global Ozone Research and Monitoring Project—Report No. 52

Scientific Assessment of

Ozone Depletion: 2010

Pursuant to Article 6 of the Montreal Protocol

on Substances that Deplete the Ozone Layer

National Oceanic and Atmospheric Administration

National Aeronautics and Space Administration

United Nations Environment Programme

World Meteorological Organization

European Commission

iii

Patricio Aceituno Chile

Ayité-Lô Nohende Ajavon Togo

Hideharu Akiyoshi Japan

Daniel L. Albritton USA

Taofiki Aminou Benin

Stephen O. Andersen USA

Julie M. Arblaster Australia/USA

Antti Arola Finland

Ghassem Asrar WMO

Elliot Atlas USA

Pieter J. Aucamp South Africa

John Austin USA

Alkiviadis F. Bais Greece

Mark P. Baldwin USA

Dimitris Balis Greece

Gufran Beig India

Slimane Bekki France

Peter F. Bernath UK

Donald R. Blake USA

Thomas Blumenstock Germany

Gregory E. Bodeker New Zealand

Rumen D. Bojkov Germany

Janet F. Bornman New Zealand

Geir O. Braathen WMO

Peter Braesicke UK

Christoph Brühl Germany

Claus Brüning Belgium

Dominik Brunner Switzerland

James B. Burkholder USA

John P. Burrows Germany/UK

Neal Butchart UK

James H. Butler USA

André Butz The Netherlands

Timothy Canty USA

Pablo O. Canziani Argentina

Bruno Carli Italy

List of International Authors,

Contributors, and Reviewers

Assessment Cochairs

Ayité-Lô Nohende Ajavon

Paul A. Newman

John A. Pyle

A.R. Ravishankara

Chapters and Coordinating Lead Authors

Chapter 1: Ozone-Depleting Substances (ODSs) and Related Chemicals (Stephen A. Montzka and Stefan Reimann)

Chapter 2: Stratospheric Ozone and Surface Ultraviolet Radiation (Anne Douglass and Vitali Fioletov)

Chapter 3: Future Ozone and Its Impact on Surface UV (Slimane Bekki and Gregory E. Bodeker)

Chapter 4: Stratospheric Changes and Climate (Piers M. Forster and David W.J. Thompson)

Chapter 5: A Focus on Information and Options for Policymakers (John S. Daniel and Guus J.M. Velders)

Twenty Questions and Answers About the Ozone Layer: 2010 Update (David W. Fahey and Michaela I. Hegglin)

Scientific Review and Advisory Group

Malcolm K.W. Ko

Theodore G. Shepherd

Susan Solomon

Coordinating Editor

Christine A. Ennis

Authors, Contributors, and Reviewers
iv

Authors, Contributors, and Reviewers

Lucy Carpenter UK

Ken Carslaw UK

Marie-Lise Chanin France

Andrew J. Charlton-Perez UK

Martyn P. Chipperfield UK

Natalia E. Chubarova Russia

Irene Cionni Germany

Hans Claude Germany

Cathy Clerbaux France

Gerrie Coetzee South Africa

William J. Collins UK

Brian J. Connor New Zealand

Raul Cordero Chile

Eugene C. Cordero USA

Derek M. Cunnold USA

Martin Dameris Germany

John S. Daniel USA

Christine David France

Hugo De Backer Belgium

Martine De Mazière Belgium

Philippe Demoulin Belgium

Peter N. den Outer The Netherlands

Dick Derwent UK

Panuganti Devara India

Sandip Dhomse UK

Roseanne Diab South Africa

Susana B. Diaz Argentina

Marcel Dorf Germany

Anne R. Douglass USA

Pierre Duchatelet Belgium

Geoffrey S. Dutton USA

Ellsworth G. Dutton USA

Kalju Eerme Estonia

James William Elkins USA

Andreas Engel Germany

Christine A. Ennis USA

Veronika Eyring Germany

David W. Fahey USA

Uwe Feister Germany

Vitali E. Fioletov Canada

Eric L. Fleming USA

Lawrence E. Flynn USA

Ian Folkins Canada

Piers M. Forster UK

James Franklin Belgium

Paul J. Fraser Australia

Melissa P. Free USA

Stacey M. Frith USA

Lucien Froidevaux USA

John C. Fyfe Canada

Annie Gabriel Australia

Rolando R. Garcia USA

Hella Garny Germany

Marvin A. Geller USA

Andrew Gettelman USA

Manuel Gil Spain

Nathan P. Gillett Canada

Marco A. Giorgetta Germany

Sophie Godin-Beekmann France

Marco González UNEP

Hans-F Graf UK

Lesley Gray UK

Kevin M. Grise USA

Jens-Uwe Grooß Germany

Joanna D. Haigh UK

Ebrahim Hajizadeh Iran

Steven C. Hardiman UK

Neil R.P. Harris UK

Dennis L. Hartmann USA

Frank Hase Germany

Birgit Hassler USA

Michaela I. Hegglin Canada

François Hendrick Belgium

Jay R. Herman USA

Ernest Hilsenrath USA

David J. Hofmann USA

Paul Horwitz UNEP

Petra Huck New Zealand

Robert D. Hudson USA

Mohammad Ilyas Malaysia

Takashi Imamura Japan

Ivar S.A. Isaksen Norway

Charles H. Jackman USA

Serm Janjai Thailand

Imre M. Jánosi Hungary

Patrick Jöckel Germany

Andreas I. Jonsson Canada

Kenneth Jucks USA

David J. Karoly Australia

Andreas Kazantzidis Greece

Philippe Keckhut France

Douglas E. Kinnison USA

Jon Klyft Sweden

Malcolm K.W. Ko USA

Kunihiko Kodera Japan

Takashi Koide Japan

Ninong Komala Indonesia

Yutaka Kondo Japan

Karin Kreher New Zealand

Mark Kroon The Netherlands

Kirstin Krüger Germany

Paul B. Krummel Australia

Janusz W. Krzyścin Poland

Anne Kubin Germany

Lambert Kuijpers The Netherlands

Michael J. Kurylo USA

Paul J. Kushner Canada

Esko Kyrö Finland
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Authors, Contributors, and Reviewers

Shyam Lal India

Jean-François Lamarque USA

Tom Land USA

Ulrike Langematz Germany

Igor Larin Russia

Katharine Law France

Franck Lefèvre France

Jos Lelieveld Germany

Yi Liu China

Jennifer Logan USA

Diego Loyola Germany

Cathrine Lund Myhre Norway

Sasha Madronich USA

Emmanuel Mahieu Belgium

Eva Mancini Italy

Gloria L. Manney USA

Alistair J. Manning UK

Elisa Manzini Germany

Marion Marchand France

Daniel R. Marsh USA

Katja Matthes Germany

Bernhard Mayer Germany

John C. McConnell Canada

C. Thomas McElroy Canada

Mack McFarland USA

Norman McFarlane Canada

Danny McKenna USA

Richard L. McKenzie New Zealand

Charles McLandress Canada

Chris A. McLinden Canada

Inna A. Megretskaia USA

Abdelwahid Mellouki France

Martine Michou France

Pauline M. Midgley Switzerland

John Miller USA

Mario J. Molina USA/Mexico

Stephen A. Montzka USA

Olaf Morgenstern New Zealand

Jens Mühle USA

Rolf Müller Germany

Nzioka John Muthama Kenya

Prijitha J. Nair France

Hideaki Nakane Japan

Eric R. Nash USA

Cindy Newberg USA

Mike Newchurch USA

Paul A. Newman USA

Ole John Nielsen Denmark

Simon O’Doherty UK

Alan O’Neill UK

Samuel J. Oltmans USA

Luke D. Oman USA

Vladimir L. Orkin USA

Mathias Palm Germany

Dimitrios Papanastasiou USA

Edward A. Parson USA

Nigel D. Paul UK

Steven Pawson USA

Stuart A. Penkett UK

Judith Perlwitz USA

Thomas Peter Switzerland

Irina Petropavlovskikh USA

Klaus Pfeilsticker Germany

Giovanni Pitari Italy

Michael Pitts USA

R. Alan Plumb USA

David Plummer Canada

Jean-Pierre Pommereau France

Michael Ponater Germany

Lamont R. Poole USA

Robert W. Portmann USA

Michael J. Prather USA

Ronald G. Prinn USA

John A. Pyle UK

Birgit Quack Germany

S. Ramachandran India

V. Ramaswamy USA

William J. Randel USA

T. Narayan Rao India

A.R. Ravishankara USA

Claire E. Reeves UK

Stefan Reimann Switzerland

Markus Rex Germany

Robert Rhew USA

Martin Riese Germany

Vincenzo Rizi Italy

Alan Robock USA

Howard K. Roscoe UK

Karen H. Rosenlof USA

Martin N. Ross USA

Eugene Rozanov Switzerland

Vladimir Ryabinin WMO

David Saint-Martin France

Ross J. Salawitch USA

Michelle L. Santee USA

K. Madhava Sarma India

Robert Sausen Germany

Adam A. Scaife UK

Sue Schauffler USA

Ulrich Schmidt Germany

Matthias Schneider Germany

Robyn Schofield Germany

Ulrich Schumann Germany

John F. Scinocca Canada

Dian J. Seidel USA

Megumi Seki UNEP

Jonathan Shanklin UK

Wafik M. Sharobiem Egypt
vi

Authors, Contributors, and Reviewers

Theodore G. Shepherd Canada

Kiyotaka Shibata Japan

Keith P. Shine UK

Masato Shiotani Japan

Michael Sigmond Canada

Peter Simmonds UK

Isobel J. Simpson USA

Björn-Martin Sinnhuber Germany

Harry Slaper The Netherlands

Dan Smale New Zealand

Anne Smith USA

Susan Solomon USA

Seok-Woo Son Canada

Johannes Staehelin Switzerland

Wolfgang Steinbrecht Germany

Georgiy L. Stenchikov Saudi Arabia

David Stevenson UK

Andreas Stohl Norway

Richard S. Stolarski USA

Frode Stordal Norway

Susan Strahan USA

Fred Stroh Germany

William T. Sturges UK

Kenshi Takahashi Japan

David W. Tarasick Canada

Susann Tegtmeier Germany

Yukio Terao Japan

Hubert Teyssèdre France

Said Ali Thaoubane Comores

David W.J. Thompson USA

Simone Tilmes USA

Darin W. Toohey USA

Kleareti Tourpali Greece

Matthew B. Tully Australia

Jéssica Valverde-Canossa Costa Rica

Ronald Van der A The Netherlands

Karel Vanicek Czech Republic

Guus J.M. Velders The Netherlands

Daniel P. Verdonik USA

Corinne Vigouroux Belgium

Martin K. Vollmer Switzerland

Marc von Hobe Germany

Dmitry I. Vyushin Canada

Timothy J. Wallington USA

Hsiang J. (Ray) Wang USA

Darryn W. Waugh USA

Elizabeth C. Weatherhead USA

Ann R. Webb UK

Mark Weber Germany

Ray F. Weiss USA

Donald J. Wuebbles USA

Masaaki Yamabe Japan

Eun-Su Yang USA

Shigeo Yoden Japan

Yoko Yokouchi Japan

Shari A. Yvon-Lewis USA

Durwood Zaelke USA

Rodolphe Zander Belgium

Christos S. Zerefos Greece

Lingxi Zhou China
vii

Remembrances

It is with sadness that we note the passing of the following scientists who have played leading roles in the international scientific assessments of the ozone layer.

Derek Cunnold (1940–2009). Derek Martin Cunnold was born July 10, 1940, in Reading, England. 

He received his B.A. and M.A. from St. John’s College in Cambridge, England, and his Ph.D. in 

Electrical Engineering from Cornell University in 1965. He was a Professor Emeritus at the Georgia 

Institute of Technology at the time of his death. He was an author and/or contributor in all of the 

Ozone Assessments since 1988, and was a Lead Author of Chapter 1 (“Long-Lived Compounds”) of 

the 2006 Assessment.

David Hofmann (1937–2009). David J. Hofmann was born January 3, 1937. He received his 

Ph.D. in Physics from the University of Minnesota in 1965. He was a scientist at the University of 

Wyoming for 25 years and then in NOAA for 17 years, directing the Global Monitoring Division 

of NOAA’s Earth System Research Laboratory for a decade. Over a period of 30 years, he traveled to Antarctica 19 times for research and as director of NOAA’s South Pole Station. He was a 

reviewer for four Ozone Assessments and Lead Author of Chapter 12 (“Predicting Future Ozone 

Changes and Detection of Recovery”) of the 1998 Assessment.

Julius London (1917–2009). Julius London was born on March 26, 1917, in Newark, New 

Jersey. He received his Ph.D. in Meteorology and Oceanography from New York University in 

1951. After working for several years at NYU, he moved to the University of Colorado in 1961 

and remained there for his entire career, chairing the Department of Astro-Geophysics from 1966 

to 1969. He was an author in NASA and WMO Assessments that predated the Montreal Protocol, 

including leading the chapter on “Long Period Changes in Stratospheric Parameters” in the 1979 

Assessment, The Stratosphere: Present and Future, and chairing the Trends working group of the 

chapter on “Model Predictions and Trend Analysis” in the 1981 Assessment, The Stratosphere 

1981: Theory and Measurements. 

ix

Contents

Scientific Assessment of Ozone Depletion: 2010

Preface ................................................................................................................................................................................ xi

Prologue ........................................................................................................................................................................... xv

EXCUTIVE SUMMARY CONTENTS ............................................................................................................................. xxv

executive summary.................................................................................................................................................ES.1

CHAPTER 1: Ozone-Depleting Substances (ODSs) and Related Chemicals

 Coordinating Lead Authors: Stephen A. Montzka and Stefan Reimann

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

1.1 Summary of the Previous Ozone Assessment.................................................................................................. 1.7

1.2 Longer-Lived Halogenated Source Gases.........................................................................................................1.7

1.3 Very Short-Lived Halogenated Substances (VSLS).......................................................................................1.37

1.4 Changes in Atmospheric Halogen...................................................................................................................1.63

1.5 Changes in Other Trace Gases that Influence Ozone and Climate .................................................................1.75

References..................................................................................................................................................................1.86

CHAPTER 2: STRATOSPHERIC OZONE AND SURFACE ULTRAVIOLET RADIATION

 Coordinating Lead Authors: Anne Douglass and Vitali Fioletov

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

Introduction..................................................................................................................................................................2.5

2.1 Ozone Observations..........................................................................................................................................2.5

2.2 Polar Ozone.....................................................................................................................................................2.17

2.3 Surface Ultraviolet Radiation..........................................................................................................................2.31

2.4 Interpretation of Observed Ozone Changes....................................................................................................2.41

References..................................................................................................................................................................2.59

CHAPTER 3: Future Ozone and its impact on surface UV

 Coordinating Lead Authors: Slimane Bekki and Gregory E. Bodeker

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

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

3.2 Factors Affecting Future Ozone and Surface UV.............................................................................................3.7

3.3 Projections of Ozone through the 21st Century...............................................................................................3.18

3.4 Projections of UV Changes Related to Ozone Changes through the 21st Century .........................................3.42

3.5 Conclusions ...................................................................................................................................................3.46

References..................................................................................................................................................................3.49

Appendix 3A: Constructing Correlative Time Series Plots......................................................................................3.59
Contents

x

CHAPTER 4: Stratospheric Changes and Climate

 Coordinating Lead Authors: Piers M. Forster and David W.J. Thompson

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

4.0 Introduction and Scope .....................................................................................................................................4.3

4.1 Observed Variations in Stratospheric Constituents that Relate to Climate.......................................................4.4

4.2 Observed Variations in Stratospheric Climate................................................................................................4.10

4.3 Simulations of Stratospheric Climate Change ................................................................................................4.16

4.4 Effects of Variations in Stratospheric Climate on the Troposphere and Surface ...........................................4.25

4.5 What to Expect in the Future ..........................................................................................................................4.41

References..................................................................................................................................................................4.46

CHAPTER 5: A Focus on Information and Options for Policymakers

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

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

5.1 Summary of Previous Assessment and Key Issues to be Addressed in the Current Assessment.....................5.5

5.2 Metrics Used to Quantify Ozone and Climate Impacts ....................................................................................5.5

5.3 Future Baseline Scenarios...............................................................................................................................5.14

5.4 Impacts of Human Activities Relevant to Ozone Policy ................................................................................5.19

5.5 The World Avoided by Ozone Policy.............................................................................................................5.33

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

Appendix 5A

Table 5A-1: Direct Global Warming Potentials for selected gases..............................................................5.47

Table 5A-2: Assumptions made in obtaining production and emission estimates for

the baseline (A1) scenario........................................................................................................5.50

Table 5A-3: Mixing ratios (ppt) of the ODSs considered in the baseline (A1) scenario.............................5.54

Table 5A-4: Halocarbon indirect GWPs from ozone depletion using the EESC-based 

method described in Daniel et al. (1995).................................................................................5.56

TWENTY QUESTIONS AND ANSWERS ABOUT THE OZONE LAYER: 2010 UPDATE

Coordinating Lead Authors: David W. Fahey and Michaela I. Hegglin

I. Ozone in Our Atmosphere ............................................................................................................................... Q.4

II. The Ozone Depletion Process........................................................................................................................ Q.14

III. Stratospheric Ozone Depletion ...................................................................................................................... Q.31

IV. Controlling Ozone-Depleting Substances...................................................................................................... Q.45

V. Implications of Ozone Depletion and the Montreal Protocol ........................................................................ Q.52

VI. Stratospheric Ozone in the Future.................................................................................................................. Q.64

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

Preface

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

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

The Charge to the Assessment Panels

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

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

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

To meet this request, the Scientific Assessment Panel, the Environmental Effects Assessment Panel, and the Technology and Economic Assessment Panel have each prepared, about every 3-4 years, major assessment reports that updated 

the state of understanding in their purviews. These reports have been scheduled 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 2010 report is the latest in a series of eleven scientific assessments prepared by the world’s leading experts in the atmospheric sciences and under the international auspices of the World Meteorological Organization (WMO) 

and/or the United Nations Environment Programme (UNEP). This report is the seventh 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.

1995 Vienna Adjustment

1997 Montreal Adjustment

 and Amendment

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

xii

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 Montreal Adjustment

2010 Scientific Assessment of Ozone Depletion: 2010. WMO No. 52.

2011 23rd Meeting of the Parties

The Current Information Needs of the Parties

The genesis of Scientific Assessment of Ozone Depletion: 2010 occurred at the 19th Meeting of the Parties to the 

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

Decision XIX/20 (4), which stated that “...for the 2010 report, the Scientific Assessment Panel should consider issues 

including:

(a) Assessment of the state of the ozone layer and its future evolution;

(b) Evaluation of the Antarctic ozone hole and Arctic ozone depletion and the predicted changes in these phenomena;

(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 and the likely implications for 

the state of the ozone layer;

(d) Assessment of the interaction between climate change and changes on the ozone-layer;

(e) Assessment of the interaction between tropospheric and stratospheric ozone;

(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 into account among other things the 

expected impacts of climate change;

(g) Assessment of consistent approaches to evaluating the impact of very short-lived substances, including potential 

replacements, on the ozone layer;

(h) Identification and reporting, as appropriate, on any other threats to the ozone layer...”

The 2010 assessment has addressed all the issues that were feasible to address to the best possible extent.

The Assessment Process

The formal planning of the current assessment was started early in 2009. The Cochairs considered suggestions 

from the Parties regarding experts from their countries who could participate in the process. Furthermore, an ad hoc international scientific advisory group also suggested participants from the world scientific community. In addition, this 

advisory group contributed to crafting the outline of the assessment report. As in previous assessments, the participants 

represented experts from the developed and developing world. The developing 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 2010 assessment is contained in five chapters associated with ozone-layer topics, which are 

preceded by a Prologue: 

Prologue. State of the Science through the 2006 WMO/UNEP Assessment

Chapter 1. Ozone-Depleting Substances (ODSs) and Related Chemicals

Chapter 2. Stratospheric Ozone and Surface Ultraviolet Radiation

Chapter 3. Future Ozone and Its Impact on Surface UV

Chapter 4. Stratospheric Changes and Climate

Chapter 5. A Focus on Information and Options for Policymakers
Preface

xiii

The initial plans for the chapters of the 2010 Scientific Assessment Panel’s report were examined at a meeting that 

occurred on 24–25 June 2009 in London, England. The Coordinating Lead Authors 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 17–19 November 2009 in Fairfax, Virginia, United States, at which the Coordinating Lead Authors, Cochairs, and a small group of international experts focused 

on the scientific content of the draft chapters.

The second drafts of the chapters were reviewed by 122 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 28 June–2 July 

2010, the responses to these mail review comments were proposed by the authors and discussed by the 74 participants. 

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

on the UNEP web site on 16 September 2010) was prepared and completed by the attendees of the Les Diablerets meeting. 

A small science advisory group assisted the Cochairs during those Les Diablerets discussions of the Executive Summary, 

and also helped with advance preparations during a meeting in Toronto on 17–18 May 2010.

The 2010 State-of-Understanding Report

In addition to the scientific chapters and the Executive Summary, the assessment also updates the 2006 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 

and published separately in a companion booklet to 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: 2010 is the product of 312 scientists from 39 countries of the developed and 

developing world who contributed to its preparation and review1 (191 scientists prepared the report and 196 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.

1 Participating were Argentina, Australia, Belgium, Benin, Canada, Chile, Comores, Costa Rica, Czech Republic, Denmark, Egypt, Estonia, Finland, 

France, Germany, Greece, Hungary, India, Indonesia, Iran, Italy, Japan, Kenya, Malaysia, Mexico, New Zealand, Norway, Poland, Russia, Saudi Arabia, 

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

xv

Prologue

Prologue: State Of The Science Through The 2006 Wmo/Unep Assessment

A.R. Ravishankara, Paul A. Newman, John A. Pyle, and Ayité-Lô Ajavon

Scientists have known for many decades that the stratospheric ozone layer screens harmful ultraviolet radiation 

(UV) from the Earth’s surface. Therefore, it has also been known that the ozone layer protects against adverse effects on 

humans (e.g., skin cancer and cataracts), the biosphere (e.g., inhibiting plant growth and damaging ecosystems), and physical infrastructure of the modern era (e.g., degradation of materials). In the early 1970s, scientists recognized that human 

actions could deplete this protective layer in connection with nitrogen oxide emissions from a proposed fleet of supersonic 

aircraft flying in the stratosphere. Around that time, it was shown that human-produced chlorofluorocarbons (CFCs) that 

had been manufactured (and emitted to the atmosphere) had remained in the atmosphere because of their stability. Soon 

afterward, scientists warned that these CFCs that are stable in the lower atmosphere would get to the stratosphere, where 

they could deplete the ozone layer. They also warned that the depletion would be large if CFC emissions continued unabated. Various national and international assessments that estimated the impact of CFCs on the ozone layer were carried 

out. For example, using the then-state-of-the-art models of the atmosphere, a 1981 Assessment sponsored by the World 

Meteorological Organization (WMO) and agencies of the United States of America estimated that up to ~15% of the column ozone would be depleted by the middle of the 21st century if the CFC emissions went unabated at 1974 emission levels 

under certain assumptions about other emissions and changes (WMO, 1982). Studies also predicted a decrease in ozone of 

5–10% if a fleet of 500 supersonic aircraft emitting nitrogen oxides were to fly routinely in the stratosphere.

In 1985, massive ozone losses in measured column abundances during the Antarctic spring (the ozone hole) were 

reported and CFCs were implicated for the loss. Extensive research efforts showed that CFCs and other ozone-depleting 

substances (ODSs) containing chlorine and bromine were the cause. Further, measured global ozone abundances showed 

a decrease between 0.5% and 1.5% by 1980. Thus, ozone depletion was not just a phenomenon expected by the middle of 

the 21st century, but was already occurring. As a result of these findings on ozone depletion, stratospheric science rapidly 

evolved during the latter part of the 20th century, allowing understanding, diagnosis, and prediction of the evolution of 

the ozone layer; these rapid scientific developments provided a sound basis for the critical policy decisions that followed.

Faced with the potential impact of human-produced long-lived halogenated chemicals on stratospheric ozone, the 

Vienna Convention for the Protection of the Ozone Layer was enacted in 1985 to protect human health and the environment against adverse effects resulting from modification of the ozone layer. The recognition that CFC use was increasing, 

and scientific evidence that this increase would cause large ozone depletions, led in 1987 to the Montreal Protocol on Substances that Deplete the Ozone Layer, a protocol that regulated and slowed the production of designated ODSs. As new 

scientific knowledge became available over the next two decades, the Protocol has been amended and adjusted to provide 

additional protection for the ozone layer. The Montreal Protocol is now more than 20 years old and has been ratified by 

all of the world’s nations.

The Montreal Protocol, at its inception, established three expert panels—the Scientific Assessment Panel (SAP), 

the Environmental Effects Assessment Panel (EEAP), and the Technology and Economic Assessment Panel (TEAP). 

These panels provide the basis for science-based decision making via periodic assessment reports. The SAP’s primary 

focus is to provide an assessment of ozone layer science, including information about the abundances and emissions of 

ozone-depleting substances, ultraviolet radiation changes, along with additional information concerning policy options 

for consideration by the Parties to the Protocol. In addition, the SAP reports also aid other customers: various nations, by 

providing information needed for their decision making; industry, by providing a basis for technology choices; the broad 

science community, the EEAP, and the TEAP, with the latest information about the ozone layer science; the ozone research 

community, with information on the current science and gaps in knowledge; and the general public, including students 

and educators, with key information about this complex issue. The “Twenty Questions and Answers About the Ozone 

Layer” and its predecessors, which are companions to the SAP assessment reports, also help by providing clear, easy-tounderstand communication of the ozone layer issues to the Parties and the general public. Further, every four years, the 
Prologue

xvi

Cochairs of the three Assessment Panels compile a Synthesis Report based on the findings of their individual Assessment 

reports. These Assessments—individual Assessments and the Synthesis Report—together provide the latest information 

to the Parties to the Protocol.

Over the past two decades, the ozone depletion assessments have provided information updates roughly every four 

years and have been interspersed with a few brief reports on special topics that addressed urgent needs of the Parties to 

the Protocol. As knowledge of ozone layer science has increased, the assessments have built a vast amount of knowledge. 

Now, the SAP is addressing some key remaining issues regarding the ozone layer and its future development. They 

include the following:

• First, the Protocol has regulated human-produced ozone-depleting substances, resulting in the reduction of their abundances in the atmosphere. This effort has brought ozone depletion science into a period of accountability. The crucial 

questions now have become: Does the Montreal Protocol continue to work as envisioned? Were the specific actions 

effective in meeting the Protocol’s goals? Are the goals of the Vienna Convention also being met? How important 

are additional actions in returning ozone to its natural level? When will ozone levels return to preindustrial values? 

When will ozone levels return to the levels seen in 1980, a level that has become a benchmark for policymakers and 

the public? When will the ozone hole disappear? As ozone levels increase, will we observe decreases in surface 

ultraviolet radiation? What is our level of understanding of the workings of the stratosphere and how confident are 

we in our predictions for the future?

• Second, since the ozone layer is an integral part of the Earth system, other important questions have emerged: What is 

the influence of climate change on the stratospheric ozone layer and its future development? Specifically, how will the 

cooling of the stratosphere due to anthropogenic carbon dioxide (CO2) increases and the warming of the troposphere 

due to the increasing abundance of greenhouse gases influence the stratospheric ozone layer? How can we disentangle 

the influences of climate change on stratospheric ozone levels from to the influences of ozone-depleting substances?

• Third, the changes in stratospheric ozone are but one component of stratospheric climate change, and this poses questions such as: What are the effects of changes in stratospheric climate on the global-climate system? In addition, how 

will decreasing concentrations of ODSs impact climate?

• Fourth, ODSs and many substitutes for the ODSs are also potent greenhouse gases. Therefore, as ODSs are phased 

out and new chemicals take their place, questions emerge on the suitability of the replacements. They include: How 

will they impact the ozone layer? Do they have appreciable effects on climate? Do they have any other unwanted 

effects on the environment?

The SAP’s goal is to provide clear scientific answers to these questions. These questions provide the major thrust of the 

research in this area and are at the center of the current Assessment.

This current document provides the latest assessment of the science of the ozone layer. Below, we very briefly 

summarize our understanding of the science going into this Assessment. We summarize the findings of the most recent 

previous report of 2006 and note the key issues for the present Assessment.

Ozone-Depleting Substances (ODSs)

Emissions of ODSs were increasing at a substantial rate before the Montreal Protocol was enacted in 1987. As a 

result of the Protocol, emissions of most of the major ODSs—the chlorofluorocarbons (CFCs) and methyl chloroform 

(CH3CCl3)—began decreasing soon thereafter. Because of the long lifetimes of CFCs, their atmospheric abundances continued to increase in the early 1990s even as their emissions were decreasing. However the abundance of the short-lived 

methyl chloroform responded quickly, as expected, and started to decrease in the atmosphere. Originally, some of the 

CFC replacements were the so-called transition substitutes (hydrochlorofluorocarbons, HCFCs); they contained chlorine 

but were shorter lived than the CFCs they replaced. This substitution led to a lower accumulation of the HCFCs and a 

smaller fraction of their emissions being transported to the stratosphere. Subsequently, the HCFCs were also selected for 

phase-out, and non-chlorine containing substitutes are now being phased in. Because of these changes, the sum of the 

abundances of chlorine and bromine ODS species in the troposphere, as measured by equivalent chlorine (ECl), reached 

a peak in the 1994–1995 time period and has continued to decrease thereafter. The majority of the decrease in the ECl is 

attributed to the rapid decline of emissions of the short-lived methyl chloroform and, to a lesser extent, methyl bromide. 
Prologue

xvii

Prologue Box 1. A Clarification of the Lexicon: Ozone Destruction, Ozone Depletion, 

Ozone-Depleting Substances, and Montreal Gases

Ozone Destruction and Ozone Depletion

The abundance of ozone at a particular point in the stratosphere, the column abundance of ozone above a given 

geographical location, and the total amount of ozone in the stratosphere are controlled by a combination of production, 

destruction, and transport (of ozone and other chemicals into and out of the region of interest). The major mechanism for 

the production of ozone in the stratosphere is the breaking up of molecular oxygen (O2) by solar UV of wavelengths less 

than 242 nanometers (photolysis) to make oxygen atoms (O), followed by the reaction of oxygen atoms with molecular 

oxygen to make ozone. The destruction of ozone occurs via the reactions of oxygen atoms (O) with ozone (O3) (the 

Chapman Mechanism), as well as through cyclic chemical reactions involving naturally occurring species such the oddhydrogen radicals (HOx: OH and HO2), nitrogen oxide radicals (NOx: mostly NO and NO2), and/or halogen radicals. The 

radicals are produced in the stratosphere by photolysis and oxidation of source gases (N2O, H2O, CH4, and a variety of 

chlorine- and bromine-containing compounds). In the absence of interference from the human emissions influencing 

the abundance of catalysts, there is a natural balance and this balance determines the ozone abundance in a location, the 

column amount over a region, and the total amount of ozone in the stratosphere. The natural amounts vary on a variety of 

timescales: daily variations in the ozone column are driven by meteorological variability (“weather”); seasonal variations 

are driven by changes in stratospheric temperature and winds; multiannual variations are driven by changes in solar input, 

by natural variations in the emissions of the source gases, and by interannual variability in stratospheric winds.

The natural abundance of stratospheric ozone can be changed by human influence. This change can be brought 

about by changes in production, destruction, and transport. The ozone abundance arises from a balance between these 

terms. Human emissions, for the most part, have led to an enhancement in the destruction term, shifting the balance to 

lower ozone abundance. Thus, any human emission of chemicals (gases or particles) that contributes to the enhancement 

of the ozone destruction term in the balance leads to a lowering of ozone, i.e., ozone layer depletion, and is evidenced 

by changes in the amount at a location, in the column amount above a location, or the total amount in the stratosphere. 

Because the destruction occurs through catalytic cycles that regenerate the ozone-destroying radicals multiple times, 

small changes in the source gases (and hence in radical concentrations) can have a large impact on ozone.

Ozone-Depleting Substances and Montreal Gases

If there is an increase in concentrations of any of these source gases that contribute nitrogen, hydrogen, or halogen radicals to the stratosphere, there will be an increase in ozone-destroying radicals and hence in stratospheric ozone 

destruction. Changes in the sources gases could occur either naturally (e.g., by biogenic processes at the surface) or 

anthropogenically (by increased industrial emissions); some source gases are emitted both naturally and anthropogenically. The response of the stratosphere does not depend on whether the changes are natural or anthropogenic; the stratosphere does not “care.” However, scientists and policymakers do care and in some circumstances it is useful to have a 

terminology that distinguishes the different origins of the source compounds. Therefore, ozone-depleting substances 

(ODSs) are those whose emissions come from human activities.

It will be important in the Assessment also to consider specifically gases that have been regulated (and which traditionally we have called ODSs). Thus, the Montreal Protocol has controlled the production (and hence their emissions 

into the atmosphere) of certain chemicals that are listed as controlled substances in Annexes A, B, C, and E of the Protocol. We will continue to call the controlled substances of the Montreal Protocol as ozone-depleting substances, or 

ODSs for short. This definition keeps the continuity in usage and will be clear to the Parties to the Montreal Protocol.

The above description yields a few key points. First, the ozone abundance can be changed not only by destruction 

but also via influence on production, transport, and stratospheric climate. There are long-standing examples of such 

production enhancements by hydrocarbons, in particular methane, via what is usually called smog chemistry (i.e., the 

chemistry that leads to the tropospheric pollutant ozone production). Second, ozone abundances can be changed by both 

changes in the concentrations of active agents, as well as by changes in the rates at which these chemical reactions occur. 

The most noteworthy way is by changes in the stratospheric climate (i.e., temperature), such as that caused by the enhancements in carbon dioxide in the atmosphere. Third, the ozone abundance can be influenced by changes in transport, 

such as that arising from a changing climate. And fourth, while the Montreal Protocol controls many substances that 

deplete ozone, not all such substances are currently controlled and, for clarity, they are not called ODSs here. Reference 

to such substances are clearly noted in this Assessment. 
Prologue

xviii

The tropospheric abundance of ECl by the end of 2005 was shown in the previous Assessment to have decreased to roughly 

92% of its maximum value seen during the period between 1992 and 1994 (i.e., about a 8% decline in roughly 14 years); 

these values will be updated in this report.

Balloon, aircraft, and satellite observations, and the interpretation of those observational data, show clearly that 

stratospheric abundances of chlorine and bromine are also decreasing. The vertical and temporal variations of the ODS 

species are generally consistent with our understanding of atmospheric dynamics and stratospheric chemical processes, 

though there are some quantitative differences between observations and calculations. Improvements in quantification of 

these variations are expected. These improvements will enable an even better definition of the stratospheric distribution 

and trends of the ODSs as well as their degradation products, which will enable a better quantification of their individual 

role in ozone layer depletion.

The CFCs, as well as some halons (which are sources of bromine to the stratosphere), have lifetimes ranging from 

several decades to a few centuries. Hence, the decline of stratospheric chlorine and bromine levels to values observed 

before 1980 will take decades.

As noted above, CFCs have been replaced by non-ozone depleting technologies, by substitutes that deplete less 

ozone (e.g., hydrochlorofluorocarbons or HCFCs), and by non-ozone depleting substances (e.g., hydrofluorocarbons or 

HFCs). The atmospheric levels of these less-depleting and non-depleting substitutes have grown rapidly over the last 

decade. HCFCs typically have shorter atmospheric lifetimes and lower Global Warming Potentials (GWPs) than CFCs, 

but HFC substitutes for HCFCs typically have comparable, and in a few cases even longer, atmospheric lifetimes and 

comparable or larger GWPs; but they have Ozone Depletion Potential (ODP) values of essentially zero. The increases 

observed for HCFCs and HFCs reflect their widespread use as ODS replacements and our understanding of their 

atmospheric lifetimes.

Global Stratospheric Ozone and Its Temporal and Spatial Trends

Global atmospheric column ozone amounts decreased over the decades from the 1970s to the 1990s, with a decrease 

amounting to 3.5% between average 1964–1980 and 2002–2005 values. Springtime Antarctic ozone levels slowly decreased in the 1970s and exhibited rapid decreases in the 1980s and early 1990s. In the 14–20 km layer of the Antarctic 

stratosphere, where most of the ozone resides, virtually all of the ozone is now destroyed every year in the late August 

to early October period. Large Arctic ozone depletions have also been observed in the spring in some years during the 

last two decades, but Arctic ozone depletion is modulated strongly by variability in atmospheric dynamics, transport, and 

temperature. The very high levels of chlorine and bromine from ODSs directly cause the observed large polar ozone depletions (both over the Antarctic and the Arctic).

Atmospheric ozone levels (often measured as a column amount) exhibit well-known and understood variations in 

space and time. Ozone amounts are influenced not only by the concentrations of ODSs but also by atmospheric transport 

(winds), incoming solar radiation, aerosols (fine particles suspended in the air), and other natural compounds. Given 

natural variability, methods used to measure stratospheric ozone must be consistent and very stable over decades if they 

are to be used to detect the changes expected over these long periods due to the changes in ODS abundances. Based on 

observations from ground-based instruments and satellites, it is clear that global ozone levels reached a minimum in the 

mid-1990s. Since then the levels have not decreased further nor have they increased substantially. Similarly, the Antarctic 

ozone hole continues to be no worse than in the mid-1990s but there also has been no discernible improvement, consistent 

with predictions from previous assessments. Both annual global ozone and the springtime Antarctic ozone levels continue 

to vary from year to year because of meteorological variability. There is no discernible ozone depletion over the tropics 

outside of the natural background variations. Vertically, ozone depletion is most evident in the lower and upper stratosphere, with minimal changes in the mid-stratosphere.

In the last few decades, ozone levels in the stratosphere have responded to volcanic eruptions that have injected 

large amounts of sulfur dioxide into the stratosphere, which then forms sulfate aerosols in this region. These sulfate aerosols enhance the ozone depletion by chlorine from ODSs. The very large ozone depletions induced by the presence of 

aerosols following the eruptions of Mt. Pinatubo (1991) and El Chichón (1982) are very clearly seen in the records in the 

Northern Hemisphere. The influence of these eruptions persisted for several years. As the stratosphere recovered from the 

volcanic emissions, there were corresponding changes in ozone. The ozone response depends on the effective abundances 

of chlorine and bromine in the stratosphere. Thus, response to future volcanic eruptions will likely be smaller because 
Prologue

xix

chlorine/bromine concentrations will be smaller (see Figure P-1). The mechanisms for these changes are qualitatively 

understood, but some uncertainties remain in their quantification.

The observed levels of ozone described above and the vertical, latitudinal, and seasonal structure of their temporal 

trends, as well as the spatial and temporal variability, are consistent with our combined understanding of the atmospheric 

motions (transport), the chemistry, and the level of ODSs in the atmosphere. Even though some details of chemical and 

dynamical processes are uncertain, atmospheric models have been largely successful in reproducing observed ozone levels 

and their temporal and spatial variations. The link between ODSs and ozone depletion was clearly established in the 1989 

Ozone Assessment (WMO, 1991) and that conclusion has only been strengthened since then.

Surface UV Changes

Ultraviolet radiation (UV) from the Sun is divided into wavelength bands. UV-B is the band that leads to serious 

medical problems. Fortunately, the majority of the UV-B is absorbed by ozone. The surface UV-B and UV-A levels 

(expressed as the UV Index) are directly related to the amount of overhead ozone. Other factors such as clouds, aerosols, 

ground reflectivity, and other tropospheric pollutants also influence surface UV-B. The data outside of the polar regions 

shows that, consistent with the observed small ozone depletion, there have not been large increases in surface UV-B over 

the last few decades. The relatively small increases of surface UV-B in the midlatitudes, which are expected based on the 

observed ozone decline, are responsible for small changes in the UV background level, which are superposed by other 

strong effects, such as changes in cloudiness. However, since medical impacts are UV-dose related, the UV changes due to 

ozone depletion are nonetheless important. In contrast, over Antarctica, and on occasion in other parts of the high latitudes 

in the Southern Hemisphere, large increases in UV-B have been seen; they are clearly associated with the ozone hole or 

the remnants of the ozone hole passing over the measurement sites.

The changes in UV-B levels are consistent with our understanding of UV transmission and the other factors that 

influence UV-B at the surface.

Factors that Influence Stratospheric Ozone and Its Future

The change in the atmospheric ODS concentrations is the most important factor in the ozone layer changes that have 

occurred over the past half a century and also in the predicted return of the ozone layer to levels that existed prior to 1980. 

However, many other aspects of the Earth system are also changing. These include changes in climate and tropospheric 

composition.

Climate change influences the stratosphere in many ways. The primary influence is a cooling of the mid- to 

upper stratosphere due to increases in carbon dioxide (CO2) via radiation to space, which is a well-understood process. 

This cooling has been clearly seen in measured temperatures. The cooling influences the ozone loss rates in the stratosphere—increasing it in the lower stratosphere and decreasing it in upper stratosphere. At the same time the warming in 

the troposphere accelerates processes of ozone formation. Further, climate change has an effect on transport between the 

stratosphere and the troposphere and within the stratosphere, and in turn, climate will influence the recovery of ozone layer 

from the effects of ODSs.

Tropospheric changes also influence stratospheric ozone levels. For example, an increased abundance of methane 

(CH4) in the troposphere will result in more methane being transported to the stratosphere, where methane interacts with 

chlorine compounds, converting active chlorine that destroys ozone to inactive hydrogen chloride (HCl) that does not 

destroy ozone. Changes in methane also lead to changes in water vapor in the stratosphere, with important consequences. 

Similarly, changes in nitrous oxide (N2O) also influence ozone destruction. Other tropospheric changes of interest include 

processes leading to increases in sulfur in the stratosphere. In some cases, changes of these tropospheric processes may 

be related to climate change. For instance, climate change may affect biogeochemical cycles and cause an increase in 

tropospheric concentrations of certain species as well as the transport rate between the troposphere and the stratosphere. 

The latter may be particularly important for the very short-lived species.

The timeline of the ozone evolution from the pre-ODS era to roughly 2100 was presented in the 2006 Assessment to 

facilitate discussion on recognition and attribution of the recovery of the ozone layer. This approach provided a pathway for 

interim conclusions on this issue, but many issues remained unresolved. They include: How should recovery be defined? 

What time period is appropriate as a baseline against which we can measure recovery? How do we separate ozone changes 
Prologue

xx

due to ODSs from those due to changes in climate and tropospheric composition? How do we describe and attribute future 

changes in levels of ozone? Given the natural variability, at which point will one be confident of the recovery from ODS 

effects? This Assessment addresses some of these issues and concepts (see Prologue Box 2 on Recovery Issues).

Influence of Stratospheric Ozone and ODS Changes on Climate

As noted above, increases in CO2 in the atmosphere have led to a clear decrease in upper stratospheric temperature. 

This temperature trend is a very clear signature of the radiative influence of increasing CO2 abundances. Changes in the 

stratosphere—be it the temperature decrease due to CO2 increases or ozone layer depletion due to ODSs—are an integral 

part of the changes to the Earth system. Further, these changes in the stratosphere influence what happens at the surface. 

Therefore, the influence of stratospheric changes on surface climate is an important issue.

Ozone is a greenhouse gas that greatly influences the Earth’s energy budget. Therefore, ozone changes—depletion 

in the stratosphere due to ODSs, recovery from the depleted state as ODSs decline, and tropospheric ozone changes—also 

influence climate. Further, many of these ODSs that deplete the ozone layer are also greenhouse gases. Consequently, 

they influenced Earth’s climate in the past as their abundances increased and will continue to do so, albeit to a lesser extent, 

as their abundances decrease in response to compliance with the Montreal Protocol. Furthermore, many of the substitutes 

for CFCs and HCFCs are also potent greenhouse gases and their contribution to climate change will depend on the their 

potency for warming and their emission rates.

These are some of the emerging issues that have been covered only briefly in the past due to a primary focus on 

ozone depletion issues. As research on the influence of stratospheric changes on the overall climate has emerged, the current Assessment is devoting more attention to this topic.

Major Findings of the Previous Assessment in 2006

The major findings of the 2010 Assessment are given in the Executive Summary that follows this Prologue. To 

place these findings in context and show the changes in our knowledge over the past four years, we provide below the 

summary of the 2006 Assessment (WMO, 2007). Further, for ease of comparison, the findings from the 2006 Assessment 

are grouped according to where they are covered in the 2010 Assessment; i.e., the 2006 Assessment is mapped on to the 

2010 Assessment’s structure. 

A major finding of the previous Assessment in 2006, the tenth in a series of Assessments dating back to 1981, was 

that the Montreal Protocol was working as intended. Some specific findings of the 2006 Assessment are summarized in 

the schematic shown as Figure P-1.

The high-level findings of the previous Assessment (WMO, 2007) include the following.

Findings of the 2006 Assessment that are related to “Ozone-Depleting Substances (ODSs) and Related Chemicals” covered in Chapter 1 of the 2010 Assessment:

1. 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.

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

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

the stratosphere.

3. Our quantitative understanding of how halogenated very short-lived substances contribute to halogen levels in 

the stratosphere has improved significantly since the 2002 Assessment (WMO, 2003), with brominated very 

short-lived substances believed to make a significant contribution to total stratospheric bromine and its effect 

on stratospheric ozone.

Findings of the 2006 Assessment that are related to “Stratospheric Ozone and Surface Ultraviolet Radiation” in the past 

and our understanding of its changes covered in Chapter 2 of the 2010 Assessment:

1. 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.
2. 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.

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

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

averaged over 60°S–60°N over roughly the 1995–2005 period are related to the near constancy of stratospheric 

ozone-depleting gases during this period.

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

been decreasing since the late 1990s. 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.

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

of low total ozone.

Findings of the 2006 Assessment that are related to “Future Ozone and Its Impact on Surface UV” covered in Chapter 3 

of the 2010 Assessment:

1. 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.

xxi

Global ozone

change

Ozone-depleting

chlorine and bromine

in the stratosphere

OD

S

production

Ultraviolet radiation

change

[Figure reproduced from the 2006 Ozone Assessment (WMO, 2007).]

Figure P-1. Ozone-Depleting Substances, 

the Ozone Layer, and UV Radiation: Past, 

Present, and Future.

(a) Production of ozone-depleting substances 

(ODSs) before and after the 1987 Montreal 

Protocol and its Amendments, from baseline 

scenario A1. Chlorofluorocarbons (CFCs) 

are shown in black; additional ODSs from 

hydrochlorofluorocarbons (HCFCs) are in gray. 

Note: HCFCs, which have been used as CFC 

replacements under the Protocol, lead to less 

ozone destruction than CFCs.

(b) Combined effective abundances of 

ozone-depleting chlorine and bromine in the 

stratosphere. The range reflects uncertainties 

due to the lag time between emission at the 

surface and the stratosphere, as well as different 

hypothetical ODS emission scenarios.

(c) Total global ozone change (outside of the 

polar regions; 60°S-60°N). Seasonal, quasibiennial oscillation (QBO), volcanic, and solar 

effects have been removed. The black line 

shows measurements. The gray region broadly 

represents the evolution of ozone predicted 

by models that encompass the range of future 

potential climate conditions. Pre-1980 values, 

to the left of the vertical dashed line, are 

often used as a benchmark for ozone and UV 

recovery.

(d) Estimated change in UV erythemal 

(“sunburning”) irradiance for high sun. The 

gray area shows the calculated response to the 

ozone changes shown in (c). The hatched area 

shows rough estimates of what might occur 

due to climate-related changes in clouds and 

atmospheric fine particles (aerosols). 
Prologue

xxii

2. 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.

3. 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.

4. Large ozone losses will likely continue to occur in cold Arctic winters during the next 15 years.

5. Chemical reaction rates in the atmosphere are dependent on temperature, and thus the concentration of ozone 

is sensitive to temperature changes caused by climate change.

Findings of the 2006 Assessment that are related to the influence of “Stratospheric Changes and Climate” covered in 

Chapter 4 of the 2010 Assessment:

1. The stratospheric cooling observed during the past two decades has slowed in the recent years up to 2005.

2. Changes to temperature and circulation of the stratosphere affect climate and weather in the troposphere.

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

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

5. 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.

Findings of the 2006 Assessment that are related to “A Focus on Options and Information for Policymakers” covered in 

Chapter 5 of the 2010 Assessment:

1. 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.

2. The dates for the return of the global ozone layer and the Antarctic ozone hole to 1980 levels were provided 

based on the best available information to be around, respectively, 2049 and 2065.

3. Many potential options for accelerating the recovery of the ozone layer were evaluated and presented. 

Organization of the Current Assessment

Much new information has been generated since the 2006 Assessment. Further, the information needs of the Parties 

to the Protocol have also changed. The specific requests of the Parties to the SAP are given in the Preface of this Assessment. Of particular note are the questions related to the influence of stratospheric changes on Earth’s climate. This is 

somewhat of a new issue to the SAP and thus demands a chapter of its own. 

This Assessment is an update to previous Assessments, and in particular the 2006 Assessment. However, as noted 

above, the changes in ozone and UV are not rapid and there are no new major findings in this area. To reflect this updating 

approach and consolidation of information, the structure of this Assessment differs from the most recent reports. In this 

Assessment, Chapter 1 deals with all issues related to ODSs; they include long-lived and very short-lived halocarbons as 

well as the replacements for the ODSs. In particular, it covers the trends and abundances of the replacements for ODSs that 

are greenhouse gases (but not ODSs), such as HFCs that are being discussed by the Parties to the Protocol for regulation. 

Chapter 2 deals with all observations of ozone and surface UV to date and our understanding of these observations, including a discussion of the current state of polar ozone. Chapter 3 focuses primarily on the future response of the ozone layer 

and UV-B radiation to reduced halocarbon emissions and other changes in an effort to focus on the question: What should 

one anticipate for ozone layer depletion and its consequences? It also picks up the issue of the definition and recognition 

of the recovery of the ozone layer first discussed in the 2006 Assessment. Of particular note are the issues related to the 

influence of stratospheric changes on climate. This issue was briefly described in the 2006 Assessment, which mostly 

focused on the influence of climate change on the recovery of the ozone layer. Because of the emergence of information 

on the influence of the stratospheric changes on Earth’s climate, we have added a new chapter—Chapter 4—to address this 

topic. Chapter 4 focuses on the two-way connection between stratospheric changes and climate changes. This places the 

effects of halocarbon-induced ozone depletion on climate in the broader context of other stratospheric changes. Chapter 5 

is expanded to include not only the policy options, often posed in hypothetical terms, available for further action but also 

other information relevant to the Parties to the Protocol.
Prologue

xxiii

Prologue Box 2. Recovery of the Ozone Layer: Concepts and Practical Issues

A conceptual diagram of the behavior of stratospheric ozone between 1960 and 2100 was presented in Chapter 6 

of the 2006 Assessment (Bodeker and Waugh et al., 2007: “The Ozone Layer in the 21st Century”). A slightly modified 

version of this diagram is shown below.

As noted in the 2006 Assessment, stratospheric ozone abundances should change in response to decreases of 

ODSs and in response to other factors that influence ozone levels in the stratosphere. The other major factors are changes 

in temperature of the stratosphere because of increases in CO2, changes in transport associated with climate change, and 

changes in tropospheric composition.

The ODS increases in the past few decades depleted the ozone layer. In the future, as ODSs decrease, the atmosphere in general—and the stratosphere in particular—should have decreasing amounts of ozone-destroying halogen 

catalysts. This decrease will follow the emissions of ODSs but will be shifted to later times because ODSs generally 

have long atmospheric lifetimes.

The past and future timeline of ozone behavior has been categorized as: stage I—slowing of ozone decline; stage 

II—onset of ozone increases; and stage III—full recovery of ozone from ODSs. In this idealization it is assumed that 

ozone production is not altered significantly, and that the climate and tropospheric changes are sufficiently small that the 

influence of ODSs is the predominant factor that controls the rate of depletion of the ozone layer. Of course, because of 

natural interannual variability, the ozone abundances do not show sufficiently clear changes to allow precise identification of these timeline stages. (Continued on following page.)

1960 End of2 1

st

ce ntury

Ti me

Range of projections 

increasing ozon e de cr ea sing ozone 

Ozone Change

1980

1980 Levels

1960 Levels

Montreal Protocol Benchmark

Stage iii: Full

recovery of ozone

from ODSs

Stage ii: Onset of

ozone increases

Stage i: Slowing

of ozone decline

Figure P-2. A conceptual diagram of the evolution of column ozone between 60°N and 60°S between 

1960 and 2100 (the x-axis is not to scale) adapted from Fig. 6-1 in the 2006 Assessment. The observations are discussed in Chapter 2. The thick red line is a representation of the ozone amounts observed to 

date and projected for the future. The red-shaded region represents the model results predicted for the 

future. The Montreal Protocol 1980 ozone level benchmark is shown as the horizontal line. The dashed 

thick gray line represents the somewhat uncertain 1960 levels. The three recovery stages are shown by 

green dashed ellipses.
Prologue

xxiv

References

Bodeker, G.E., and D.W. Waugh (Lead Authors), H. Akiyoshi, P. Braesicke, V. Eyring, D.W. Fahey, E. Manzini, M.J. 

Newchurch, R.W. Portmann, A. Robock, K.P. Shine, W. Steinbrecht, and E.C. Weatherhead, The ozone layer in the 

21st century, Chapter 6 in Scientific Assessment of Ozone Depletion: 2006, Global Ozone Research and Monitoring 

Project–Report No. 50, 572 pp., World Meteorological Organization, Geneva, Switzerland, 2007.

WMO (World Meteorological Organization), The Stratosphere 1981: Theory and Measurements, Global Ozone Research 

and Monitoring Project–Report No. 11, 516 pp., Geneva, Switzerland, 1982.

WMO (World Meteorological Organization), Scientific Assessment of Stratospheric Ozone: 1989, Global Ozone Research 

and Monitoring Project–Report No. 20, Geneva, Switzerland, 1991. [Referred to as the 1989 Assessment.]

WMO (World Meteorological Organization), Scientific Assessment of Ozone Depletion: 2002, Global Ozone Research and 

Monitoring Project–Report No. 47, Geneva, Switzerland, 2003. [Referred to as the 2002 Assessment.]

WMO (World Meteorological Organization), Scientific Assessment of Ozone Depletion: 2006, Global Ozone Research 

and Monitoring Project–Report No. 50, 572 pp., Geneva, Switzerland, 2007. [Referred to as the 2006 Assessment.]

Prologue Box 2, continued.

This three-stage timeline is a very useful conceptual picture for understanding ozone changes, diagnosing the 

current and future trends, and attempting to predict future ozone levels. However, as noted above, the ozone timeline is 

also influenced by other changes—climate change, volcanic eruptions that introduce sulfate aerosols in the stratosphere, 

and tropospheric composition changes. Further, the natural (and forced) variability in the Earth system will lead to difficulties in identifying as well as attributing these changes. These variabilities occur not only in the ozone abundances 

but also in the ODS levels, as climate change and other changes will alter when the ODS levels will reach values seen 

prior to 1980.

For all practical purposes, the Montreal Protocol has used 1980 levels as the time when there was little perturbation of the ozone layer by ODSs. This does not mean that there was no ozone depletion in 1980. Indeed, retroactive 

analyses of observations show that the ozone hole was growing prior to 1980. Yet we use 1980 levels of ODSs as the 

level when the ozone layer was not significantly influenced by ODSs and we will continue to use this date as a benchmark in this Assessment.

Because of factors other than ODSs, the ozone levels in the future could easily go above the values that were 

present either in the 1980s or even the 1960s. This situation was described in the previous Assessment as a “superrecovery.” Of course, this is not recovery from the influence of ODSs but due to other factors, primarily CO2. Therefore, the use of the term “super-recovery” differs from references to recovery from ODS-forced ozone depletion.
Executive Summary

Contents

OVERVIEW .............................................................................................................................................................ES.1

CHANGES IN GASES THAT AFFECT STRATOSPHERIC OZONE AND CLIMATE .....................................ES.1

Ozone-Depleting Substances and Substitutes: Tropospheric Abundances and Emissions.............................ES.1

CFCs, HCFCs, HFCs, and Climate Change....................................................................................................ES.2

Total Chlorine and Bromine and Implications for Ozone Depletion ..............................................................ES.3

Figure ES-1: Emissions of ODSs and Their Substitutes...............................................................................ES.3

OZONE AND CLIMATE: ANTARCTIC................................................................................................................ES.4

OZONE AND CLIMATE: GLOBAL AND ARCTIC.............................................................................................ES.4

Figure ES-2: Schematic of the Influence of Ozone-Depleting Substances and Climate Change

 on the Stratospheric Ozone Layer, and the Influence of Ozone Changes on

 Surface Ultraviolet Radiation ..................................................................................................ES.5

INFORMATION FOR POLICYMAKERS AND OPTIONS FOR POLICY FORMULATION............................ES.7

Information for Policymakers .........................................................................................................................ES.7

Options for Policy Formulation.......................................................................................................................ES.8

Table ES-1: Hypothetical Cases ..................................................................................................................ES.9

APPENDIX: SCIENTIFIC SUMMARIES OF THE CHAPTERS (included in the individual chapters)

Chapter 1: Ozone-Depleting Substances (ODSs) and Related Chemicals

Figure S1-1: Stratospheric EESC Relative to Peak Abundances Versus Time

Table S1-1: Radiative Forcings of ODSs and Other Gases and Their Recent Changes

Chapter 2: Stratospheric Ozone and Surface Ultraviolet Radiation

Table S2-1: Summary of Ozone Changes Estimated from Observations

Chapter 3: Future Ozone and Its Impact on Surface UV

Chapter 4: Stratospheric Changes and Climate

Chapter 5: A Focus on Information and Options for Policymakers

Table S5-1: Summary of Hypothetical Cases for Accelerating the Recovery

 of the Ozone Layer and Reducing Carbon-Equivalent Emissions

xxv

ES.1

Executive Summary

OVERVIEW

It has been recognized since the 1970s that a number of compounds emitted by human activities deplete stratospheric ozone. The Montreal Protocol on Substances that Deplete the Ozone Layer was adopted in 1987 to protect global 

ozone and, consequently, protect life from increased ultraviolet (UV) radiation at Earth’s surface. Chlorine- and brominecontaining substances that are controlled by the Montreal Protocol are known as ozone-depleting substances (ODSs). 

ODSs are responsible for the depletion of stratospheric ozone observed in polar regions (for example, the “ozone hole” 

above Antarctica) and in middle latitudes. The severe depletion of stratospheric ozone observed in the Antarctic has 

increased UV at the surface and affected climate at southern high latitudes.

The Montreal Protocol and its Amendments and Adjustments have successfully controlled the global production 

and consumption of ODSs over the last two decades, and the atmospheric abundances of nearly all major ODSs that were 

initially controlled are declining. Nevertheless, ozone depletion will continue for many more decades because several key 

ODSs last a long time in the atmosphere after emissions end.

In contrast to the diminishing role of ODSs, changes in climate are expected to have an increasing influence on stratospheric ozone abundances in the coming decades. These changes derive principally from the emissions of long-lived greenhouse gases, mainly carbon dioxide (CO2), associated with human activities. An important remaining scientific challenge is 

to project future ozone abundances based on an understanding of the complex linkages between ozone and climate change.

Most ODSs are potent greenhouse gases. The buildup of ODS abundances over the last decades contributes to 

global warming. The actions taken under the Montreal Protocol have reduced the substantial contributions these gases 

would have made to global warming.

There is now new and stronger evidence of the effect of stratospheric ozone changes on Earth’s surface climate, 

and of the effects of climate change on stratospheric ozone. These results are an important part of the new assessment of 

the depletion of the ozone layer presented here.

CHANGES IN GASES THAT AFFECT STRATOSPHERIC OZONE AND CLIMATE

Changes in the global atmospheric abundance of a substance are determined by the balance between its emissions 

and removals from the atmosphere. Declines observed for ozone-depleting substances controlled under the Montreal 

Protocol are due to global emission reductions that have made emissions smaller than removals. Most ODSs are potent 

greenhouse gases. As the majority of ODSs have been phased out, demand for hydrochlorofluorocarbon (HCFC) and 

hydrofluorocarbon (HFC) substitutes for the substances controlled under the Montreal Protocol has increased; these are 

also greenhouse gases. HCFCs deplete much less ozone per kilogram emitted than chlorofluorocarbons (CFCs), while 

HFCs are essentially non-ozone depleting gases.

Ozone-Depleting Substances and Substitutes: Tropospheric Abundances and Emissions

•	 The amended and adjusted Montreal Protocol continues to be successful at reducing emissions (Figure ES-1)

and thereby abundances of most controlled ozone-depleting substances in the lower atmosphere (troposphere),

as well as abundances of total chlorine and total bromine from these ozone-depleting substances. By 2008, the 

total tropospheric abundance of chlorine from ODSs and methyl chloride had declined to 3.4 parts per billion (ppb) 

from its peak of 3.7 ppb. However, the rate of decline in total tropospheric chlorine by 2008 was only two-thirds as 

fast as was expected. This is because HCFC abundances increased more rapidly than expected, while CFCs decreased 

more slowly than expected. The discrepancy in CFC decreases is most likely because of emissions from “banks” in 

existing applications such as refrigerators, air conditioners, and foams. The rapid HCFC increases are coincident 
ES.2

Executive Summary

with increased production in developing countries, particularly in East Asia. The rate of decline of total tropospheric 

bromine from controlled ODSs was close to that expected and was driven by changes in methyl bromide.

•	 Declines in CFCs made the largest contribution to the observed decrease in total tropospheric chlorine during

the past few years and are expected to continue to do so through the rest of this century. Observations show 

that CFC-12 tropospheric abundances have decreased for the first time. The decline of methyl chloroform (CH3CCl3) 

abundances made a smaller contribution to the decrease in total chlorine than described in past Assessments, because 

this short-lived substance has already been largely removed from the atmosphere.

•	 Carbon tetrachloride (CCl4) tropospheric abundances have declined less rapidly than expected. Emissions derived from data reported to the United Nations Environment Programme (UNEP) are highly variable and on average 

appear smaller than those inferred from observed abundance trends. Although the size of this discrepancy is sensitive 

to uncertainties in our knowledge of how long CCl4 persists in the atmosphere (its “lifetime”), the variability cannot 

be explained by lifetime uncertainties. Errors in reporting, errors in the analysis of reported data, and/or unknown 

sources are likely responsible for the year-to-year discrepancies.

•	 Observations near the tropical tropopause suggest that several very short-lived industrial chlorinated chemicals, not presently controlled under the Montreal Protocol (e.g., methylene chloride, CH2Cl2; chloroform,

CHCl3; 1,2 dichloroethane, CH2ClCH2Cl; perchloroethylene, CCl2CCl2), reach the stratosphere. However, 

their contribution to stratospheric chlorine loading is not well quantified.

•	 Bromine from halons stopped increasing in the troposphere during 2005–2008. As expected, abundances of 

halon-1211 decreased for the first time during 2005–2008, while halon-1301 continued to increase but at a slower rate 

than in the previous Assessment.

•	 Tropospheric methyl bromide abundances continued to decline during 2005–2008, as expected due to reductions in industrial production, consumption, and emission. About half of the remaining methyl bromide consumption was for uses not controlled by the Montreal Protocol (quarantine and pre-shipment applications).

•	 Tropospheric abundances and emissions of some HCFCs are increasing faster now than four years ago. Abundances of HCFC-22, the most abundant HCFC, increased more than 50% faster in 2007–2008 than in 2003–2004, 

while HCFC-142b abundances increased about twice as fast as in 2003–2004. HCFC-141b abundances increased at 

a similar rate to that observed in 2003–2004. Total emissions of HCFCs are projected to begin to decline during the 

coming decade due to measures already agreed to under the Montreal Protocol (Figure ES-1).

•	 Tropospheric abundances and emissions of HFCs, used mainly as substitutes for CFCs and HCFCs, continue

to increase. For example, abundances of HFC-134a, the most abundant HFC, have been increasing by about 10% per 

year in recent years. Abundances of other HFCs, including HFC-125, -143a, -32, and -152a, have also been increasing. Regional studies suggest significant HFC emissions from Europe, Asia, and North America.

CFCs, HCFCs, HFCs, and Climate Change

•	 The Montreal Protocol and its Amendments and Adjustments have made large contributions toward reducing

global greenhouse gas emissions (Figure ES-1). In 2010, the decrease of annual ODS emissions under the Montreal 

Protocol is estimated to be about 10 gigatonnes of avoided CO2-equivalent1 emissions per year, which is about five times 

larger than the annual emissions reduction target for the first commitment period (2008–2012) of the Kyoto Protocol.

•	 The sum of the HFCs currently used as ODS replacements contributes about 0.4 gigatonnes of CO2-equivalent

per year to total global CO2-equivalent emissions, while the HCFCs contribute about 0.7 gigatonnes. CO2-

equivalent emissions of HFCs are increasing by about 8% per year and this rate is expected to continue to grow, while 

the contribution from HCFCs is expected to start decreasing in the next decade.

•	 Emissions of HFC-23, a by-product of HCFC-22 production, contributed about 0.2 gigatonnes of CO2-equivalent

1

 GWP-weighted emissions, also known as CO2-equivalent emissions, are defined as the amount of gas emitted multiplied by its 100-year 

Global Warming Potential (GWP).
ES.3

Executive Summary

per year in 2006–2008. HFC-23 is a particularly potent greenhouse gas with a lifetime of about 220 years. Its emissions 

have increased in the past decade despite global emissions reduction measures, including those covered by the Kyoto 

Protocol’s Clean Development Mechanism projects.

Total Chlorine and Bromine and Implications for Ozone Depletion

•	 Total chlorine has continued to decline from its 1990s peak values in both the troposphere and the stratosphere.

Total tropospheric bromine is decreasing from its peak values, which occurred comparatively recently, while

stratospheric bromine is no longer increasing.

•	 Relative declines in the sum of stratospheric chlorine and bromine from peak values are largest in midlatitudes

and smallest in Antarctica (refer to Figure S1-1 in the Scientific Summary of Chapter 1 of this Assessment). These 

declines are not as pronounced as observed in their tropospheric abundances. Differences between declines in the 

troposphere and different regions of the stratosphere are primarily associated with the time required for air to move 

from the troposphere to those regions. The relative declines are smallest in Antarctica primarily because the transport 

times to polar regions are the largest.

0

1

2

3

4

Megatonnes per year

Mass-Weighted Emissions

Without Montreal Protocol

HFCs

ODSs

HCFCs

CFCs-only

Without accelerated

HCFC phaseout

high

low

0.0

0.5

1.0

1.5

2.0

Megatonnes CFC-11-eq per year

ODP-Weighted

Emissions

Without Montreal

Protocol

ODSs

HCFCs

CFCs, 

halons, others

1950 1970 1990 2010 2030 2050

Year

0

10

20

30

Gigatonnes CO2-eq per year

HFCs

ODSs

HCFCs

GWP-Weighted Emissions

Without Montreal Protocol

IPCC-SRES CO2 range

Magnitude of

Kyoto Protocol reduction target

 2.0

high

low

CFCs, 

halons, others

Figure ES-1. Emissions of ODSs and their substitutes. Global 

emissions of ODSs (CFCs, halons, HCFCs, and others) and their 

non-ozone depleting substitutes (HFCs) from 1950 to 2050. Emissions are the total from developing and developed countries. The 

legends identify the specific groups of substances included in each 

panel. The high and low HFC labels identify the upper and lower 

limits, respectively, in global baseline scenarios. The blue hatched 

regions indicate the emissions that would have occurred, in the 

absence of the Montreal Protocol, with 2–3% annual production 

increases in all ODSs.

Top panel: Global mass-weighted emissions expressed as megatonnes per year. The yellow dashed line shows HCFC emissions 

calculated without the provisions of the 2007 accelerated HCFC 

phase-out under the Montreal Protocol.

Middle panel: Global Ozone Depletion Potential-weighted emissions expressed as megatonnes of CFC-11-equivalent per year.

The emissions of individual gases are multiplied by their respective ODPs (CFC-11 = 1) to obtain aggregate, equivalent CFC-11 

emissions. The dashed line marks 1987, the year of the Montreal 

Protocol signing.

Bottom panel: Global GWP-weighted emissions expressed as 

gigatonnes of CO2-equivalent per year. The emissions of individual 

gases are multiplied by their respective GWPs (direct, 100-year time 

horizon; CO2 = 1) to obtain aggregate, equivalent CO2 emissions. 

Shown for reference are emissions for the range of CO2 scenarios from the Intergovernmental Panel on Climate Change (IPCC) 

Special Report on Emission Scenarios (SRES). The CO2 emissions 

for 1950–2007 are from global fossil fuel use and cement production. Beyond 2007, the shaded region for CO2 reflects the maximum 

(A1B) and minimum (B2) SRES scenarios. The dashed line marks 

2010, the middle year of the first commitment period of the Kyoto 

Protocol. Also shown is the magnitude of the reduction target of the 

first commitment period of the Kyoto Protocol, which is based on a 

1990–2010 projection of global greenhouse gas emission increases 

and the reduction target for participating countries.
ES.4

Executive Summary

OZONE AND CLIMATE: ANTARCTIC

The Antarctic ozone hole is the clearest manifestation of the effect of ODSs on the ozone layer. The depletion far 

exceeds natural variability and has occurred without exception since 1980. The ozone hole also provides the most visible 

example of how ozone depletion affects surface climate.

•	 Springtime Antarctic total column ozone losses (the ozone hole), first recognizable around 1980, continue to

occur every year (Figure ES-2c). Although the ozone losses exhibit year-to-year variations that are primarily driven 

by year-to-year changes in meteorology, October mean column ozone within the vortex has been about 40% below 

1980 values for the past fifteen years. The average erythemal (“sunburning”) UV measured at the South Pole between 

1991 and 2006 was 55–85% larger than the estimated values for the years 1963–1980.

•	 Doubts raised since the previous Assessment regarding our understanding of the cause of the Antarctic ozone

hole have been dispelled. New laboratory measurements on the key chemistry involved in polar ozone depletion 

have reaffirmed that past changes in ODSs are indeed the cause of the ozone hole. This is also supported by quantification of the chemicals responsible for the ozone hole via field observations.

•	 There is increased evidence that the Antarctic ozone hole has affected the surface climate in the Southern

Hemisphere. Climate models demonstrate that the ozone hole is the dominant driver of the observed changes in 

surface winds over the Southern Hemisphere mid and high latitudes during austral summer. These changes have 

contributed to the observed warming over the Antarctic Peninsula and cooling over the high plateau. The changes in 

the winds have also been linked to regional changes in precipitation, increases in sea ice around Antarctica, warming 

of the Southern Ocean, and a local decrease in the ocean sink of CO2.

•	 The trends in the summertime winds in the Southern Hemisphere are not expected to persist over the next few

decades. This is because of the expected offsetting influences on the surface winds of increasing greenhouse gases 

and the recovering ozone hole.

•	 Observed Antarctic springtime column ozone does not yet show a statistically significant increasing trend (Figure ES-2c). Year-to-year variability, due to meteorology, is much larger than the expected response to the small ODS 

decreases in the Antarctic vortex to date. This is consistent with simulations using chemistry-climate models (CCMs).

•	 The evolution of Antarctic springtime column ozone over the rest of the century is expected to be dominated by

the decline in ODS abundance (Figure ES-2c). CCM simulations show that greenhouse gas changes have had, and 

will continue to have, a small impact on the ozone hole compared to the effects of the ODS changes. There are some 

indications that small episodic Antarctic ozone holes may occur even through the end of the century. In spring and 

early summer, Antarctica will continue to experience excess surface UV.

OZONE AND CLIMATE: GLOBAL AND ARCTIC

As a result of the controls introduced by the Montreal Protocol and its Amendments and Adjustments, it is expected 

that the decline in ODSs will lead to an increase in stratospheric ozone abundances. However, it will be challenging to 

attribute ozone increases to the decreases in ODSs during the next few years because of natural variability, observational 

uncertainty, and confounding factors, such as changes in stratospheric temperature or water vapor. A feature of this 

Assessment is the coordinated use by the community of chemistry-climate models (CCMs) with integrations covering 

the period from 1960–2100, which has allowed more detailed study of the long-term changes in the stratosphere and of 

the relative contributions of ODSs and greenhouse gases (GHGs).

•	 Average total ozone values in 2006–2009 remain at the same level as the previous Assessment, at roughly 3.5%

and 2.5% below the 1964–1980 averages respectively for 90°S–90°N and 60°S–60°N. Midlatitude (35°–60°) 

annual mean total column ozone amounts in the Southern Hemisphere [Northern Hemisphere] over the period 2006–

2009 have remained at the same level as observed during 1996–2005, at ~6% [~3.5%] below the 1964–1980 average.

•	 The ozone loss in Arctic winter and spring between 2007 and 2010 has been variable, but has remained in a

range comparable to the values prevailing since the early 1990s. Substantial chemical loss continues to occur 

during cold Arctic winters.
ES.5

Executive Summary

Figure ES-2. Schematic of the influence of ozonedepleting substances (ODSs) and climate change 

on the stratospheric ozone layer, and the influence 

of ozone changes on surface ultraviolet radiation.

The red lines are based on observations to date. The 

blue dashed lines represent one commonly accepted 

scenario for the future. Shaded areas represent yearto-year variability and uncertainties in simulations of 

the past and future. The dashed vertical line at 1980, a 

year used as a benchmark for ozone and UV recovery, 

demarcates the situation before and after significant 

changes to the ozone layer. The curve for carbon dioxide (CO2), a greenhouse gas important to Earth’s climate, is shown because its changes can affect stratospheric temperatures as well as wind patterns, both of 

which affect stratospheric ozone.

(a) Combined effective abundance of ozone-depleting chlorine and bromine in the lower atmosphere 

(troposphere). The red line is a representation 

of the measured atmospheric abundances. The 

blue dashed line is the expected combined effective abundance of chlorine and bromine based on 

the most likely ODS scenario used in this report 

and current understanding of the workings of the 

atmosphere. Because of the Montreal Protocol, a 

continued decline is expected through the end of 

this century, with a return to the 1980 benchmark 

value occurring around the middle of this century. A similar curve for the stratosphere would 

be shifted to the right (later dates) by a few years 

because of the time lag in the transport of substances from the surface to the stratosphere.

(b) The atmospheric abundance of carbon dioxide, 

the major anthropogenic greenhouse gas that 

changes Earth’s climate, including in the stratosphere; CO2 abundance is a proxy for climate change. The gray dotted/shaded 

area represents expectations of increasing future CO2 abundance based on different scenarios used in this Assessment.

(c) The extent of the Antarctic ozone hole, as measured by the amount of ozone in the total overhead column averaged for 

the month of October. The ozone hole is the clearest indicator of ozone layer depletion by ODSs, and the ODSs in the atmosphere have been and are expected to continue to be the primary control on the extent and duration of the ozone hole. 

Antarctic ozone is expected to return to pre-1980 benchmark values in the late 21st century. The blue shaded area shows 

the estimated year-to-year variability of ozone for one scenario that includes changes in ODSs (panel a), CO2 (panel b, blue 

dashed line), and changes in nitrous oxide and methane (not shown), but does not capture all uncertainties. The gray dotted/

shaded area shows the uncertainty due to different climate scenarios, but again does not capture all uncertainties.

(d) The extent of northern midlatitude ozone depletion, as measured by the amount of ozone in the total overhead column 

between 30°N and 60°N averaged over each year; blue and gray shaded areas as in panel c. Panels c and d show the approximate relative magnitudes of the northern midlatitude ozone depletion and the Antarctic ozone hole. Influences of the 

quasi-biennial oscillation (QBO), volcanoes, and solar cycle have been removed from the observational data. The future projections do not include the influence of any volcanic eruptions or solar cycle variations. Natural variability makes it difficult to 

identify the projected return of northern midlatitude ozone levels to pre-1980 levels, but the expectation is that climate change 

will hasten this return by several decades, such that it will occur before the middle of the 21st century (before the return of 

stratospheric chlorine and bromine to the 1980 benchmark value, and before the return of Antarctic ozone, panel c).

(e) Changes in clear-sky surface UV radiation at northern midlatitudes that accompany the ozone changes of the ODS scenario 

above. Because the ozone depletion in the northern midlatitudes has been small, the UV changes also have been small. 

The blue shaded area shows the year-to-year variability of surface UV for the ozone changes of panel d. Clouds, aerosols, 

and air pollution significantly affect surface UV, but it is difficult to project their future changes. The uncertainties in these 

changes, which are larger than the uncertainties due to ozone changes, are not represented in the figure. The expectation 

is that climate change will result in northern midlatitude clear-sky surface UV radiation levels well below 1980 values by the 

second half of this century.

(c) Antarctic total ozone in October

(e) Northern midlatitude surface ultraviolet radiation

(b) Carbon dioxide

1980 now ~2100

(a) Ozone-depleting chlorine and bromine in the lower atmosphere

(d) Northern midlatitude total ozone
ES.6

Executive Summary

•	 Robust linkages between Arctic stratospheric ozone depletion and tropospheric and surface climate trends

have not been established, as expected from the smaller ozone depletion compared with the Antarctic.

•	 Chemistry-climate models reproduce both the latitudinal and vertical structure of the observed ozone trends

in both northern and southern midlatitudes during the past periods of increase of the ODSs, confirming our

basic understanding of ozone change. Simulations agree with observations that the last decade has shown flattening 

of the time series of global total ozone.

•	 Analyses based on surface and satellite measurements show that erythemal UV irradiance over midlatitudes

has increased since the late 1970s (Figure ES-2e). This is in qualitative agreement with the observed decrease in 

column ozone, although other factors (mainly clouds and aerosols) have influenced long-term changes in erythemal 

irradiance. Clear-sky UV observations from unpolluted sites in midlatitudes show that since the late 1990s, UV irradiance levels have been approximately constant, consistent with ozone column observations over this period.

•	 New analyses of both satellite and radiosonde data give increased confidence in changes in stratospheric temperatures between 1980 and 2009. The global-mean lower stratosphere cooled by 1–2 K and the upper stratosphere 

cooled by 4–6 K between 1980 and 1995. There have been no significant long-term trends in global-mean lower 

stratospheric temperatures since about 1995. The global-mean lower-stratospheric cooling did not occur linearly but 

was manifested as downward steps in temperature in the early 1980s and the early 1990s. The cooling of the lower 

stratosphere includes the tropics and is not limited to extratropical regions as previously thought.

•	 The evolution of lower stratospheric temperature is influenced by a combination of natural and human factors

that has varied over time. Ozone decreases dominate the lower stratospheric cooling since 1980. Major volcanic 

eruptions and solar activity have clear shorter-term effects. Models that consider all of these factors are able to 

reproduce this temperature time history.

•	 Changes in stratospheric ozone, water vapor, and aerosols all radiatively affect surface temperature. The 

radiative forcing2 of climate in 2008 due to stratospheric ozone depletion (−0.05 ± 0.1 Watts per square meter, W/m2) 

is much smaller than the positive radiative forcing due to the CFCs and HCFCs largely responsible for that depletion 

(about +0.3 W/m2). For context, the current forcing by CO2 is approximately +1.7 W/m2. Radiative calculations 

and climate modeling studies suggest that the radiative effects of variability in stratospheric water vapor (± ~0.1 W/

m2 per decade) can contribute to decadal variability in globally averaged surface temperature. Climate models and 

observations show that major volcanic eruptions (e.g., Mt. Pinatubo in 1991, roughly −3 W/m2) can cool the surface 

for several years.

•	 The global middle and upper stratosphere are expected to cool in the coming century, mainly due to CO2

increases. Stratospheric ozone recovery will slightly offset the cooling. HFCs could warm the tropical lower stratosphere and tropopause region by about 0.3°C if stratospheric abundances reach the 1 ppb level.

•	 Emerging evidence from model simulations suggests that increasing greenhouse gases lead to an acceleration

of the stratospheric circulation usually referred to as the Brewer-Dobson circulation. Such an acceleration 

could have important consequences, particularly decreases in column ozone in the tropics and increases in column 

ozone elsewhere. However, responsible mechanisms remain unclear and observational evidence for the circulation 

increase is lacking.

•	 Global ozone is projected to increase approximately in line with the ODS decline, and the increase is accelerated by cooling of the upper stratosphere. Global ozone is not very sensitive to circulation changes, so high 

confidence can be placed in this projection.

•	 The evolution of ozone in the Arctic is projected to be more sensitive to climate change than in the Antarctic.

The projected strengthening of the stratospheric Brewer-Dobson circulation is expected to significantly increase 

lower stratospheric ozone in the Arctic, augmenting the GHG-induced ozone increase from upper stratospheric cooling and hastening the return to 1980 levels.

2

 Positive radiative forcings generally warm the surface; negative radiative forcings generally cool the surface.
ES.7

Executive Summary

•	 GHG-induced temperature and circulation changes are projected to hasten the return of midlatitude total

column ozone to 1980 levels by several decades, rising well above 1980 levels by the end of the century. The 

effect is most pronounced in northern midlatitudes (Figure ES-2d), where it would result in clear-sky surface UV 

radiation levels well below 1980 values by the second half of the century (Figure ES-2e). In southern midlatitudes, 

the effect of circulation changes is projected to be weaker and ozone is also influenced by depletion in the Antarctic, 

where the return to 1980 levels occurs much later.

INFORMATION FOR POLICYMAKERS AND OPTIONS FOR POLICY FORMULATION

Cases related to the elimination of future emissions, production, and banks for various ozone-depleting substances 

(ODSs) can be formulated starting from a baseline future emission scenario. The baseline scenario here has been developed to account for past and present levels of ODSs along with emission projections. This scenario projects that stratospheric chlorine and bromine levels are likely to return to 1980 levels in midcentury for the midlatitudes and about 25 years 

later in the Antarctic vortex. These additional cases are used to evaluate the impact of various hypothetical policy options.

Information for Policymakers

•	 The Montreal Protocol has both protected the ozone layer and provided substantial co-benefits by reducing

climate change (see Figure ES-1, bottom two panels). It has protected the stratospheric ozone layer by phasing out 

production and consumption of ozone-depleting substances. Simulations show that unchecked growth in the emissions of ODSs would have led to global ozone depletion in the coming decades very much larger than current levels. 

Solar UV radiation at the surface would also have increased substantially.

•	 Projections of hydrofluorocarbon (HFC) growth in scenarios that assume no controls suggest that by 2050,

Global Warming Potential–weighted emissions from these substances could be comparable to the GWPweighted emissions of chlorofluorocarbons (CFCs) at their peak in 1988 (see Figure ES-1, bottom panel). The 

highest projection assumes that developing countries use HFCs with GWPs comparable to those currently in use.

•	 The accelerated hydrochlorofluorocarbon (HCFC) phase-out agreed to by the Parties to the Montreal Protocol

in 2007 is projected to reduce ozone depletion and to help reduce climate forcing (see Figure ES-1). This acceleration is expected to reduce cumulative HCFC emissions by about 0.7 million Ozone Depletion Potential–tonnes 

between 2011 and 2050 and would bring forward the year equivalent effective stratospheric chlorine (EESC) returns 

to 1980 levels by 4–5 years. The accelerated HCFC phasedown is projected to reduce greenhouse gas emissions by 

about 0.5 gigatonnes of carbon dioxide (CO2)-equivalent per year averaged over 2011 through 2050. The projected 

benefit would be determined by the climate impact of the replacements. In comparison, global anthropogenic emissions of CO2 were greater than 30 gigatonnes per year in 2008.

•	 Since the previous Assessment, new fluorocarbons have been suggested as possible replacements for potent

HCFC and HFC greenhouse gases. For example, HFC-1234yf (Ozone Depletion Potential (ODP) = 0; 100-year 

GWP = 4) is proposed to replace HFC-134a (ODP = 0; 100-year GWP = 1370) in mobile air conditioning. To fully 

assess the environmental impacts, each proposed substance would need to be evaluated for its ODP, GWP, atmospheric fate, safety, and toxicity. Preliminary analyses indicate that global replacement of HFC-134a with HFC1234yf at today’s level of use is not expected to contribute significantly to tropospheric ozone formation or produce 

harmful levels of the degradation product TFA (trifluoroacetic acid). It is well established that TFA is a ubiquitous 

component of the environment, but uncertainties remain regarding its natural and anthropogenic sources, long-term 

fate, and abundances.

Due to the success of the Montreal Protocol and its Amendments and Adjustments in reducing the production, emissions, and abundances of controlled ODSs, emissions from other compounds and activities not controlled by the Montreal 

Protocol have become relatively more important to stratospheric ozone.
ES.8

Executive Summary

•	 Increasing abundances of radiatively important gases, especially carbon dioxide (CO2) and methane (CH4), are

expected to significantly affect future stratospheric ozone through effects on temperature, winds, and chemistry. CO2 increased in the atmosphere at 2.1 parts per million per year from 2005–2008, while CH4 increased by 

about 6.7 parts per billion per year from 2006–2008.

•	 Nitrous oxide (N2O) is known to both deplete global ozone and warm the climate. The current ODP-weighted

anthropogenic emission is larger than that of any ODS.

•	 Deliberate large injections of sulfur-containing compounds into the stratosphere, which have been suggested

as a climate intervention approach (geoengineering), would alter the radiative, dynamical, and chemical state

of the stratosphere and could be expected to have substantial unintended effects on stratospheric ozone levels.

Options for Policy Formulation

Additional cases have been developed to show the impact of further control measures on various substances. Table 

ES-1 shows the percentage reductions in integrated chlorine and bromine levels and integrated GWP-weighted emissions, 

relative to the baseline scenario, that can be achieved in these hypothetical cases.

•	 Halons and CFCs: Leakage from banks is the largest source of current ODP-weighted emissions of ODSs. A 

delay in the capture and destruction of estimated CFC banks from 2011 to 2015 is currently thought to reduce the 

possible ozone and climate benefits that could be achieved by about 30%.

•	 Carbon tetrachloride (CCl4): Elimination of future CCl4 emissions after 2010 would have an EESC impact comparable to the capture and destruction of CFC and halon banks. This is a much larger effect than was estimated in 

the previous Assessment because of a revision in the estimated emissions.

•	 HCFCs: The recent growth in reported HCFC production in developing countries was larger than projected in the 

previous Assessment. This alone would have resulted in a larger projected HCFC production in the new baseline 

scenario compared to the previous Assessment, but is expected to be more than compensated for by the accelerated 

HCFC phasedown agreed to by the Parties to the Montreal Protocol in 2007.

•	 Elimination of all emissions of chlorine- and bromine-containing ODSs after 2010: This would bring forward 

the return of EESC to 1980 levels by about 13 years. The elimination of these ODS emissions would have a climate 

impact equivalent to about a 0.7 gigatonnes of CO2-equivalent per year reduction from 2011 through 2050, on average. The sum of current banks of CFCs plus HCFCs contributes about the same amount to these CO2-equivalent 

emissions as future HCFC production.

•	 Methyl bromide: Two methyl bromide cases were examined. Case 1: A phase-out of quarantine and pre-shipment 

emissions beginning in 2011 would accelerate the return of EESC to 1980 levels by 1.5 years, relative to a case of 

maintaining emissions at 2004–2008 average levels. Case 2: Continuing critical-use exemptions at the approved 

2011 level indefinitely would delay the return of EESC to 1980 levels by 0.2 year.
ES.9

Executive Summary

Table ES-1. Hypothetical cases. Reductions in integrated chlorine and bromine levels (as measured by 

equivalent effective stratospheric chlorine, EESC) and integrated GWP-weighted emissions, relative to the 

baseline scenario, that can be achieved in hypothetical cases developed to show the impact of further control 

measures on various substances.

Substance or Group of

Substances

Reductions (%) in

Integrated EESC

(equivalent effective

stratospheric chlorine)

Reduction in Cumulative

GWP-Weighted Emissions

from 2011 to 2050

(gigatonnes of

CO2-equivalent)

Bank capture and 

destruction in 2011 and 2015:

2011 2015 2011 2015

CFCs 11 7.0 7.9 5.5

Halons 14 9.1 0.4 0.3

HCFCs 4.8 5.3 1 4.9 5.5 1

Production elimination after 2010:

HCFCs 8.8 13.2

CH3Br for quarantine and pre-shipment 6.7 0.002

Total emissions elimination after 2010:

CCl4

 2 7.6 0.9

CH3CCl3 0.1 0.004

HFCs 0.0 Up to 170 3

1 The impact of a 2015 HCFC bank recovery is larger than a 2011 bank recovery because this calculation assumes destruction of the bank in only a 

single year, and because the bank in 2015 is larger than the bank in 2011 owing to continued annual production that is larger than the annual bank 

release. 2 Banks are assumed to be zero. Emissions include uncertain sources such as possible fugitive emissions and unintended by-product emissions. 3 Strongly dependent on future projections and does not consider HFC-23 emissions. Currently HFCs are not controlled by the Montreal Protocol, but 

are included in the basket of gases of the Kyoto Protocol.
Chapter 1

Ozone-Depleting Substances (ODSs) and

Related Chemicals

Coordinating Lead Authors: Contributors:

S.A. Montzka E. Atlas

S. Reimann P. Bernath

T. Blumenstock

Lead Authors: J.H. Butler

A. Engel A. Butz

K. Krüger B. Connor

S. O’Doherty P. Duchatelet

W.T. Sturges G. Dutton

F. Hendrick

Coauthors: P.B. Krummel

D. Blake L.J.M. Kuijpers

M. Dorf E. Mahieu

P. Fraser A. Manning

L. Froidevaux J. Mühle

K. Jucks K. Pfeilsticker

K. Kreher B. Quack

M.J. Kurylo M. Ross

A. Mellouki R.J. Salawitch

J. Miller S. Schauffler

O.-J. Nielsen I.J. Simpson

V.L. Orkin D. Toohey

R.G. Prinn M.K. Vollmer

R. Rhew T.J. Wallington

M.L. Santee H.J.R. Wang

A. Stohl R.F. Weiss

D. Verdonik M. Yamabe

Y. Yokouchi

S. Yvon-Lewis

Chapter 1

OZONE-DEPLETING SUBSTANCES (ODSs) AND RELATED CHEMICALS

Contents

SCIENTIFIC SUMMARY .............................................................................................................................................1

1.1 SUMMARY OF THE PREVIOUS OZONE ASSESSMENT................................................................................7

1.2 LONGER-LIVED HALOGENATED SOURCE GASES ......................................................................................7

1.2.1 Updated Observations, Trends, and Emissions.........................................................................................7

1.2.1.1 Chlorofluorocarbons (CFCs) ......................................................................................................7

Box 1-1. Methods for Deriving Trace Gas Emissions............................................................................14

1.2.1.2 Halons.......................................................................................................................................15

1.2.1.3 Carbon Tetrachloride (CCl4).....................................................................................................16

Box 1-2. CCl4 Lifetime Estimates...........................................................................................................18

1.2.1.4 Methyl Chloroform (CH3CCl3).................................................................................................19

1.2.1.5 Hydrochlorofluorocarbons (HCFCs) ........................................................................................20

1.2.1.6 Methyl Bromide (CH3Br) .........................................................................................................23

1.2.1.7 Methyl Chloride (CH3Cl)..........................................................................................................27

Box 1-3. Atmospheric Lifetimes and Removal Processes.....................................................................34

1.2.2 Loss Processes.........................................................................................................................................35

1.3 VERY SHORT-LIVED HALOGENATED SUBSTANCES (VSLS)..................................................................37

1.3.1 Emissions, Atmospheric Distributions, and Abundance Trends of Very Short-Lived Source Gases.....37

1.3.1.1 Chlorine-Containing Very Short-Lived Source Gases .............................................................37

Box 1-4. Definition of Acronyms Related to Short-Lived Gases..........................................................39

1.3.1.2 Bromine-Containing Very Short-Lived Source Gases .............................................................41

1.3.1.3 Iodine-Containing Very Short-Lived Source Gases.................................................................44

1.3.1.4 Halogen-Containing Aerosols...................................................................................................44

1.3.2 Transport of Very Short-Lived Substances into the Stratosphere...........................................................44

1.3.2.1 VSLS Transport from the Surface in the Tropics to the Tropical Tropopause Layer (TTL) ...45

1.3.2.2 VSLS Transport from the TTL to the Stratosphere ..................................................................46

1.3.2.3 VSLS Transport from the Surface to the Extratropical Stratosphere .......................................46

1.3.3 VSLS and Inorganic Halogen Input to the Stratosphere .........................................................................47

1.3.3.1 Source Gas Injection (SGI).......................................................................................................47

1.3.3.2 Product Gas Injection (PGI)......................................................................................................49

1.3.3.3 Total Halogen Input into the Stratosphere from VSLS and Their Degradation Products........51

1.3.4 Potential Influence of VSLS on Ozone ...................................................................................................53

1.3.5 The Potential for Changes in Stratospheric Halogen from Naturally Emitted VSLS.............................54

1.3.6 Environmental Impacts of Anthropogenic VSLS, Substitutes for Long-Lived ODSs, and HFCs .........54

1.3.6.1 Evaluation of the Impact of Intensified Natural Processes on Stratospheric Ozone ................55

1.3.6.2 Very Short-Lived New ODSs and Their Potential Influence on Stratospheric Halogen..........55

1.3.6.3 Evaluation of Potential and In-Use Substitutes for Long-Lived ODSs....................................55

1.4 CHANGES IN ATMOSPHERIC HALOGEN......................................................................................................63

1.4.1 Chlorine in the Troposphere and Stratosphere........................................................................................63

1.4.1.1 Tropospheric Chlorine Changes ...............................................................................................63

1.4.1.2 Stratospheric Chlorine Changes................................................................................................64

1.4.2 Bromine in the Troposphere and Stratosphere........................................................................................66

1.4.2.1 Tropospheric Bromine Changes ...............................................................................................66
1.4.2.2 Stratospheric Bromine Changes................................................................................................67

1.4.3 Iodine in the Upper Troposphere and Stratosphere.................................................................................73

1.4.4 Equivalent Effective Chlorine (EECl) and Equivalent Effective Stratospheric Chlorine (EESC) .........73

1.4.5 Fluorine in the Troposphere and Stratosphere ........................................................................................75

1.5 CHANGES IN OTHER TRACE GASES THAT INFLUENCE OZONE AND CLIMATE ...............................75

1.5.1 Changes in Radiatively Active Trace Gases that Directly Influence Ozone...........................................76

1.5.1.1 Methane (CH4)..........................................................................................................................76

1.5.1.2 Nitrous Oxide (N2O).................................................................................................................79

1.5.1.3 COS, SO2, and Sulfate Aerosols...............................................................................................80

1.5.2 Changes in Radiative Trace Gases that Indirectly Influence Ozone.......................................................81

1.5.2.1 Carbon Dioxide (CO2) ..............................................................................................................81

1.5.2.2 Fluorinated Greenhouse Gases .................................................................................................82

1.5.3 Emissions of Rockets and Their Impact on Stratospheric Ozone ...........................................................85

REFERENCES .............................................................................................................................................................86
1.1

ODSs and Related Chemicals

SCIENTIFIC SUMMARY

The amended and adjusted Montreal Protocol continues to be successful at reducing emissions and atmospheric abundances of most controlled ozone-depleting substances (ODSs).

Tropospheric Chlorine

•	 Total tropospheric chlorine from long-lived chemicals (~3.4 parts per billion (ppb) in 2008) continued to 

decrease between 2005 and 2008. Recent decreases in tropospheric chlorine (Cl) have been at a slower rate than in 

earlier years (decreasing at 14 parts per trillion per year (ppt/yr) during 2007–2008 compared to a decline of 21 ppt/

yr during 2003–2004) and were slower than the decline of 23 ppt/yr projected in the A1 (most likely, or baseline) 

scenario of the 2006 Assessment. The tropospheric Cl decline has recently been slower than projected in the A1 

scenario because chlorofluorocarbon-11 (CFC-11) and CFC-12 did not decline as rapidly as projected and because 

increases in hydrochlorofluorocarbons (HCFCs) were larger than projected.

•	 The contributions of specific substances or groups of substances to the decline in tropospheric Cl have changed 

since the previous Assessment. Compared to 2004, by 2008 observed declines in Cl from methyl chloroform 

(CH3CCl3) had become smaller, declines in Cl from CFCs had become larger (particularly CFC-12), and increases in 

Cl from HCFCs had accelerated. Thus, the observed change in total tropospheric Cl of −14 ppt/yr during 2007–2008 

arose from:

• −13.2 ppt Cl/yr from changes observed for CFCs

• −6.2 ppt Cl/yr from changes observed for methyl chloroform

• −5.1 ppt Cl/yr from changes observed for carbon tetrachloride

• −0.1 ppt Cl/yr from changes observed for halon-1211

• +10.6 ppt Cl/yr from changes observed for HCFCs

•	 Chlorofluorocarbons (CFCs), consisting primarily of CFC-11, -12, and -113, accounted for 2.08 ppb (about 

62%) of total tropospheric Cl in 2008. The global atmospheric mixing ratio of CFC-12, which accounts for about 

one-third of the current atmospheric chlorine loading, decreased for the first time during 2005–2008 and by mid-2008 

had declined by 1.3% (7.1 ± 0.2 parts per trillion, ppt) from peak levels observed during 2000–2004.

•	 Hydrochlorofluorocarbons (HCFCs), which are substitutes for long-lived ozone-depleting substances, 

accounted for 251 ppt (7.5%) of total tropospheric Cl in 2008. HCFC-22, the most abundant of the HCFCs, 

increased at a rate of about 8 ppt/yr (4.3%/yr) during 2007–2008, more than 50% faster than observed in 2003–2004 

but comparable to the 7 ppt/yr projected in the A1 scenario of the 2006 Assessment for 2007–2008. HCFC-142b mixing ratios increased by 1.1 ppt/yr (6%/yr) during 2007–2008, about twice as fast as was observed during 2003–2004 

and substantially faster than the 0.2 ppt/yr projected in the 2006 Assessment A1 scenario for 2007–2008. HCFC141b mixing ratios increased by 0.6 ppt/yr (3%/yr) during 2007–2008, which is a similar rate observed in 2003–2004 

and projected in the 2006 Assessment A1 scenario.

•	 Methyl chloroform (CH3CCl3) accounted for only 32 ppt (1%) of total tropospheric Cl in 2008, down from a 

mean contribution of about 10% during the 1980s.

•	 Carbon tetrachloride (CCl4) accounted for 359 ppt (about 11%) of total tropospheric Cl in 2008. Mixing ratios 

of CCl4 declined slightly less than projected in the A1 scenario of the 2006 Assessment during 2005–2008.

Stratospheric Chlorine and Fluorine

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

inorganic chlorine continued to decline during 2005–2008. This burden agrees within ±0.3 ppb (±8%) 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 contributions from shorter-lived source gases and product gases of 80 (40–130) 
1.2

Chapter 1

ppt. Declines since 1996 in total column and stratospheric abundances of inorganic chlorine compounds are reasonably consistent with the observed trends in long-lived source gases over this period.

•	 Measured column abundances of hydrogen fluoride increased during 2005–2008 at a smaller rate than in earlier years. This is qualitatively consistent with observed changes in tropospheric fluorine (F) from CFCs, HCFCs, 

hydrofluorocarbons (HFCs), and perfluorocarbons (PFCs) that increased at a mean annual rate of 40 ± 4 ppt/yr (1.6 ± 

0.1%/yr) since late 1996, which is reduced from 60–100 ppt/yr observed during the 1980s and early 1990s.

Tropospheric Bromine

•	 Total organic bromine from controlled ODSs continued to decrease in the troposphere and by mid-2008 was 

15.7 ± 0.2 ppt, approximately 1 ppt below peak levels observed in 1998. This decrease was close to that expected 

in the A1 scenario of the 2006 Assessment and was driven by declines observed for methyl bromide (CH3Br) that 

more than offset increased bromine (Br) from halons.

•	 Bromine from halons stopped increasing during 2005–2008. Mixing ratios of halon-1211 decreased for the first 

time during 2005–2008 and by mid-2008 were 0.1 ppt below levels observed in 2004. Halon-1301 continued to 

increase in the atmosphere during 2005–2008 but at a slower rate than observed during 2003–2004. The mean rate 

of increase was 0.03–0.04 ppt/yr during 2007–2008. A decrease of 0.01 ppt/yr was observed for halon-2402 in the 

global troposphere during 2007–2008.

•	 Tropospheric methyl bromide (CH3Br) mixing ratios continued to decline during 2005–2008, and by 2008 had 

declined by 1.9 ppt (about 20%) from peak levels measured during 1996–1998. Evidence continues to suggest 

that this decline is the result of reduced industrial production, consumption, and emission. This industry-derived 

emission is estimated to have accounted for 25–35% of total global CH3Br emissions during 1996–1998, before 

industrial production and consumption were reduced. Uncertainties in the variability of natural emissions and in the 

magnitude of methyl bromide stockpiles in recent years limit our understanding of this anthropogenic emissions fraction, which is derived by comparing the observed atmospheric changes to emission changes derived from reported 

production and consumption.

•	 By 2008, nearly 50% of total methyl bromide consumption was for uses not controlled by the Montreal 

Protocol (quarantine and pre-shipment applications). From peak levels in 1996–1998, industrial consumption in 

2008 for controlled and non-controlled uses of CH3Br had declined by about 70%. Sulfuryl fluoride (SO2F2) is used 

increasingly as a fumigant to replace methyl bromide for controlled uses because it does not directly cause ozone 

depletion, but it has a calculated direct, 100-year Global Warming Potential (GWP100) of 4740. The SO2F2 global 

background mixing ratio increased during recent decades and had reached about 1.5 ppt by 2008.

Stratospheric Bromine

•	 Total bromine in the stratosphere was 22.5 (19.5–24.5) ppt in 2008. It is no longer increasing and by some 

measures has decreased slightly during recent years. Multiple measures of stratospheric bromine monoxide (BrO) 

show changes consistent with tropospheric Br trends derived from observed atmospheric changes in CH3Br and the 

halons. Slightly less than half of the stratospheric bromine derived from these BrO observations is from controlled 

uses of halons and methyl bromide. The remainder comes from natural sources of methyl bromide and other bromocarbons, and from quarantine and pre-shipment uses of methyl bromide not controlled by the Montreal Protocol.

Very Short-Lived Halogenated Substances (VSLS)

VSLS are defined as trace gases whose local lifetimes are comparable to, or shorter than, tropospheric transport 

timescales and that have non-uniform tropospheric abundances. In practice, VSLS are considered to be those compounds 

having atmospheric lifetimes of less than 6 months.
1.3

ODSs and Related Chemicals

•	 The amount of halogen from a very short-lived source substance that reaches the stratosphere depends on the 

location of the VSLS emissions, as well as atmospheric removal and transport processes. Substantial uncertainties remain in quantifying the full impact of chlorine- and bromine-containing VSLS on stratospheric ozone. 

Updated results continue to suggest that brominated VSLS contribute to stratospheric ozone depletion, particularly 

under enhanced aerosol loading. It is unlikely that iodinated gases are important for stratospheric ozone loss in the 

present-day atmosphere.

•	 Based on a limited number of observations, very short-lived source gases account for 55 (38–80) ppt chlorine 

in the middle of the tropical tropopause layer (TTL). From observations of hydrogen chloride (HCl) and carbonyl 

chloride (COCl2) in this region, an additional ~25 (0–50) ppt chlorine is estimated to arise from VSLS degradation. 

The sum of contributions from source gases and these product gases amounts to ~80 (40–130) ppt chlorine from 

VSLS that potentially reaches the stratosphere. About 40 ppt of the 55 ppt of chlorine in the TTL from source gases 

is from anthropogenic VSLS emissions (e.g., methylene chloride, CH2Cl2; chloroform, CHCl3; 1,2 dichloroethane, 

CH2ClCH2Cl; perchloroethylene, CCl2CCl2), but their contribution to stratospheric chlorine loading is not well 

quantified.

•	 Two independent approaches suggest that VSLS contribute significantly to stratospheric bromine. Stratospheric 

bromine derived from observations of BrO implies a contribution of 6 (3–8) ppt of bromine from VSLS. Observed, 

very short-lived source gases account for 2.7 (1.4–4.6) ppt Br in the middle of the tropical tropopause layer. By 

including modeled estimates of product gas injection into the stratosphere, the total contribution of VSLS to stratospheric bromine is estimated to be 1–8 ppt.

•	 Future climate changes could affect the contribution of VSLS to stratospheric halogen and its influence on 

stratospheric ozone. Future potential use of anthropogenic halogenated VSLS may contribute to stratospheric halogen in a similar way as do present-day natural VSLS. Future environmental changes could influence both anthropogenic and natural VSLS contributions to stratospheric halogens.

Equivalent Effective Stratospheric Chlorine (EESC)

EESC is a sum of chlorine and bromine derived from ODS tropospheric abundances weighted to reflect their potential 

influence on ozone in different parts of the stratosphere. The growth and decline in EESC varies in different regions of the atmosphere because a given tropospheric abundance propagates to the stratosphere with varying time lags associated with transport. 

Thus the EESC abundance, when it peaks, and how much it has declined from its peak vary in different regions of the atmosphere.

0%

50%

60%

70%

80%

90%

100%

110%

1980 1985 1990 1995 2000 2005 2010

EESC abundance 

(relative to the peak)

Midlatitude stratosphere

Polar stratosphere

-10%

-28%

% return to

1980 level

by the end

of 2008

Date of peak in the 

troposphere

1980 levels

Year 

Figure S1-1. Stratospheric EESC derived 

for the midlatitude and polar stratospheric 

regions relative to peak abundances, plotted as a function of time. Peak abundances 

are ~1950 ppt for the midlatitude stratosphere and ~4200 ppt for the polar stratosphere. Percentages shown to the right indicate the observed change in EESC by the 

end of 2008 relative to the change needed 

for EESC to return to its 1980 abundance. 

A significant portion of the 1980 EESC level 

is from natural emissions.
1.4

Chapter 1

•	 EESC has decreased throughout the stratosphere.

• By the end of 2008, midlatitude EESC had decreased by about 11% from its peak value in 1997. This drop is 

28% of the decrease required for EESC in midlatitudes (red curve in figure) to return to the 1980 benchmark 

level.

• By the end of 2008, polar EESC had decreased by about 5% from its peak value in 2002. This drop is 10% of 

the decrease required for EESC in polar regions (blue curve in figure) to return to the 1980 benchmark level.

•	 During the past four years, no specific substance or group of substances dominated the decline in the total 

combined abundance of ozone-depleting halogen in the troposphere. In contrast to earlier years, the long-lived 

CFCs now contribute similarly to the decline as do the short-lived CH3CCl3 and CH3Br. Other substances contributed 

less to this decline, and HCFCs added to this halogen burden over this period.

Emission Estimates and Lifetimes

•	 While global emissions of CFC-12 derived from atmospheric observations decreased during 2005–2008, those 

for CFC-11 did not change significantly over this period. Emissions from banks account for a substantial fraction 

of current emissions of the CFCs, halons, and HCFCs. Emissions inferred for CFCs from global observed changes 

did not decline during 2005–2008 as rapidly as projected in the A1 scenario of the 2006 Assessment, most likely 

because of underestimates of bank emissions.

•	 Global emissions of CCl4 have declined only slowly over the past decade.

• These emissions, when inferred from observed global trends, were between 40 and 80 gigagrams per year (Gg/

yr) during 2005–2008 given a range for the global CCl4 lifetime of 33–23 years. By contrast, CCl4 emissions 

derived with a number of assumptions from data reported to the United Nations Environment Programme 

(UNEP) ranged from 0–30 Gg/yr over this same period.

• In addition, there is a large variability in CCl4 emissions derived from data reported to UNEP that is not 

reflected in emissions derived from measured global mixing ratio changes. This additional discrepancy cannot be explained by scaling the lifetime or by uncertainties in the atmospheric trends. If the analysis of data 

reported to UNEP is correct, unknown anthropogenic sources may be partly responsible for these observed 

discrepancies.

•	 Global emissions of HCFC-22 and HCFC-142b derived from observed atmospheric trends increased during 

2005–2008. HCFC-142b global emissions increased appreciably over this period, compared to a projected emissions 

decline of 23% from 2004 to 2008. By 2008, emissions for HCFC-142b were two times larger than had been projected 

in the A1 scenario of the 2006 Assessment. These emission increases were coincident with increasing production of 

HCFCs in developing countries in general and in East Asia particularly. It is too soon to discern any influence of the 

2007 Adjustments to the Montreal Protocol on the abundance and emissions of HCFCs.

•	 The sum of CFC emissions (weighted by direct, 100-year GWPs) has decreased on average by 8 ± 1%/yr from 

2004 to 2008, and by 2008 amounted to 1.1 ± 0.3 gigatonnes of carbon dioxide-equivalent per year (GtCO2-

eq/yr). The sum of GWP-weighted emissions of HCFCs increased by 5 ± 2%/yr from 2004 to 2008, and by 2008 

amounted to 0.74 ± 0.05 GtCO2-eq/yr.

•	 Evidence is emerging that lifetimes for some important ODSs (e.g., CFC-11) may be somewhat longer than 

reported in past assessments. In the absence of corroborative studies, however, the CFC-11 lifetime reported in this 

Assessment remains unchanged at 45 years. Revisions in the CFC-11 lifetime would affect estimates of its global 

emission derived from atmospheric changes and calculated values for Ozone Depletion Potentials (ODPs) and bestestimate lifetimes for some other halocarbons.
1.5

ODSs and Related Chemicals

Other Trace Gases That Directly Affect Ozone and Climate

•	 The methane (CH4) global growth rate was small, averaging 0.9 ± 3.3 ppb/yr between 1998–2006, but increased 

to 6.7 ± 0.6 ppb/yr from 2006–2008. Analysis of atmospheric data suggests that this increase is due to wetland 

sources in both the high northern latitudes and the tropics. The growth rate variability observed during 2006–2008 

is similar in magnitude to that observed over the last two decades.

•	 In 2005–2008 the average growth rate of nitrous oxide (N2O) was 0.8 ppb/yr, with a global average tropospheric mixing ratio of 322 ppb in 2008. A recent study has suggested that at the present time, Ozone Depletion 

Potential-weighted anthropogenic emissions of N2O are the most significant emissions of a substance that depletes 

ozone.

•	 Long-term changes in carbonyl sulfide (COS) measured as total columns above the Jungfraujoch (46.5°N) and 

from surface flasks sampled in the Northern Hemisphere show that atmospheric mixing ratios have increased 

slightly during recent years concurrently with increases in “bottom-up” inventory-based emissions of global 

sulfur. Results from surface measurements show a mean global surface mixing ratio of 493 ppt in 2008 and a mean 

rate of increase of 1.8 ppt/yr during 2000–2008. New laboratory, observational, and modeling studies indicate that 

vegetative uptake of COS is significantly larger than considered in the past.

Other Trace Gases with an Indirect Influence on Ozone

•	 The carbon dioxide (CO2) global average mixing ratio was 385 parts per million (ppm) in 2008 and had 

increased during 2005–2008 at an average rate of 2.1 ppm/yr. This rate is higher than the average growth rate 

during the 1990s of 1.5 ppm/yr and corresponds with increased rates of fossil fuel combustion.

•	 Hydrofluorocarbons (HFCs) used as ODS substitutes continued to increase in the global atmosphere. HFC134a is the most abundant HFC; its global mixing ratio reached about 48 ppt in 2008 and was increasing at 4.7 ppt/

yr. Other HFCs have been identified in the global atmosphere at <10 ppt (e.g., HFC-125, -143a, -32, and -152a) and 

were increasing at ≤1 ppt/yr in 2008.

•	 Emissions of HFC-23, a by-product of HCFC-22 production, have increased over the past decade even as 

efforts at minimizing these emissions were implemented in both developed and developing countries. These 

emission increases are concurrent with rapidly increasing HCFC-22 production in developing countries and are 

likely due to increasing production of HCFC-22 in facilities not covered by the Kyoto Protocol’s Clean Development 

Mechanism projects. Globally averaged HFC-23 mixing ratios reached 21.8 ppt in 2008, with a yearly increase of 

0.8 ppt/yr (3.9%/yr).

•	 The sum of emissions (weighted by direct, 100-year GWPs) of HFCs used as ODS replacements has increased 

by 8–9%/yr from 2004 to 2008, and by 2008 amounted to 0.39 ± 0.03 GtCO2-eq/yr. Regional studies suggest 

significant contributions of HFC-134a and -152a emissions during 2005–2006 from Europe, North America, and 

Asia. Emissions of HFC-23, most of which do not arise from use of this substance as an ODS replacement, added an 

additional 0.2 Gt CO2-eq/yr, on average, during 2006–2008.

•	 Sulfur hexafluoride (SF6) and nitrogen trifluoride (NF3): Global averaged mixing ratios of SF6 reached 6.4 ppt in 

2008, with a yearly increase of 0.2 ppt/yr. NF3 was detected in the atmosphere for the first time, with a global mean 

mixing ratio in 2008 of 0.45 ppt and a growth rate of 0.05 ppt/yr, or 11%/yr.
1.6

Chapter 1

Direct Radiative Forcing

The abundances of ODSs as well as many of their replacements contribute to radiative forcing of the atmosphere. 

These climate-related forcings have been updated using the current observations of atmospheric abundances and are 

summarized in Table S1-1. This table also contains the primary Kyoto Protocol gases as reference.

•	 Over these 5 years, radiative forcing from the sum of ODSs and HFCs has increased but, by 2008, remained 

small relative to the forcing changes from CO2 (see Table S1-1).

Table S1-1. Direct radiative forcings of ODSs and other gases, and their recent changes.

Specific Substance or Group

of Substances

Direct Radiative Forcing

(2008), milliWatts per

square meter (mW/m2)

Change in Direct Radiative

Forcing (2003.5–2008.5),

mW/m2

CFCs * 262 −6

Other ODSs * 15 −2

HCFCs * 45 8

HFCs #,a 12 5

HFC-23 # 4 0.9

CO2

 # 1740 139

CH4

 # 500 4

N2O # 170 12

PFCs # 5.4 0.5

SF6

 # 3.4 0.7

Sum of Montreal Protocol gases * 322 0

Sum of Kyoto Protocol gases # 2434 163

*

 Montreal Protocol Gases refers to CFCs, other ODSs (CCl4, CH3CCl3, halons, CH3Br), and HCFCs. #

 Kyoto Protocol Gases (CO2, CH4, N2O, HFCs, PFCs, and SF6). a

 Only those HFCs for which emissions arise primarily through use as ODS replacements (i.e., not HFC-23).
1.7

ODSs and Related Chemicals

1.1 SUMMARY OF THE PREVIOUS 

OZONE ASSESSMENT

The 2006 Assessment report (WMO, 2007) documented the continued success of the Montreal Protocol in 

reducing the atmospheric abundance of ozone-depleting 

substances (ODSs). Tropospheric abundances and emissions of most ODSs were decreasing by 2004, and tropospheric chlorine (Cl) and bromine (Br) from ODSs were 

decreasing as a result. Methyl chloroform contributed 

more to the decline in tropospheric chlorine than other 

controlled gases. ODS substitute chemicals containing 

chlorine, the hydrofluorochlorocarbons (HCFCs), were 

still increasing during 2000–2004, but at reduced rates 

compared to earlier years.

A significant mismatch between expected and 

atmosphere-derived emissions of carbon tetrachloride 

(CCl4) was identified. For the first time a decline was observed in the stratospheric burden of inorganic Cl as measured both by ground- and space-based instrumentation. 

The amount and the trend observed for stratospheric chlorine was consistent with abundances and trends of longlived ODSs observed in the troposphere, though lag times 

and mixing complicated direct comparisons.

Tropospheric bromine from methyl bromide and 

halons was determined in the previous Assessment to be 

decreasing. Changes derived for stratospheric inorganic 

bromine (Bry) from observations of BrO were consistent 

with tropospheric trends measured from methyl bromide 

and the halons, but it was too early to detect a decline 

in stratospheric Bry. Amounts of stratospheric Bry were 

higher than expected from the longer-lived, controlled 

gases (methyl bromide and halons). This suggested a significant contribution of 5 (3–8) parts per trillion (ppt) of Br 

potentially from very short-lived substances (VSLS) with 

predominantly natural sources. Large emissions of very 

short-lived brominated substances were found in tropical 

regions, where rapid transport from Earth’s surface to the 

stratosphere is possible. Quantitatively accounting for this 

extra Br was not straightforward given our understanding 

at that time of timescales and heterogeneity of VSLS emissions and oxidation product losses as these compounds 

become transported from Earth’s surface to the stratosphere. It was concluded that this additional Br has likely 

affected stratospheric ozone levels, and the amount of Br 

from these sources would likely be sensitive to changes 

in climate that affect ocean conditions, atmospheric loss 

processes, and atmospheric circulation.

By 2004, equivalent effective chlorine (EECl), a 

simple metric to express the overall effect of these changes on ozone-depleting halogen abundance, continued to 

decrease. When based on measured tropospheric changes 

through 2004, EECl had declined then by an amount that 

was 20% of what would be needed to return EECl values to those in 1980 (i.e., before the ozone hole was observed).

In the past, the discussions of long-lived and shortlived compounds were presented in separate chapters but 

are combined in this 2010 Assessment. 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, are all considered to be 

equivalent.

1.2 LONGER-LIVED HALOGENATED 

SOURCE GASES

1.2.1 Updated Observations, Trends, 

and Emissions

1.2.1.1 Chlorofluorocarbons (CFCs)

The global surface mean mixing ratios of the three 

most abundant chlorofluorocarbons (CFCs) declined significantly during 2005–2008 (Figure 1-1 and Table 1-1). 

After reaching its peak abundance during 2000–2004, 

the global annual surface mean mixing ratio of CFC-12 

(CCl2F2) had declined by 7.1 ± 0.2 ppt (1.3%) by mid2008. Surface means reported for CFC-12 in 2008 by the 

three independent global sampling networks (532.6–537.4 

ppt) agreed to within 5 ppt (0.9%). The consistency for 

CFC-12 among these networks has improved since the 

previous Assessment and stems in part from a calibration 

revision in the National Oceanic and Atmospheric Administration (NOAA) data. The 2008 annual mean mixing ratio of CFC-11 (CCl3F) from the three global sampling networks (243.4–244.8 ppt) agreed to within 1.4 ppt (0.6%) 

and decreased at a rate of −2.0 ± 0.6 ppt/yr from 2007 to 

2008. Global surface means observed by these networks 

for CFC-113 (CCl2FCClF2) during 2008 were between 

76.4 and 78.3 ppt and had decreased from 2007 to 2008 at 

a rate of −0.7 ppt/yr.

Long-term CFC-11 and CFC-12 data obtained from 

ground-based infrared solar absorption spectroscopy are 

available from the Jungfraujoch station (Figure 1-2; an 

update of Zander et al., 2005). Measured trends in total 

vertical column abundances during 2001 to 2008 indicate 

decreases in the atmospheric burdens of these gases that 

are similar to the declines derived from the global sampling networks over this period. For example, the mean 

decline in CFC-11 from the Jungfraujoch station column 

data is −0.83(± 0.06)%/yr during 2001–2009 (relative to 

2001), and global and Northern Hemisphere (NH) surface 

trends range from −0.78 to −0.88%/yr over this same period (range of trends from different networks). For CFC-
1.8

Chapter 1

12, the rate of change observed at the Jungfraujoch station 

was −0.1(± 0.05)% during 2001–2008 (relative to 2001), 

while observed changes at the surface were slightly larger 

at −0.2%/yr over this same period.

Additional measurements of CFC-11 in the upper 

troposphere and stratosphere with near-global coverage 

have been made from multiple satellite-borne instruments 

(Kuell et al., 2005; Hoffmann et al., 2008; Fu et al., 2009). 

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CFC-11

HCFC-124

HCFC-22 HCFC-141b HCFC-142b

halon-2402

halon-1211 halon-1301

CH3CCl3 CH3Br

CFC-12 CFC-113

CCl4

Global surface mixing ratio (parts per trillion or ppt)

Year

Figure 1-1. Mean global surface mixing ratios (expressed as dry air mole fractions in parts per trillion or ppt) of 

ozone-depleting substances from independent sampling networks and from scenario A1 of the previous Ozone 

Assessment (Daniel and Velders et al., 2007) over the past 18 years. Measured global surface monthly means 

are shown as red lines (NOAA data) and blue lines (AGAGE data). Mixing ratios from scenario A1 from the 

previous Assessment (black lines) were derived to match observations in years before 2005 as they existed 

in 2005 (Daniel and Velders et al., 2007). The scenario A1 results shown in years after 2004 are projections 

made in 2005.
1.9

ODSs and Related Chemicals

Table 1-1. Measured mole fractions and growth rates of ozone-depleting gases from ground-based 

sampling.

Chemical 

Formula

Common or 

Industrial 

Name

Annual Mean

Mole Fraction (ppt)

Growth

(2007–2008) Network, Method

2004 2007 2008 (ppt/yr) (%/yr)

CFCs

CCl2F2 CFC-12 543.8 539.6 537.4 −2.2 −0.4 AGAGE, in situ (Global)

542.3 537.8 535.5 −2.3 −0.4 NOAA, flask & in situ (Global)

539.7 535.1 532.6 −2.5 −0.5 UCI, flask (Global)

541.5 541.2 541.0 −0.3 −0.05 NIES, in situ (Japan)

- 542.9 540.1 −2.9 −0.5 SOGE-A, in situ (China)

CCl3F CFC-11 251.8 245.4 243.4 −2.0 −0.8 AGAGE, in situ (Global)

253.8 247.0 244.8 −2.2 −0.9 NOAA, flask & in situ (Global) 

253.7 246.1 244.2 −1.9 −0.8 UCI, flask (Global)

253.6 247.7 247.6 −0.1 0.0 NIES, in situ (Japan)

254.7 247.4 244.9 −2.6 −1.1 SOGE, in situ (Europe)

- 246.8 245.0 −1.8 −0.7 SOGE-A, in situ (China)

CCl2FCClF2 CFC-113 79.1 77.2 76.5 −0.6 −0.8 AGAGE, in situ (Global)

81.1 78.9 78.3 −0.6 −0.8 NOAA, in situ (Global)

79.3 77.4 76.4 −1.0 −1.3 NOAA, flask (Global)

79.1 77.8 77.1 −0.7 −0.9 UCI, flask (Global)

79.7 78.1 78.0 −0.1 −0.1 NIES, in situ (Japan)

- 77.5 76.7 −0.8 −1.1 SOGE-A, in situ (China)

CClF2CClF2 CFC-114 16.6 16.5 16.4 −0.04 −0.2 AGAGE, in situ (Global)

16.2 16.4 16.2 −0.2 −1.3 UCI, flask (Global)

16.0 15.9 16.0 0.05 0.3 NIES, in situ (Japan)

- 16.7 - - - SOGE, in situ (Europe)

CClF2CF3 CFC-115 8.3 8.3 8.4 0.02 0.3 AGAGE, in situ (Global)

8.6 8.3 8.3 0.05 0.6 NIES, in situ (Japan)

8.3 8.5 8.5 0.0 0.0 SOGE, in situ (Europe)

HCFCs

CHClF2 HCFC-22 163.4 183.6 192.1 8.6 4.6 AGAGE, in situ (Global)

162.1 182.9 190.8 7.9 4.2 NOAA, flask (Global) 

160.0 180.7 188.3 7.6 4.2 UCI, flask (Global)

- 190.7 200.6 9.9 5.2 NIES, in situ (Japan)

- 197.3 207.3 10.0 5.0 SOGE-A, in situ (China)

CH3CCl2F HCFC-141b 17.5 18.8 19.5 0.7 3.6 AGAGE, in situ (Global)

17.2 18.7 19.2 0.5 2.6 NOAA, flask (Global) 

- 18.2 18.8 0.6 3.2 UCI, flask (Global)

- 20.2 21.2 0.9 4.6 NIES, in situ (Japan)

- 20.8 21.2 0.5 2.2 SOGE, in situ (Europe)

CH3CClF2 HCFC-142b 15.1 17.9 18.9 1.1 5.9 AGAGE, in situ (Global)

14.5 17.3 18.5 1.2 6.7 NOAA, flask (Global) 

- 17.0 18.0 1.0 5.7 UCI, flask (Global)
1.10

Chapter 1

Chemical 

Formula

Common or 

Industrial 

Name

Annual Mean

Mole Fraction (ppt)

Growth

(2007–2008) Network, Method

2004 2007 2008 (ppt/yr) (%/yr)

CH3CClF2 HCFC-142b - 18.9 20.2 1.3 6.5 NIES, in situ (Japan)

- 19.7 21.0 1.4 6.8 SOGE, in situ (Europe)

- 20.9 21.8 0.9 4.1 SOGE-A, in situ (China)

CHClFCF3 HCFC-124 1.43 1.48 1.47 −0.01 −0.8 AGAGE, in situ (Global)

- 0.81 0.80 −0.01 −1.2 NIES, in situ (Japan)

Halons

CBr2F2 halon-1202 0.038 0.029 0.027 −0.002 −7.0 UEA, flasks (Cape Grim only)

CBrClF2 halon-1211 4.37 4.34 4.30 −0.04 −0.9 AGAGE, in situ (Global)

4.15 4.12 4.06 −0.06 −1.4 NOAA, flasks (Global)

4.31 4.29 4.25 −0.04 −0.8 NOAA, in situ (Global)

- 4.30 4.23 −0.06 −1.4 UCI, flasks (Global)

4.62 4.50 4.40 −0.1 −2.0 SOGE, in situ (Europe)

- 4.40 4.31 −0.1 −2.0 SOGE-A, in situ (China)

4.77 4.82 4.80 −0.02 −0.4 UEA, flasks (Cape Grim only)

CBrF3 halon-1301 3.07 3.17 3.21 0.04 1.3 AGAGE, in situ (Global)

2.95 3.09 3.12 0.03 1.1 NOAA, flasks (Global) 

3.16 3.26 3.29 0.03 1.2 SOGE, in situ (Europe)

- 3.15 3.28 0.1 3.8 SOGE-A, in situ (China)

2.45 2.48 2.52 0.03 1.3 UEA, flasks (Cape Grim only)

CBrF2CBrF2 halon-2402 0.48 0.48 0.47 −0.01 −1.2 AGAGE, in situ (Global)

0.48 0.47 0.46 −0.01 −2.0 NOAA, flasks (Global)

0.43 0.41 0.40 −0.01 −1.2 UEA, flasks (Cape Grim only)

Chlorocarbons

CH3Cl methyl

chloride

533.7 541.7 545.0 3.3 0.6 AGAGE, in situ (Global)

545 550 - - - NOAA, in situ (Global)

537 548 547 −0.7 −0.1 NOAA, flasks (Global) 

526 541 547 5.9 1.1 SOGE, in situ (Europe)

CCl4 carbon

tetrachloride

92.7 89.8 88.7 −1.1 −1.3 AGAGE, in situ (Global)

95.7 92.3 90.9 −1.4 −1.5 NOAA, in situ (Global)

95.1 92.6 91.5 −1.1 −1.2 UCI, flask (Global)

- 90.2 88.9 −1.3 −1.5 SOGE-A, in situ (China)

CH3CCl3 methyl

chloroform

21.8 12.7 10.7 −2.0 −17.6 AGAGE, in situ (Global)

22.5 13.2 11.4 −1.9 −15.1 NOAA, in situ (Global)

22.0 12.9 10.8 −2.1 −17.8 NOAA, flasks (Global) 

23.9 13.7 11.5 −2.2 −17.5 UCI, flask (Global)

22.2 13.1 11.0 −2.2 −18.0 SOGE, in situ (Europe)

- 13.3 11.7 −1.6 −12.8 SOGE-A, in situ (China)

Table 1-1, continued.
1.11

ODSs and Related Chemicals

These results uniquely characterize the interhemispheric, 

interannual, and seasonal variations in the CFC-11 upperatmosphere distribution, though an analysis of the consistency in trends derived from these platforms and from 

surface data has not been performed.

The global mixing ratios of the two less abundant 

CFCs, CFC-114 (CClF2CClF2) and CFC-115 (CClF2CF3), 

have not changed appreciably from 2000 to 2008 (Table 

1-1) (Clerbaux and Cunnold et al., 2007). During 2008, 

global mixing ratios of CFC-114 were between 16.2 and 

16.4 ppt based on results from the Advanced Global Atmospheric Gases Experiment (AGAGE) and University 

of California-Irvine (UCI) networks, and AGAGE measurements show a mean global mixing ratio of 8.4 ppt for 

CFC-115 (Table 1-1). For these measurements, CFC-114 

measurements are actually a combination of CFC-114 and 

CFC-114a (see notes to Table 1-1).

Observed mixing ratio declines of the three most 

abundant CFCs during 2005–2008 were slightly slower 

than projected in scenario A1 (baseline scenario) from the 

2006 WMO Ozone Assessment (Daniel and Velders et al., 

2007) (Figure 1-1). The observed declines were smaller 

than projected during 2005–2008 in part because release 

rates from banks were underestimated in the A1 scenario 

Chemical 

Formula

Common or 

Industrial 

Name

Annual Mean

Mole Fraction (ppt)

Growth

(2007–2008) Network, Method

2004 2007 2008 (ppt/yr) (%/yr)

Bromocarbons

CH3Br methyl

bromide

8.2 7.7 7.5 −0.2 −2.7 AGAGE, in situ (Global)

7.9 7.6 7.3 −0.3 −3.6 NOAA, flasks (Global)

- 8.5 8.1 −0.4 −5.2 SOGE, in situ (Europe)

Notes:

Rates are calculated as the difference in annual means; relative rates are this same difference divided by the average over the two-year period. Results 

given in bold text and indicated as “Global” are estimates of annual mean global surface mixing ratios. Those indicated with italics are from a single 

site or subset of sites that do not provide a global surface mean mixing ratio estimate. Measurements of CFC-114 are a combination of CFC-114 and 

the CFC-114a isomer. The CFC-114a mixing ratio has been independently estimated as being ~10% of the CFC-114 mixing ratio (Oram, 1999) and 

has been subtracted from the results presented here assuming it has been constant over time.

These observations are updated from the following sources:

Butler et al. (1998), Clerbaux and Cunnold et al. (2007), Fraser et al. (1999), Maione et al. (2004), Makide and Rowland (1981), Montzka et al. (1999, 

2000, 2003, 2009), O’Doherty et al. (2004), Oram (1999), Prinn et al. (2000, 2005), Reimann et al. (2008), Rowland et al. (1982), Stohl et al. (2010), 

Sturrock et al. (2001), Reeves et al. (2005), Simmonds et al. (2004), Simpson et al. (2007), Xiao et al. (2010a, 2010b), and Yokouchi et al. (2006).

AGAGE, Advanced Global Atmospheric Gases Experiment; NOAA, National Oceanic and Atmospheric Administration; SOGE, System for Observation 

of halogenated Greenhouse gases in Europe; SOGE-A, System for Observation of halogenated Greenhouse gases in Europe and Asia; NIES, National 

Institute for Environmental Studies; UEA, University of East Anglia; UCI, University of California-Irvine.

Calendar Year

1986.0 1989.0 1992.0 1995.0 1998.0 2001.0 2004.0 2007.0

Column Abundance (x10 molecules/cm ) 15 2

1

2

3

5

6

7

8

CFC-12

HCFC-22

CFC-11

Pressure normalized monthly means

June to November monthly means

Polynomial fit to filled datapoints

NPLS fit (20%)

CFC-12, -11 and HCFC-22 above Jungfraujoch Figure 1-2. The time evolution of the 

monthly-mean total vertical column 

abundances (in molecules per square 

centimeter) of CFC-12, CFC-11, and 

HCFC-22 above the Jungfraujoch station, Switzerland, through 2008 (update 

of Zander et al., 2005). Note discontinuity in the vertical scale. Solid blue lines 

show polynomial fits to the columns 

measured only in June to November 

so as to mitigate the influence of variability caused by atmospheric transport 

and tropopause subsidence during winter and spring (open circles) on derived 

trends. Dashed green lines show nonparametric least-squares fits (NPLS) to 

the June to November data.

Table 1-1, continued.
1.12

Chapter 1

during this period (Daniel and Velders et al., 2007). For 

CFC-12, some of the discrepancy is due to revisions to 

the NOAA calibration scale. In the A1 scenario, CFC-11 

and CFC-12 release rates from banks were projected to 

decrease over time based on anticipated changes in bank 

sizes from 2002–2015 (IPCC/TEAP 2005). The updated 

observations of these CFCs, however, are more consistent 

with emissions from banks having been fairly constant 

during 2005–2008, or with declines in bank emissions being offset by enhanced emissions from non-bank-related 

applications. Implications of these findings are further 

discussed in Chapter 5 of this Assessment.

The slight underestimate of CFC-113 mixing ratios during 2005–2008, however, is not likely the result of 

inaccuracies related to losses from banks, since banks of 

CFC-113 are thought to be negligible (Figure 1-1). The 

measured mean hemispheric difference (North minus 

South) was ~0.2 ppt during 2005–2008, suggesting the 

potential presence of only small residual emissions (see 

Figure 1-4). The mean exponential decay time for CFC113 over this period is 100–120 years, slightly longer than 

the steady-state CFC-113 lifetime of 85 years. This observation is consistent with continuing small emissions 

(≤10 gigagrams (Gg) per year). Small lifetime changes 

are expected as atmospheric distributions of CFCs respond 

to emissions becoming negligible, but changes in the atmospheric distribution of CFC-113 relative to loss regions (the stratosphere) suggest that the CFC-113 lifetime 

should become slightly shorter, not longer, as emissions 

decline to zero (e.g., Prather, 1997).

CFC Emissions and Banks

Releases from banks account for a large fraction of 

current emissions for some ODSs and will have an important influence on mixing ratios of many ODSs in the 

future. Banks of CFCs were 7 to 16 times larger than 

amounts emitted in 2005 (Montzka et al., 2008). Implications of bank sizes, emissions from them, and their influence on future ODS mixing ratios are discussed further in 

Chapter 5.

Global, “top-down” emissions of CFCs derived 

from global surface observations and box models show 

rapid declines during the early 1990s but only slower 

changes in more recent years (Figure 1-3) (see Box 1-1 for 

a description of terms and techniques related to deriving 

emissions). Emission changes derived for CFC-11, for example, are small enough so that different model approaches (1-box versus 12-box) suggest either slight increases or 

slight decreases in emissions during 2005–2008. Considering the magnitude of uncertainties on these emissions, 

changes in CFC-11 emissions are not distinguishable from 

zero over this four-year period. “Bottom-up” estimates of 

emissions derived from production and use data have not 

been updated past 2003 (UNEP/TEAP, 2006), but projections made in 2005 indicated that CFC-11 emissions from 

banks of ~25 Gg/yr were not expected to decrease substantially from 2002 to 2008 (IPCC/TEAP, 2005) (Figure 1-3).

“Top-down” emissions derived for CFC-11 during 

2005–2008 averaged 80 Gg/yr. These emissions are larger 

than derived from “bottom-up” estimates by an average 

of 45 (37–60) Gg/yr over this same period. The discrepancy between the atmosphere-derived and “bottom-up” 

emissions for CFC-11 is not fully understood but could 

suggest an underestimation of releases from banks or fastrelease applications (e.g., solvents, propellants, or opencell foams). Emissions from such short-term uses were 

estimated at 15–26 Gg/yr during 2000–2003 (UNEP/

TEAP, 2006; Figure 1-3) and these accounted for a substantial fraction of total CFC-11 emissions during those 

years. The discrepancy may also arise from errors in the 

CFC-11 lifetime used to derive “top-down” emissions. 

New results from models that more accurately simulate air 

transport rates through the stratosphere suggest a steadystate lifetime for CFC-11 of 56–64 years (Douglass et al., 

2008), notably longer than 45 years. A relatively longer 

lifetime for CFC-12 was not suggested in this study. A 

longer CFC-11 lifetime of 64 years would bring the atmosphere-derived and “bottom-up” emissions into much 

better agreement (see light blue line in Figure 1-3).

Global emissions of CFC-12 derived from observed 

atmospheric changes decreased from ~90 to ~65 Gg/yr 

during 2005–2008 (Figure 1-3). These emissions and 

their decline from 2002–2008 are well accounted for by 

leakage from banks as projected in a 2005 report (IPCC/

TEAP, 2005). Global emissions of CFC-113 derived from 

observed global trends and 1-box or 12-box models and a 

global lifetime of 85 years were small compared to earlier 

years, and averaged <10 Gg/yr during 2005–2008 (Figure 

1-3).

Summed emissions from CFCs have declined during 2005–2008. When weighted by semi-empirical Ozone 

Depletion Potentials (ODPs) (Chapter 5), the sum of emissions from CFCs totaled 134 ± 30 ODP-Kt in 2008 (where 

1 kilotonne (Kt) = 1 Gg = 1 × 109 g). The sum of emissions of CFCs weighted by direct, 100-yr Global Warming 

Potentials (GWPs) has decreased on average by 8 ± 1%/

yr from 2004 to 2008, and by 2008 amounted to 1.1 ± 0.3 

gigatonnes of CO2-equivalents per year (Gt CO2-eq/yr).

Emission trends and magnitudes can also be 

inferred from measured hemispheric mixing ratio differences. This approach is valid when emissions are predominantly from the Northern Hemisphere and sink processes 

are symmetric between the hemispheres. Hemispheric 

mixing ratio differences for CFC-11 and CFC-12 averaged 

between 10 and 20 ppt during the 1980s when emissions 

were large, but since then as emissions declined, hemispheric differences also became smaller. United Nations 
1.13

ODSs and Related Chemicals

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HCFC-22 HCFC-141b HCFC-142b

halon-1211 halon-1301 halon-2402

CH3CCl3 CH3Br

CFC-12 CFC-113

CCl4

Global annual emissions (Gg/yr)

Year

Figure 1-3. “Top-down” and “bottom-up” global emission estimates for ozone-depleting substances (in Gg/

yr). “Top-down” emissions are derived with NOAA (red lines) and AGAGE (blue lines) global data and a 1-box 

model. These emissions are also derived with a 12-box model and AGAGE data (gray lines with uncertainties 

indicated) (see Box 1-1). Halon and HCFC emissions derived with the 12-box model in years before 2004 are 

based on an analysis of the Cape Grim Air Archive only (Fraser et al., 1999). A1 scenario emissions from the 

2006 Assessment are black lines (Daniel and Velders et al., 2007). “Bottom-up” emissions from banks (refrigeration, air conditioning, foams, and fire protection uses) are given as black plus symbols (IPCC/TEAP, 2005; 

UNEP, 2007a), and total, “bottom-up” emissions (green lines) including fast-release applications are shown for 

comparison (UNEP/TEAP, 2006). A previous bottom-up emission estimate for CCl4 is shown as a brown point 

for 1996 (UNEP/TEAP, 1998). The influence of a range of lifetimes for CCl4 (23–33 years) and a lifetime of 64 

years for CFC-11 are given as light blue lines.
1.14

Chapter 1

Box 1-1. Methods for Deriving Trace Gas Emissions

a) Emissions derived from production, sales, and usage (the “bottom-up” method). Global and national emissions of trace gases can be derived from ODS global production and sales magnitudes for different applications 

and estimates of application-specific leakage rates. For most ODSs in recent years, production is small or insignificant compared to historical levels and most emission is from material in use. Leakage and releases from 

this “bank” of material (produced but not yet emitted) currently dominate emissions for many ozone-depleting 

substances (ODSs). Uncertainties in these estimates arise from uncertainty in the amount of material in the bank 

reservoir and the rate at which material is released or leaks from the bank. Separate estimates of bank magnitudes 

and loss rates from these banks have been derived from an accounting of devices and appliances in use (IPCC/

TEAP, 2005). Emissions from banks alone account for most, if not all, of the “top-down,” atmosphere-derived 

estimates of total global emission for some ODSs (CFC-12, halon-1211, halon-1301, HCFC-22; see Figure 1-3).

b) Global emissions derived from observed global trends (the “top-down” method). Mass balance considerations allow estimates of global emissions for long-lived trace gases based on their global abundance, changes in 

their global abundance, and their global lifetime. Uncertainties associated with this “top-down” approach stem 

from measurement calibration uncertainty, imperfect characterization of global burdens and their change from 

surface observations alone, uncertain lifetimes, and modeling errors. The influence of sampling-related biases 

and calibration-related biases on derived emissions is small for most ODSs, given the fairly good agreement observed for emissions derived from different measurement networks (Figure 1-3). Hydroxyl radical (OH)-derived 

lifetimes are believed to have uncertainties on the order of ±20% for hydrochlorofluorocarbons (HCFCs), for 

example (Clerbaux and Cunnold et al., 2007). Stratospheric lifetimes also have considerable uncertainty despite 

being based on model calculations (Prinn and Zander et al., 1999) and observational studies (Volk et al., 1997). 

Recent improvements in model-simulated stratospheric transport suggest that the lifetime of CFC-11, for example, 

is 56–64 years instead of the current best estimate of 45 years (Douglass et al., 2008).

Global emissions derived for long-lived gases with different models (1-box and 12-box) show small differences in most years (Figure 1-3) (UNEP/TEAP, 2006). Though a simple 1-box approach has been used extensively in past Assessment reports, emissions derived with a 12-box model have also been presented. The 12-box 

model emissions estimates made here are derived with a Massachusetts Institute of Technology-Advanced Global 

Atmospheric Gases Experiment (MIT-AGAGE) code that incorporates observed mole fractions and a Kalman filter applying sensitivities of model mole fractions to 12-month semi-hemispheric emission pulses (Chen and Prinn, 

2006; Rigby et al., 2008). This code utilizes the information contained in both the global average mole fractions 

and their latitudinal gradients. Uncertainties computed for these annual emissions enable an assessment of the 

statistical significance of interannual emission variations.

c) Continental and global-scale emissions derived from measured global distributions. Measured mixing 

ratios (hourly through monthly averages) can be interpreted using inverse methods with global Eulerian threedimensional (3-D) chemical transport models (CTMs) to derive source magnitudes for long-lived trace gases 

such as methane (CH4), methyl chloride (CH3Cl), and carbon dioxide (CO2) on continental scales (e.g., Chen 

and Prinn, 2006; Xiao, 2008; Xiao et al., 2010a; Peylin et al., 2002; Rödenbeck et al., 2003; Meirink et al., 2008; 

Peters et al., 2007; Bergamaschi et al., 2009). Although much progress has been made with these techniques in 

recent years, some important obstacles limit their ability to retrieve unbiased fluxes. The first is the issue that the 

underdetermined nature of the problem (many fewer observations than unknowns) means that extra information, 

in the form of predetermined and prior constraints, is typically required to perform an inversion but can potentially 

impose biases on the retrieved fluxes (Kaminski et al., 2001). Second, all of these methods are only as good as the 

atmospheric transport models and underpinning meteorology they use. As Stephens et al. (2007) showed for CO2, 

biases in large-scale flux optimization can correlate directly with transport biases.

d) Regional-scale emissions derived from high-frequency data at sites near emission regions. High-frequency 

measurements (e.g., once per hour) near source regions can be used to derive regional-scale (~104–106 km2) trace 

gas emission magnitudes. The method typically involves interpreting measured mixing ratio enhancements above 

a background as an emissive flux using Lagrangian modeling concepts.
1.15

ODSs and Related Chemicals

Environment Programme consumption data suggest that 

CFC emissions continue to be dominated by releases in 

the Northern Hemisphere (UNEP, 2010). Furthermore, 

the small (0.2 ppt) hemispheric difference (North minus 

South) measured for CFC-113 since 2004—when emissions derived from atmospheric trends of this compound 

were very small (<10 Gg/yr)—indicates at most only a 

small contribution of loss processes to hemispheric differences for long-lived CFCs (Figure 1-4). By contrast, 

mean annual hemispheric differences for CFC-11 and 

CFC-12 have remained between 1.5 and 3 ppt since 2005 

and suggest the presence of continued substantial Northern 

Hemispheric emissions of these chemicals. For CFC-11, 

the hemispheric difference (North minus South) measured 

in both global networks has not declined appreciably since 

2005 (Figure 1-4), consistent with fairly constant emissions over that period (Figure 1-3).

Polar Firn and Volcano Observations

New CFC observations in firn air collected from the 

Arctic (two sites) and Antarctic (three sites) show small 

but detectable CFC-11, -12, and -114 levels increasing after ~1940 and CFC-113 and -115 appearing and increasing after ~1970 (Martinerie et al., 2009). These results 

add to conclusions from earlier firn-air studies (Butler 

et al., 1999; Sturrock et al., 2002) indicating that natural 

sources of CFCs, CCl4, CH3CCl3, halons, and HCFCs, if 

they exist, must be very small. Such conclusions rely on 

these compounds being stable in firn. Consistent with this 

result, studies of fumarole discharges from three Central 

American volcanoes over two years show that volcanic 

emissions are not a significant natural source of CFCs 

(Frische et al., 2006).

1.2.1.2 Halons

Updated observations show that the annual global 

surface mean mixing ratio of halon-1301 (CBrF3) increased at a rate of 0.03–0.04 ppt/yr during 2007–2008 

and reached 3.1–3.2 ppt in mid-2008 (Figure 1-1; Table 

1-1). Revised calibration procedures and reliance on gas 

chromatography with mass spectrometric detection analyses of flasks within NOAA have improved the consistency 

(within 5% in 2008) among results from independent laboratories compared to past reports (Clerbaux and Cunnold 

et al., 2007).

The global surface mean mixing ratio of halon1211 (CBrClF2) began to decrease during 2004–2005 and 

changed by −0.05 ± 0.01 ppt/yr during 2007–2008 (Figure 1-1; Table 1-1). The global surface mean in 2008 

was only about 0.1 ppt lower than peak levels measured 

in 2004.

Updated halon-2402 (CBrF2CBrF2) observations 

indicate that global surface mixing ratios declined from 

0.48 ppt in 2004 to 0.46–0.47 ppt in 2008 at a rate of −0.01 

ppt/yr in 2007–2008 (Table 1-1).

Updated halon-1202 (CBr2F2) measurements (Fraser et al., 1999; Reeves et al., 2005) show a substantial 

decrease in mixing ratios of this chemical since 2000. 

Southern Hemispheric mixing ratios decreased from 0.038 

ppt in 2004 to 0.027 ppt in 2008 at a rate of −0.002 ppt/yr 

in 2007–2008.

The observed changes in halon-1211 mixing ratios 

during 2005–2008 are similar to those projected in the A1 

scenario of the previous Assessment (Daniel and Velders 

et al., 2007); halon-1301 has increased at a slightly higher 

rate than projected. Observed surface mixing ratios of 

halon-2402 are notably higher than scenarios from past 

Qualitative indications of emission source regions or “hotspots” are provided by correlating observed mixing ratio enhancements with back trajectories (typically 4- to 10-day) calculated for the sampling time and location 

(e.g., Maione et al., 2008; Reimann et al., 2004; Reimann et al., 2008).

Quantitative emission magnitudes have been derived with the “ratio-method” (e.g., Dunse et al., 2005; 

Yokouchi et al., 2006; Millet et al., 2009; Guo et al., 2009). In this straightforward approach, trace-gas emissions 

are derived from correlations between observed enhancements above background for a trace gas of interest and a 

second trace gas (e.g., carbon monoxide or radon) whose emissions are independently known. Uncertainties in this 

approach are reduced when the emissions of the reference substance are well known, co-located, and temporally 

covarying with the halocarbon of interest, and when the mixing ratio enhancements of both chemicals are well correlated and large relative to uncertainties in the background levels.

More complex and powerful tools combine Lagrangian 3-D models and inverse methods to estimate regional emission fluxes (e.g., Manning et al., 2003; Stohl et al., 2009). As with larger-scale inversions, the challenge with these methods is that the inversion problem is underdetermined and results are dependent on transport 

accuracy. Furthermore, each station is sensitive to emissions only from a restricted region in its vicinity. To obtain 

global coverage of emissions from regional measurements, global transport models are used. Stohl et al. (2009) 

have recently developed a formal analytical method that takes into account a priori information for halocarbon 

emissions, which allows for regional-scale inversions with a global coverage.
1.16

Chapter 1

Assessments because those scenarios were not based on 

actual measurement data (Figure 1-1).

Halon Emissions, Stockpiles, and Banks

Stockpiles and banks of halons, which are used 

primarily as fire extinguishing agents, represent a substantial reservoir of these chemicals. The amounts 

of halons present in stockpiles or banks are not well 

quantified, but were estimated to be 15–33 times larger 

than emissions of halon-1211 and halon-1301 in 2008 

(UNEP, 2007a). “Bottom-up” estimates of halon emissions derived from production and use data were recently 

revised based on a reconsideration of historic release 

rates and the implications of this reanalysis on current 

bank sizes (UNEP, 2007a). The magnitude and trends 

in these emission estimates compare well with those derived from global atmospheric data and best-estimate 

lifetimes for halon-1211 and halon-1301 (Figure 1-3). 

“Bottom-up” emission estimates of halon-2402 are 

significantly lower than those derived from global atmospheric trends. Bank-related emissions are thought to 

account for nearly all halon emissions (plusses in Figure 

1-3). Halons are also used in small amounts in non-fire 

suppressant applications and as chemical feedstocks, but 

these amounts are not included in the “bottom-up” emissions estimates included in Figure 1-3.

Summed emissions of halons, weighted by semiempirical ODPs, totaled 90 ± 19 ODP-Kt in 2008. When 

weighted by 100-yr direct GWPs, summed halon emissions totaled 0.03 Gt CO2-eq in 2008.

1.2.1.3 Carbon Tetrachloride (CCl4)

The global mean surface mixing ratio of CCl4

continued to decrease during 2005–2008 (Figure 1-1). 

By 2008, the surface mean from the three global surface 

networks was approximately 90 ± 1.5 ppt and had de0

1

2

3

4

5

6

1990 1995 2000 2005 2010

NH – SH (ppt)

CFC-11 (N)

CFC-11 (A)

CCl4 (N)

CCl4 (A)

CFC-113 (N)

CFC-113 (A)

Year

Figure 1-4. Mean hemispheric mixing ratio differences (North minus South, in parts per trillion) measured for 

some ODSs in recent years from independent sampling networks (AGAGE data (A) as plusses, Prinn et al., 

2000; and NOAA data (N) as crosses, Montzka et al., 1999). Points are monthly-mean differences; lines are 

running 12-month means of the monthly differences.
1.17

ODSs and Related Chemicals

creased during 2007–2008 at a rate of −1.1 to −1.4 ppt/

yr (Table 1-1).

Global CCl4 Emissions

Though the global surface CCl4 mixing ratio decreased slightly faster during 2005–2008 than during 

2000–2004 (Figure 1-1), the observations imply only a 

slight decrease in CCl4 emissions over time (Figure 1-3). 

The measured global CCl4 mixing ratio changes suggest 

“top-down,” global emissions between 40 and 80 Gg/yr 

during 2005–2008 for a lifetime of 33–23 years (see Box 

1-2). Similar emission magnitudes and trends are derived 

for recent years from the independent global sampling 

networks and with different modeling approaches (Figure 

1-3). The decline observed for CCl4 mixing ratios during 

2005–2008 was slightly less rapid than that projected in 

the A1 scenario of the previous Assessment (Daniel and 

Velders et al., 2007), which was derived assuming a linear decline in emissions from 65 Gg/yr in 2004 to 0 Gg/

yr in 2015 and a 26-year lifetime (Daniel and Velders et 

al., 2007).

As with other compounds, “top-down” emissions 

for CCl4 are sensitive to errors in global lifetimes. A lifetime at the upper (or lower) end of the current uncertainty 

range yields a smaller (larger) emission than when calculated with the current best-estimate lifetime of 26 years 

(Figure 1-5). The magnitude of uncertainties that remain 

in the quantification of CCl4 sinks (i.e., stratosphere, 

ocean, and soil), however, do not preclude revisions to the 

CCl4 lifetime in the future that could significantly change 

the magnitude of “top-down” emissions.

Global CCl4 emission magnitudes and their trends 

also can be qualitatively assessed from measured hemispheric differences, which are roughly proportional to 

emissions for long-lived trace gases emitted primarily in 

the Northern Hemisphere (see Section 1.2.1.1). This difference has remained fairly constant for CCl4 at 1.25–1.5 

ppt over the past decade (Figure 1-4). These differences 

are independent of measured year-to-year changes in 

atmospheric mixing ratios and so provide a first-order 

consistency check on emission magnitudes and trends. 

Though the hemispheric difference (NH minus SH) expected for CCl4 in the absence of emissions is not well 

defined, it is expected to be small because the asymmetry in loss fluxes between the hemispheres due to oceans 

and soils is likely small (<10 Gg/yr) and offsetting. This 

analysis suggests, based on the considerations discussed 

above for CFCs, that the significant and sustained NH minus SH difference observed for CCl4 mixing ratios arises 

from substantial NH CCl4 emissions that have not changed 

substantially over the past decade (Figure 1-3).

In contrast to the CCl4 emissions derived from 

“top-down” methods, those derived with “bottom-up” 

techniques suggest substantially smaller emissions in most 

years. In the past two Assessment reports, for example, 

the “bottom-up” emissions estimate for CCl4 during 1996 

was 41 (31–62) Gg/yr (UNEP/TEAP 1998), or 20–50 Gg/

yr lower than the “top-down,” atmosphere-derived estimate for that year (Montzka and Fraser et al., 2003; Clerbaux and Cunnold et al., 2007). Because similar estimates 

have not been made for subsequent years, we derive here 

“potential emissions” from the difference between CCl4

production magnitudes in excess of amounts used as feedstock and amounts destroyed that are reported to UNEP 

(UNEP, 2010) (for the European Union (EU), feedstock 

magnitudes were determined from numbers reported by 

the individual EU countries and not from the EU as a 

whole). An upper limit to these “potential emissions” was 

also derived from this same data by filling apparent gaps 

in production and feedstock data reported to UNEP, including a 2% fugitive emissions rate from quantities used 

as feedstock, and incorporating an efficiency for reported 

CCl4 destruction of only 75% (Figure 1-5). This approach 

yields emissions of 0–30 Gg/yr during 2005–2008. As is 

apparent from this figure, CCl4 continues to be produced 

in substantial quantities (reported production in 2008 was 

156 Gg), but the primary use of this production is for feedstocks (e.g., in the production of some hydrofluorocarbons 

(HFCs), see Table 1-11, and other chemicals) that should 

yield only small emissions.

This “bottom-up” “potential emission” history derived from reported production, feedstock, and destruction 

data is inconsistent with the magnitude and variability in 

the “top-down,” atmosphere-derived emissions. Importantly, these differences cannot be reconciled by a simple 

scaling of the CCl4 atmospheric lifetime (Figure 1-5). This 

is particularly true during 2003–2008, when large declines 

are derived for “bottom-up” “potential emissions” but 

are not suggested by the atmosphere-derived “top-down” 

estimates or by the relatively constant NH minus SH difference measured during these years. The discrepancies 

during 2005–2008 between the “top-down” and “potential 

emission” estimates are between 30 and 50 Gg/yr. Even 

when the upper limits to “potential emissions” are compared to the lower range of “top-down” emissions (lifetime 

= 33 years), a discrepancy during 2005–2008 of 15–30 Gg/

yr remains (Figure 1-5). Though the magnitude of this discrepancy is sensitive to the uncertain CCl4 lifetime (Box 

1-2), the different time-dependent changes implied for 

emissions particularly during 2003–2008 with the different 

estimation methods are not sensitive to this lifetime.

Regional Studies

Observational studies can potentially provide information about regional source distributions and magnitudes 

(Box 1-1). Recently, Xiao et al. (2010b) have inversely 
1.18

Chapter 1

Box 1-2. CCl4 Lifetime Estimates

The loss of carbon tetrachloride (CCl4) in the global atmosphere is dominated by photolytic destruction in the 

stratosphere but also occurs by surface ocean uptake and uptake by soils. The atmospheric lifetime due to photolysis 

is estimated to be 35 years based on older modeling data and measured gradients in the lower stratosphere (Prinn and 

Zander et al., 1999; Volk et al., 1997). This number is based in part on the atmospheric lifetime of CFC-11. For 

example, in an analysis of global CCl4 distributions measured by satellite, Allen et al. (2009) derived a lifetime for 

CCl4 of 34 ± 5 years relative to a lifetime for CFC-11 of 45 years. An updated analysis of model-derived lifetimes, 

however, indicates that current models that more accurately simulate different stratospheric metrics (age of air, for 

example) calculate a substantially longer stratospheric lifetime for CFC-11 of 56–64 years (Douglass et al., 2008). Although CCl4 was not explicitly considered in the Douglass et al. study, the results could suggest a longer CCl4 partial 

lifetime with respect to stratospheric loss of ~ 44–50 yr. A recent independent analysis provides further evidence that 

the CCl4 stratospheric lifetime may be as long as 50 years (Rontu Carlon et al., 2010).

Undersaturations of CCl4 have been observed in many different regions of the world’s ocean waters and are indicative of a CCl4 sink (Wallace et al., 1994; Lee et al., 1999; Huhn et al., 2001; Yvon-Lewis and Butler, 2002; Tanhua 

and Olsson, 2005). These undersaturations are larger than can be explained from laboratory-determined hydrolysis 

rates and the responsible loss mechanism is not known. In the absence of new results, the partial lifetime with respect 

to oceanic loss of 94 (82–191) years (Yvon-Lewis and Butler, 2002) remains unchanged.

Losses of atmospheric CCl4 to a subset of terrestrial biomes have been observed in independent studies published since the previous Assessment (Liu, 2006; Rhew et al., 2008). These results confirm that terrestrial biomes can 

act as a sink for CCl4, but the magnitude of this loss remains uncertain. Losses to tropical soils account for most of the 

calculated soil sink (62%), but are based primarily on results from one tropical forest (Happell and Roche, 2003) and 

have yet to be remeasured. The new flux estimates reported for temperate forest and temperate grassland areas were 

derived with soil gas gradient methods and, in semi-arid and arid shrublands, with flux chamber methods. In these remeasured biomes, Liu (2006) and Rhew et al. (2008) found a mean sink about half as strong as in the original Happell 

and Roche (2003) study. When a range of partial lifetimes with respect to soil losses and other losses is considered, a 

mid-range estimate for CCl4 lifetime remains at 26 (23–33) years (see Table 1).

Box 1-2 Table 1. Sensitivity of total global CCl4 lifetime to component sink magnitudes.

Partial Lifetimes (years) with

Respect to Loss to:

Stratosphere Ocean Soil Total Lifetime Description and Notes

35 ∞ ∞ 35 pre-2003 Ozone 

Assessments

35 94 ∞ 26 post-2003 Ozone 

Assessments1

35 94 90 20 a

35 94 101 20 b

35 94 195 23 c

44–50 94 ∞ 30–33 d

44–50 94 195 26–28 c & d

44–50 94 101 23–25 b & d

Notes:

1

 Montzka and Fraser et al. (2003); Clerbaux and Cunnold et al. (2007).

(a) Partial lifetime from loss to soils derived from Happell and Roche (2003) from measurements in seven different biomes.

(b) Rhew et al. (2008) and Liu (2006) estimates of loss to arid land, temperate forest, and temperate grasslands soils used instead of those estimated 

by Happell and Roche (2003), and Happell and Roche (2003) loss estimates for soils in the other four biomes.

(c) All soil losses in Happell and Roche (2003) were scaled by 0.5 to account for updated results (Rhew et al. (2008) and Liu (2006)) in three of 

the seven biomes originally studied being only half as large, on average, as originally found.

(d) A longer stratospheric CCl4 lifetime (of 50 years) is considered based on a CFC-11 lifetime of 56-64 years rather than 45 years (Douglass et 

al., 2008) (44 yr = 35×56/45; 50 yr = 35×64/45), consistent with the recent work of Rontu Carlon et al. (2010).
1.19

ODSs and Related Chemicals

estimated regional and global annual CCl4 emissions and 

sinks using the three-dimensional Model of Atmospheric 

Transport and Chemistry (3D-MATCH) model, a monthly 

applied Kalman filter, a priori industrial emission patterns 

for 8 regions in the world, and observed monthly-mean 

mixing ratios from 1996–2004 at multiple, globally distributed AGAGE and NOAA/Earth System Research Laboratory (ESRL) sites. The average 1996–2004 East Asian 

(including China) emissions accounted for 53.3 ± 3.6% 

of the global total industrial emissions during this period. 

The fraction of global emissions inferred from South Asia 

(including India) were estimated at 22.5 ± 3.0%, those 

from Africa at 9.0 ± 1.2%, those from North America at 

6.6 ± 1.9%, and those from Europe at 4.0 ± 2.2%.

Regional emissions of CCl4 have also been estimated from measured mixing ratio enhancements in pollution events near source regions. These studies have 

suggested small or no detectable (ND) emissions from 

North America during 2003–2006 (Millet et al., 2009: 

ND; Hurst et al., 2006: <0.6 Gg/yr), Australia (Dunse 

et al., 2005: none detected in most years during 1995–

2000), and Japan (Yokouchi et al., 2005: 1.8 Gg/yr in 

2001 decreasing to 0.27 Gg/yr in 2003). Larger emissions have been inferred as coming from China (Vollmer 

et al., 2009: 15 (10–22) Gg/yr for the period Oct 2006–

March 2008). While these studies and those of Xiao et al. 

(2010b) point to countries in South East Asia (including 

China) as providing a large fraction of CCl4 emissions 

in the past, it is not possible to gauge their consistency 

relative to one another or relative to “top-down” global 

emission magnitudes owing to the different time periods 

included and incomplete geographic coverage addressed 

by these studies.

1.2.1.4 Methyl Chloroform (CH3CCl3)

Surface mixing ratios of methyl chloroform 

(CH3CCl3) continued to decrease at a near-constant exponential rate in the remote global atmosphere during 

2005–2008. By mid-2008 the global surface mean mixing ratio had decreased to about 11 ppt (Figure 1-1). This 

decline represents more than a factor of 10 decrease in 

global surface mixing ratios of this chemical since the 

early 1990s. As CH3CCl3 mixing ratios have declined in 

response to reduced emissions, hemispheric differences 

have also diminished. The mean annual hemispheric difference (NH minus SH) was a few percent of the global 

mean during 2005–2008, much smaller than during the 

1980s when emissions were substantial. Measured abundances and rates of change in 2008 agree to within about 

Figure 1-5. Atmospheric (“top-down”) 

global CCl4 emissions (Gg/yr) derived 

from observations (blue, red, and orange lines, some of which are shown 

in Figure 1-3) compared to “potential 

emissions” derived from UNEP production data (green lines). The lower “potential emissions” green line is derived 

from the difference between total CCl4

production (solid black line labeled “P”) 

reported to UNEP and the sum of feedstock and amounts destroyed (dashed 

line labeled “F&D”) (production magnitudes to feedstock alone are indicated 

with the dotted line labeled “F”). The 

upper “potential emission” green line 

was derived similarly as the lower 

line but was augmented by additional 

amounts to fill apparent gaps in UNEP 

reporting, plus an allotment for fugitive 

emissions of 2% of reported CCl4 feedstock use, plus an efficiency of only 75% for reported destruction. Topdown emission estimates are derived from a 1-box model of NOAA atmospheric data (red line labeled N1) 

and a 12-box analysis of the AGAGE data (orange line labeled A12) with a lifetime of 26 years (see Box 1-1). 

The influence of lifetimes between 23 and 33 years on emissions derived with the 1-box model from AGAGE 

data are also indicated (blue lines labeled A1, tau=23 and A1, tau=33). The TEAP “bottom-up” emission estimate for 

1996 is shown as a brown diamond (UNEP/TEAP, 1998).

0

50

100

150

200

250

300

1985 1990 1995 2000 2005 2010

Emission or production magnitude (Gg/yr)

P

F

F&D

A12

Potential 

Emission

N1

Year

A1, tau=33

A1, tau=23
1.20

Chapter 1

12% among the different ground-based measurement 

networks (Table 1-1).

Losses of CH3CCl3 are dominated by oxidation by 

the hydroxyl radical (OH). Other processes such as photolysis and oceanic removal also are significant sinks for 

CH3CCl3 (Clerbaux and Cunnold et al., 2007; Yvon-Lewis 

and Butler, 2002). Accurate quantification of all CH3CCl3

loss processes is particularly important because budget 

analyses of this chemical (e.g., Prinn et al., 2005) provide 

estimates of global abundance of the hydroxyl radical, an 

important oxidant for many reduced atmospheric gases.

The potential for significant terrestrial losses of 

CH3CCl3 has been further explored since the previous 

Assessment. Aerobic soils had been previously identified as a sink for CH3CCl3, accounting for 5 ± 4% (26 ± 

19 Gg/yr) of global removal rates in 1995 (Happell and 

Wallace, 1998). This estimate was based on soil gas profiles measured in Long Island, New York. A more recent 

study using flux chamber methods in southern California 

salt marshes and shrublands showed average net fluxes for 

CH3CCl3 that were <10% of uptake rates in the Long Island 

study, suggesting a less significant role for soils as a sink 

for this compound (Rhew et al., 2008). No new studies related to oceanic losses have been published since 2006; the 

estimated partial atmospheric lifetime for CH3CCl3 with 

respect to ocean loss remains 94 (81–145) years.

CH3CCl3 Emissions and Banks

Emissions of CH3CCl3 have declined substantially 

since the early 1990s. When derived from atmospheric 

changes and a 5-year lifetime, inferred emissions of 

CH3CCl3 during 2005–2008 are calculated to be less than 

10 Gg/yr (Figure 1-3). Banks for CH3CCl3 are thought to 

be negligible owing to its dominant past use in applications resulting in rapid release to the atmosphere.

CH3CCl3 emissions arise primarily from industrial 

production, though some fugitive emissions may arise 

from its use as a feedstock in the production of HCFC141b and HCFC-142b. Previous laboratory-based studies 

of biomass combustion combined with limited field sampling of biomass burning plumes suggested the potential 

for nonindustrial emissions of 2.5–11.5 Gg/yr CH3CCl3

from biomass burning (Rudolph et al., 1995; Rudolph et 

al., 2000). A recent study of many ambient air samples 

significantly affected by biomass burning suggests very 

small or negligible CH3CCl3 emissions from this source 

(<<1 Gg/yr) (Simpson et al., 2007).

1.2.1.5 Hydrochlorofluorocarbons (HCFCs)

HCFCs are regularly measured in three global 

ground-based networks and at a number of additional 

sites around the world using grab-sampling techniques 

(Montzka et al., 2009; O’Doherty et al., 2004; Stemmler 

et al., 2007; Yokouchi et al., 2006). Results from all three 

networks indicate that global mean surface mixing ratios 

of the three most abundant HCFCs (i.e., HCFC-22, -142b, 

and -141b) continued to increase during 2005–2008 (Figure 1-6). Mixing ratios have also been determined for 

HCFC-22 and HCFC-142b from Fourier transform infrared (FTIR) instruments onboard the Envisat (the Michelson Interferometer for Passive Atmospheric Sounding, or 

Figure 1-6. Global surface monthly-mean mixing 

ratios (parts per trillion) measured by NOAA (red) 

and AGAGE (blue) for the three most abundant 

HCFCs (left-hand scale) (Montzka et al., 2009; 

O’Doherty et al., 2004). Growth rates (ppt/yr) estimated as 12-month differences from AGAGE (black 

points) and NOAA (gray points) are shown relative 

to the right-hand scale and are plotted relative to 

the midpoint of the 12-month interval. Growth rates 

smoothed over 12-month periods appear as black 

and gray lines. Tic marks correspond to the beginning of each year.

60

80

100

120

140

160

180

200

2

4

6

8

10

12

14

16

18

0

5

10

15

20

0.0

0.4

0.8

1.2

1.6

2.0

0

5

10

15

20

0.0

0.5

1.0

1.5

2.0

2.5

HCFC-141b (ppt) HCFC-142b (ppt) HCFC-22 (ppt)

Rate of change (ppt/yr)

Year

Rate of change (ppt/yr)

Rate of change (ppt/yr)

1990 1995 2000 2005 2010
1.21

ODSs and Related Chemicals

MIPAS-E) and Atmospheric Chemistry Experiment (the 

ACE-Fourier Transform Spectrometer or ACE-FTS instrument) satellites, respectively (Moore and Remedios, 

2008; Rinsland et al., 2009).

The global mean surface mixing ratio of HCFC-22 

(CHClF2) was 188–192 ppt in 2008, with an averaged annual growth rate of 8.0 ± 0.5 ppt/yr (4.3 ± 0.3%/yr) during 

2007–2008 (Table 1-1; Figure 1-6). This increase is approximately 60% larger than the mean rate of change during 1992–2004 or the rate of change reported from global 

surface sampling networks during 2003–2004 (Clerbaux 

and Cunnold et al., 2007). Though the rate of HCFC-22 

increase from 2007–2008 was comparable to that projected in the A1 scenario of the previous Assessment report (7 

ppt/yr; Daniel and Velders et al., 2007), the mixing ratio 

increase during the entire 2005–2008 period was notably 

larger than in the scenario projection (Figure 1-1).

Moore and Remedios (2008) report a 2003 global 

mean HCFC-22 mixing ratio from MIPAS-E at 300 hPa 

of 177 ± 18 ppt (uncertainty includes 0.5 ppt of random 

error on the mean and an additional systematic uncertainty); this value is in fairly good agreement with the 2003 

global mean surface mixing ratio of 160 ± 2 ppt (Clerbaux 

and Cunnold et al., 2007). They also deduce an average 

HCFC-22 growth rate of 3.5 ± 0.4%/yr (5.4 ± 0.7 ppt/yr) in 

the northern midlatitude (20°N–50°N) lower stratosphere 

(50–300 hPa) between November 1994 and October 2003 

from the Atmospheric Trace Molecule Spectroscopy 

(ATMOS) (Atmospheric Laboratory for Applications and 

Science, ATLAS-3) based on measured HCFC-22/nitrous 

oxide (N2O) correlations. This rate is similar to the 3.92 ± 

2.08%/yr derived using a similar approach with ATMOS 

and ACE-FTS (from 2004) HCFC-22 data near 30°N 

(Rinsland et al., 2005). A slightly larger mean growth rate 

(4.3 ± 0.5%/yr or 6.0 ± 0.7 ppt/yr) is estimated for the lower 

stratosphere from the MIPAS-E HCFC-22 data at southern 

high latitudes (60°S–80°S) (Moore and Remedios, 2008). 

This averaged rate is comparable to global mean HCFC-22 

trends at the surface during this period (~5.2 ppt/yr).

Total vertical column abundances of HCFC-22 

above the Jungfraujoch station (Figure 1-2, an update of 

Zander et al., 2005) also indicate an increase of 4.31 ± 

0.17%/yr with respect to 2005 values over the 2005–2008 

period, which is comparable with NH trends from surface 

networks (4.2–4.5%/yr calculated similarly). Moreover, 

Gardiner et al. (2008) applied a bootstrap resampling method to aggregated total and partial column data sets from six 

European remote sensing sites to quantify long-term trends 

across the measurement network; they found a mean tropospheric increase for HCFC-22 at these sites of 3.18 ± 

0.24%/yr, which is slightly smaller than determined from 

ground-level grab samples at surface sites in high northern 

latitudes such as Mace Head, Barrow, or Alert during the 

analyzed period (1999–2003 rates of 3.7–3.9%/yr).

The global mean surface mixing ratio of HCFC142b (CH3CClF2) increased to 18.0–18.9 ppt in 2008 with 

an averaged annual growth rate of about 1.0–1.2 ppt/yr 

(6.1 ± 0.6%/yr) during 2007–2008 (Table 1-1; Figure 1-6). 

After declining from the late 1990s to 2003, the growth 

rate of HCFC-142b increased substantially during 2004–

2008. During 2007–2008 this rate was approximately two 

times faster than reported for 2003–2004 (Montzka et al., 

2009). This accelerated accumulation of HCFC-142b was 

not projected in the A1 scenario of the 2006 Assessment 

(the projected 2007–2008 rate was 0.2 ppt/yr); a substantial divergence occurred between projected and observed 

mixing ratios after 2004 (Figure 1-1). The mean difference in reported mixing ratios from AGAGE and NOAA 

of 3.3% (with AGAGE being higher) is primarily related 

to calibration differences of ~2.9% reported previously 

(O’Doherty et al., 2004). Global means from UCI are 

approximately 2% lower than NOAA (Table 1-1).

The first satellite measurements of HCFC-142b 

have been made from the ACE-FTS instrument (Rinsland 

et al., 2009). Monthly-mean ACE-FTS HCFC-142b mixing ratios over 13–16 kilometers (km) altitude, with an estimated total (random and systematic) error of ~20%, were 

used to derive trends at northern (25–35°N) and southern 

(25–35°S) midlatitudes of 4.94 ± 1.51%/yr and 6.63 ± 

1.23%/yr, respectively, over the interval from February 

2004 to August 2008. The ACE-FTS trends are consistent 

with those computed from flask sampling measurements 

over a similar time period (5.73 ± 0.14%/yr at Niwot 

Ridge (40°N) and 5.46 ± 0.08%/yr at Cape Grim (40°S) 

over the interval from July 2003 to July 2008) (Rinsland 

et al., 2009).

The global mean surface mixing ratio of HCFC141b (CH3CCl2F) continued to increase during 2005–

2008. By 2008, mean, global surface mixing ratios were 

18.8–19.5 ppt (Table 1-1). The growth rate of HCFC141b decreased from approximately 2 ppt/yr in the mid1990s to <0.5 ppt/yr in 2004–2005 (Figure 1-6). Since 

2005 the growth rate has varied between 0.2–0.8 ppt/yr, 

similar to the mean 0.5 ppt/yr increase projected in the A1 

scenario over this period (Daniel and Velders et al., 2007). 

The mean increase during 2007–2008 was 0.6 (±0.1) ppt/

yr (or 3.2 ± 0.5%/yr).

The annual global surface mean mixing ratio of 

HCFC-124 (CHClFCF3) has been updated from AGAGE 

measurements (Prinn et al., 2000) and has decreased to 1.5 ± 

0.1 ppt in 2008, with an averaged annual growth rate of 

−0.01 ± 0.01 ppt/yr (−0.8 ± 0.8%/yr) for 2007–2008. No updated HCFC-123 (CHCl2CF3) measurements are available.

Recent changes in atmospheric growth rates of 

the three most abundant HCFCs can be explained qualitatively with UNEP (2010) production and consumption 

data (Figure 1-7). Global HCFC production for dispersive 

uses increased rapidly in developed countries during the 
1.22

Chapter 1

1990s. But as this production was being phased out in 

developed countries, global totals decreased slightly from 

2000–2003. This trend reversed during 2003–2008 as production and consumption grew substantially in developing 

countries (those operating under Article 5 of the Montreal 

Protocol, also referred to as A5 countries). In 2008 HCFC 

data reported to UNEP, developing (A5) countries accounted for 74% and 73% of total, ODP-weighted HCFC 

consumption and production, respectively (UNEP, 2010).

HCFC Emissions and Banks

Global emissions of HCFC-22 continued to increase 

during 2005–2008. By 2008, “top-down” emissions inferred from global atmospheric changes totaled 320–380 

Gg/yr, up from approximately 280 Gg/yr in 2004. These 

emissions are reasonably consistent with the emissions 

derived from banks for 2002 and projected for 2008 in a 

2005 study (IPCC/TEAP, 2005) (Figure 1-3). These results suggest that the dominant emission for HCFC-22 is 

from banks contained in current, in-use applications. Total “bottom-up” emissions derived for past years (UNEP/

TEAP, 2006; estimates available through 2003 only) show 

a similar trend to emissions derived from atmospheric 

data, but are larger in most years (Figure 1-3).

While atmosphere-derived global emissions for 

HCFC-141b and HCFC-142b decreased slightly during 

the 2000–2004 period as production in developed countries was diminishing (Montzka et al., 2009), emissions of 

both of these HCFCs increased during 2005–2008 (Figure 

1-3). The substantial increase in HCFC-142b emissions 

was not projected in the A1 scenario of the previous Assessment (Daniel and Velders, 2007). In that scenario, 

a 23% emissions decline was projected during 2004 to 

2008 (relative to 2004 levels). Instead, HCFC-142b emissions in 2008 derived from observed global mixing ratio 

changes (37 ± 7 Gg/yr) were approximately two times 

larger than had been projected for that year. Changes in 

HCFC-141b emissions during 2005–2008 were quite consistent with those projected in the A1 scenario (Daniel and 

Velders, 2007).

“Bottom-up” estimates of HCFC-141b emissions 

(UNEP/TEAP, 2006) have captured the overall increase 

in emissions of this compound derived from “top-down” 

calculations, but with a slightly different time lag. Similar 

“bottom-up” estimates of HCFC-142b have also captured 

the rough changes implied from atmospheric data, but in 

this case, the “bottom-up” estimates are substantially lower than implied from year-to-year atmospheric observations during 2000–2004 (Figure 1-3). For both HCFCs, it 

is apparent that emissions from banks estimated for 2002 

and projected for 2008 (IPCC/TEAP, 2005) account for 

<50% of total emissions for these compounds. About 10–

20% of annual production of HCFC-141b is for solvent 

uses that result in release to the atmosphere shortly after 

production (UNEP/TEAP, 2006). Based on production 

data this would yield emissions of 10–20 Gg/yr and explain some of the difference between atmosphere-derived 

emissions and bank-related emissions (UNEP/TEAP, 

2006). Rapid losses of HCFC-142b during or shortly after production would also provide an explanation for only 

some of the shortfall in emissions not explained by bank 

releases. HCFC bank magnitudes have increased in recent 

years given that reported production in recent years has 

substantially exceeded emissions (Montzka et al., 2009).

Summed, “top-down” emissions from HCFCs have 

increased during 2005–2008. When weighted by semiempirical ODPs (Chapter 5), the sum of emissions from 

HCFCs totaled 22 ± 2 ODP-Kt in 2008. The sum of emissions of HCFCs weighted by direct, 100-yr GWPs has 

increased on average by 5 ± 2%/yr from 2004 to 2008, 

and by 2008 amounted to 0.74 ± 0.05 Gt CO2-eq/yr.

The 2007 Adjustments to the Montreal Protocol are 

expected to have a discernable influence on HCFC emissions in the coming decade (see Chapter 5). But because 

those Adjustments are scheduled to affect HCFC production and consumption only after 2009, it is too soon to 

discern any influence of these Protocol adjustments in 

the mixing ratios or emissions derived for 2008 or earlier 

years reported in this Assessment.

Regional emissions for HCFCs from atmospheric 

measurements (Box 1-1) have been derived in different 

studies since the previous Assessment report. Compari0

5

10

15

20

25

30

35

40

45

1990 1995 2000 2005

HCFC production 

and consumption (ODP-Gg)

Consumption

Production

Non A5

A5

Total

Year

Figure 1-7. Production and consumption (in ODPGg) of HCFCs in developed (Non A5) and developing (A5) countries as reported to UNEP for dispersive uses only. Also shown (thin line) is the global 

total reported production (UNEP, 2010). Note that 

1 ODP-Gg = 1 ODP-Kt.
1.23

ODSs and Related Chemicals

sons with “bottom-up” estimates provide useful information on the accuracy of individual country accounting of 

their emissions of ODSs and regional estimates derived 

from atmospheric measurements. In an analysis of United 

States (U.S.) HCFC emissions, HCFC-22 mean emissions 

during 2004–2006 were estimated from aircraft measurements to be 46 (21–69) Gg/yr, or substantially lower than 

amounts derived from “bottom-up” estimates (Millet et al., 

2009). U.S. HCFC-22 emissions estimated with “bottomup” inventory methods by the U.S. Environmental Protection Agency (EPA) were estimated at between 89 and 97 

Gg/yr during 2004–2008 and have been used to derive a 

U.S. contribution to global HCFC-22 atmospheric mixing 

ratios of between 21 and 45% during 2006 (Montzka et 

al., 2008).

Detailed analyses of HCFC emissions have also 

been reported for China in recent years, concurrent with 

a substantial increase in reported HCFC production and 

consumption in this country. Inventory-based, “bottomup” estimates suggest HCFC-22 emissions increasing 

from 34 to 69 Gg/yr during 2004–2007 (Wan et al., 2009), 

or 12–20% of total global emissions during these years. 

Atmosphere-derived emissions attributed to China based 

on correlations to carbon monoxide (CO) and inversion 

modeling of elevated mixing ratios at down-wind sampling 

locations suggest slightly larger emissions than the inventory approach (52 ± 34 Gg/yr as the average of 2004–2005 

from Yokouchi et al., 2006; 60 and 71 Gg/yr for 2005 and 

2006 from Stohl et al. (2009); and 165 (140–213) Gg/yr in 

2007 from Vollmer et al. (2009), though the Vollmer et al. 

(2009) estimate may be biased high because of relatively 

higher per-capita HCFC-22 emissions near Beijing than in 

other regions of China (Stohl et al., 2009)).

1.2.1.6 Methyl Bromide (CH3Br)

The global, annual mean surface mixing ratio 

of methyl bromide (CH3Br) had reached 7.3 to 7.5 ppt 

in 2008 (Figure 1-8; Table 1-1), down from the 9.2 ppt 

measured during the three years (1996–1998) before industrial production declined as a result of the Montreal 

Protocol (Yvon-Lewis et al., 2009). Global mixing ratios declined during 2005–2008 at a rate of −0.14 ppt/yr, 

which is slightly slower than the mean decline observed 

since 1999 when industrial production was first reduced. 

Since 1999, the annual mean hemispheric difference (NH 

minus SH) has decreased by nearly 50%: this measured 

difference was 1.2 ppt in 2008 compared to 2.3 ± 0.1 ppt 

during 1996–1998 (Figure 1-8).

Declines in the global tropospheric abundance and 

hemispheric difference of CH3Br have coincided with 

decreases in global industrial production and subsequent 

emission. Reported global methyl bromide consumption 

in 2008 for all uses including uncontrolled quarantine and 

pre-shipment (QPS) uses was 73% below peak amounts 

reported in the late 1990s. An emission history derived 

from these reported data suggests a reduction in total 

fumigation-related CH3Br emissions of 71% by the end of 

2008 (see Figure 1-8).

The concurrent decline in the measured hemispheric difference and industrially derived emissions suggests 

a mean hemispheric mixing ratio difference close to 0 ppt 

6

7

8

9

10

11

12

mean

0

10

20

30

40

50

1994 1996 1998 2000 2002 2004 2006 2008 2010

0.0

0.5

1.0

1.5

2.0

2.5

3.0

Mixing ratio (ppt) NH – SH difference (ppt) Emissions from fumigation (Gg/yr)

SH

NH

Year

Figure 1-8. Top panel: Monthly hemispheric means 

for CH3Br mixing ratios (ppt) (Montzka et al., 2003 updated). Middle panel: NH - SH difference by month 

(points) and smoothed over 12-month periods (bold 

red line). Bottom panel: Fumigation-related emissions (Gg/yr) of CH3Br derived from reported regulated consumption × 0.65 plus reported quarantine and 

pre-shipment consumption × 0.84 (UNEP 2007b), 

where 0.65 and 0.84 are the estimated mean fractions of reported production to these different uses 

that ultimately become emitted to the atmosphere 

(UNEP 2007b).
1.24

Chapter 1

Table 1-2. Summary of the estimated source and sink strengths (Gg/yr) of methyl bromide (CH3Br) for 

periods 1996–1998 and 2008.

1996–1998 Range 2008 Range Reference Note

SOURCES

Fumigation- dispersive (soils) 41.5 (28.1 to 55.6) 6.7 (4.6 to 9.0) 1, 2 a

Fumigation- quarantine/

 pre-shipment

7.9 (7.4 to 8.5) 7.6 (7.1 to 8.1) 1, 2 b

Ocean 42 (34 to 49) 42 (34 to 49) 3, 4 c

Biomass Burning 29 (10 to 40) 29 (10 to 40) 5, 6 d

Leaded gasoline 5.7 (4.0 to 7.4) < 5.7 7 e

Temperate peatlands* 0.6 (−0.1 to 1.3) 0.6 (−0.1 to 1.3) 8, 9, 10 f

Rice paddies* 0.7 (0.1 to 1.7) 0.7 (0.1 to 1.7) 11, 12 g

Coastal salt marshes* 7 (0.6 to 14) 7 (0.6 to 14) h

based on California saltmarshes 14 (7 to 29) 14 (7 to 29) 13, 14 i

based on Scottish saltmarsh 1 (0.5 to 3.0) 1 (0.5 to 3.0) 15 j

based on Tasmania saltmarsh 0.6 (0.2 to 1.0) 0.6 (0.2 to 1.0) 16 k

Mangroves 1.3 (1.2 to 1.3) 1.3 (1.2 to 1.3) 17 l

Shrublands* 0.2 (0 to 1) 0.2 (0 to 1) 18 m

Rapeseed 4.9 (3.8 to 5.8) 5.1 (4.0 to 6.1) 19 n

Fungus (litter decay) 1.7 (0.5 to 5.2) 1.7 (0.5 to 5.2) 20 o

Fungus (leaf-cutter ants) 0.5 0.5 21 p

Potential terrestrial sources q

Tropical trees n.q. n.q. 22, 23 r

Temperate woodlands n.q. n.q. 24, 25 s

Tropical ferns n.q. n.q. 26

Abiotic decomposition n.q. n.q. 27 t

Subtotal (Sources) 143 111.5

SINKS

Ocean 56 (49 to 64) 49 (45 to 52) 3 u

OH and photolysis 77 63.6 3 v

Soils 40 (23 to 56) 32 (19 to 44) 28–33 w

Subtotal (Sinks) 177 147.6

Total (SOURCES−SINKS) −34 −36.1

* All asterisked items were estimated from measurements of net fluxes and may be influenced by sinks within them, thus they represent minimum gross 

fluxes. n.q. = not quantified. 

Notes:

a. Soil fumigation emission rates estimated as 65% (46–91%) of reported consumption rates (ref 2).

b. QPS emission rates estimated as 84% (78–90%) of reported consumption rates (ref 2).

c. Oceanic production rate calculated based on saturation state of pre-phase-out ocean: production = net oceanic flux − ocean sink, where net flux is −14 

(−7 to −22) Gg/yr.

d. Biomass burning estimates unchanged from the previous Assessment (Clerbaux and Cunnold et al., 2007), which is slightly higher than the 18–23 Gg/

yr estimate in ref (3).

e. 2006–2008 values are not separately quantified but expected to be lower with the phase-out of leaded gasoline use.

f. Temperate peatlands net flux calculated by updating (ref 7) with 3-year average fluxes at same New Hampshire sites (ref 8). Range includes median 

estimate of 0.9 Gg/yr from Irish peatland study (ref 10).
1.25

ODSs and Related Chemicals

for CH3Br in preindustrial times. Accordingly, the pre1990 global mixing ratio trend used in creating scenario 

A1 for CH3Br in Chapter 5 was derived from Southern 

Hemisphere firn data by including a time-varying hemispheric ratio that increased linearly from 1.0 in 1940 to 

1.3 in 1995.

In the past, much research related to methyl bromide focused on refining our understanding of source and 

sink magnitudes in order to understand the relative contribution of anthropogenic methyl bromide emissions to 

methyl bromide atmospheric abundance and, therefore, 

provide more accurate projections of the atmospheric 

response to reduced industrial production. For example, 

the global measured decline since 1996–1998 of ~2 ppt 

in response to a decline in fumigation emissions of 60–

70% suggests a total contribution from fumigation-related 

production before the phase-out (i.e., during 1996–1998) 

of 2.8–3.2 ppt (provided other sources or loss frequencies did not change appreciably). Considering that peak 

global mixing ratios were 9.2 ppt during 1996–1998, this 

suggests that industrially derived emissions accounted for 

31–36% of total CH3Br emissions in the three years before 

the phase-out began.

A recent modeling analysis of global atmospheric 

CH3Br observations provided additional constraints to our 

understanding (Yvon-Lewis et al., 2009). It suggested that 

the observed global declines are well explained given our 

understanding of CH3Br sources and sinks and the known 

changes in anthropogenic emissions, though a substantial 

source (~35 Gg/yr) is still unaccounted for in current budget compilations (Yvon-Lewis et al., 2009) (Table 1-2). 

The best-estimate budget derived in this work (based on 

observed global and hemispheric mixing ratio trends and 

seasonal variations together with time-varying sources 

and sinks) suggested a pre-phase-out anthropogenic fumigation contribution of ~28%. Though uncertainties in the 

variability of natural emissions and in the magnitude of 

methyl bromide stockpiles in recent years add uncertainty 

to our understanding of this ratio, when these new model 

results are considered together with the more simple analysis of methyl bromide mixing ratio changes since 1999 

(see previous paragraph), a pre-phase-out anthropogenic 

contribution of 25–35% is estimated.

Our understanding of preindustrial mixing ratios of 

methyl bromide in the Southern Hemisphere has improved 

since the 2006 Assessment. A 2000-year record derived 

for methyl bromide from an ice core collected at South 

Pole (Saltzman et al., 2008) shows no systematic trend and 

a mean mixing ratio in samples with mean ages from 160 

Before the Common Era (BCE) to 1860 CE of 5.39 ± 0.06 

g. Re-evaluation of global emission rates that were previously estimated as 3.5 Gg/yr in ref (11).

h. Salt marsh net flux estimated as the mid-range of the best estimates provided from four different studies.

i. Estimates were 14 (7–29) Gg/yr (from ref 13) and 8 to 24 Gg/yr (ref 14) depending on whether extrapolation incorporated or excluded mudflats, respectively.

j. Low and high range based on lowest and highest emitting of eight sites in ref (15).

k. Extrapolations calculated using mean fluxes reported in ref (16).

l. Based on study of two mangrove species, with range as results of two different methods of extrapolation.

m. Shrublands range as reported in Montzka and Fraser et al. (2003).

n. Rapeseed flux as reported in ref (19) with year-by-year data from author. 2007–2008 data uses results for 2003, the last year estimated.

o. These emission rates may possibly incorporate emission rates reported for woodlands and forest soils.

p. No range provided.

q. Terrestrial sources are poorly quantified and based on very limited studies so are not included in the tabulated sources.

r. Global extrapolations based on measurements of Malaysian trees (18 Gg/yr in ref 22) exceed the upper limit on net flux estimated in a study of a tropical South American rainforest (17 Gg/yr in ref 23), so no discrete estimate is included here.

s. Range of temperate woodland fluxes of 0.4 to 3.1 Gg/yr can be estimated by extrapolating net fluxes in ref (24) and gross production rates in ref (25) 

to global area of 13 × 1012 m2. Eucalyptus forest floor shows no net emissions (ref 16).

t. The largest abiotic production rates, observed from the decomposition of saltwort leaves (ref 27), are roughly 2% of the emission rates from the live 

saltwort plant (ref 14).

u. Oceanic consumption calculated assuming production rates constant, with a decrease in saturation anomaly because of decreasing atmospheric concentrations, with global average net flux in 2007 predicted to be −6.6 (−3.3 to −10.4) Gg/yr. Range assumes the same net flux error range as 1996–1998 

(as percent of flux).

v. A range has not been provided here.

w. Soil sink (refs. 28, 29) scaled to updated background Northern Hemisphere concentrations of 10.3 ppt (1996–1998) and 8.2 ppt (2006–2008), and 

includes a new tundra sink (refs. 30, 31) and a revised average temperate grasslands flux (refs. 28, 32, 33).

References: 

1. UNEP, 2010. 2. UNEP, 2007b. 3. Yvon-Lewis et al., 2009. 4. King et al., 2002. 5. Clerbaux and Cunnold et al., 2007. 6. Andreae and Merlet, 2001. 

7. Thomas et al., 1997. 8. Varner et al., 1999a. 9. White et al., 2005. 10. Dimmer et al., 2001. 11. Redeker and Cicerone, 2004. 12. Lee-Taylor and 

Redeker, 2005. 13. Rhew et al., 2000. 14. Manley et al., 2006. 15. Drewer et al., 2006. 16. Cox et al., 2004. 17. Manley et al., 2007. 18. Rhew et al., 

2001. 19. Mead et al., 2008b. 20. Lee-Taylor and Holland, 2000. 21. Mead et al., 2008a. 22. Blei et al., 2010. 23. Gebhardt et al., 2008. 24. Drewer

et al., 2008. 25. Rhew et al., 2010. 26. Saito and Yokouchi, 2006. 27. Wishkerman et al., 2008. 28. Shorter et al., 1995. 29. Varner et al., 1999b. 30. 

Rhew et al., 2007. 31. Hardacre et al., 2009. 32. Rhew and Abel, 2007. 33. Teh et al., 2008.

Table 1-2, continued (notes).
1.26

Chapter 1

ppt (uncertainty represents 1 standard error here). This 

preindustrial mixing ratio is similar to the 5.1–5.5 ppt observed in the deepest firn-air samples at South Pole and 

Law Dome, Antarctica (Butler et al., 1999; Trudinger et 

al., 2004) and previous ice-core results at a different site 

in Antarctica in samples dated 1671–1942 CE (5.8 ppt; 

Saltzman et al., 2004). Based on these preindustrial SH 

mixing ratio results, the observed decline in SH mixing 

ratios from their peak (8.0 ppt) through the end of 2008 

(6.7 ppt) suggests that SH mixing ratios have declined 50–

60% of the way back to preindustrial levels as industrial 

production declined by a similar magnitude (60–70%).

The concurrence between global atmospheric 

changes relative to expected emissions declines, the decreased NH–SH differences, and the decline in the mean 

SH mixing ratio much of the way back to its preindustrial 

value, all suggest that production restrictions on CH3Br 

have been successful at substantially reducing the global 

atmospheric abundance of CH3Br. It is worth noting that 

this benefit was achieved despite substantial emissions 

from natural sources and an incomplete understanding of 

the global budget of CH3Br.

Budget

Significant uncertainties remain in the detailed atmospheric budget of methyl bromide despite additional 

research since the previous Assessment. As indicated 

above, known sinks still outweigh best estimates of known 

sources by about 35 Gg/yr or roughly 20% of the total annual flux. This discrepancy remains even though the tropospheric burden, known sources, and known sinks have 

quantifiably changed in the last decade.

In light of the changing atmospheric concentrations 

of CH3Br, separate budgets are created for pre-phase-out 

(1996–1998) and the 2008 atmospheres (Table 1-2). Reported consumption of CH3Br from fumigation (dispersive 

and quarantine/pre-shipment uses) declined 73% between 

these periods, from 70.5 Gg/yr to 18.9 Gg/yr (UNEP, 

2010). Before phase-out, pre-plant soil fumigation was 

the major use for CH3Br, but this use had declined by 

84% by 2008. Over the same period, CH3Br consumption for quarantine and pre-shipment (QPS) applications 

has ranged between 7.5 and 12.5 Gg/yr, as this particular 

application is an exempted use (UNEP, 2010). As a result, consumption for QPS use accounted for nearly 50% 

of total global CH3Br uses during 2007 and 2008 (UNEP, 

2010). These values do not include production for use as 

a chemical feedstock, which is assumed to be completely 

consumed and which averaged 3.9 Gg/yr from 1995–1997 

and 6.6 Gg/yr from 2003–2005 (the last three years reported in UNEP, 2007b).

Other anthropogenically influenced sources of 

CH3Br include leaded gasoline combustion, biomass 

burning, and growth of certain methyl bromide-emitting 

crops. Our understanding of the amount of methyl bromide emitted from biomass burning remains unchanged 

from the previous Assessment report as 29 (10 to 40) 

Gg/yr. Biomass burning emissions, however, vary substantially from year to year and peaked in 1998 during 

the strong El Niño event and enhanced burning then (van 

der Werf et al., 2004; Yvon-Lewis et al., 2009). Emissions from biofuel burning in the developing world are 

included in the above estimate, although they have been 

derived separately to be 6.1 ± 3.1 Gg/yr (Yvon-Lewis et 

al., 2009). Crop production levels of rapeseed (canola) 

and rice, two known crop sources of methyl bromide, 

have been increasing steadily due to demand for food 

supply and biofuel (FAO, 2009). A recent extrapolation 

of rapeseed CH3Br emissions using crop harvest, growth 

rate, and global production data suggest a three- to fourfold increase from 1980 to 2003 (Mead et al., 2008a). In 

this new study, estimated average emission rates ranged 

from 4.3 to 6.2 Gg/yr between 1996 and 2003 (the final 

year estimated). These values are slightly less than the 

previous estimate of 6.6 Gg/yr (Clerbaux and Cunnold 

et al., 2007; Gan et al., 1998). While cabbage and mustard production also are increasing, total emission from 

these crops is estimated at <0.1 Gg/yr. A re-evaluation 

of CH3Br emission from rice crops using a model incorporating temperature, seasonality, and soil moisture 

effects yields a lower source estimate (0.5 to 0.9 Gg/

yr) than derived previously (3.5 Gg/yr) (Lee-Taylor and 

Redeker, 2005; Redeker and Cicerone, 2004).

Emission rates from the three known major natural sources (oceans, freshwater wetlands, and coastal salt 

marshes) have been revised downward since the 2002 Assessment (compare Table 1-2 to Table 1-9 in Montzka and 

Fraser et al., 2003). Ocean production rates have been 

revised from 63 (23–119) Gg/yr down to 42 (34–49) Gg/

yr (Yvon-Lewis et al., 2009). Freshwater wetlands were 

previously estimated as a 4.6 (2.3 to 9.2) Gg/yr net source 

based on a partial season of measurements from two New 

Hampshire peatlands (Varner et al., 1999a). This source, 

specified in Table 1-2 as temperate peatlands, has been 

revised downward to 0.6 (−0.1 to 1.3) Gg/yr based on 

a 3-year study at the same New Hampshire sites, which 

showed much lower average net fluxes (White et al., 

2005). Coastal salt marshes were previously estimated as 

a 14 (7–29) Gg/yr source based on a study of two southern 

California salt marshes (33°N) (Rhew et al., 2000). While 

a separate, nearby study (34°N) found similar net fluxes, 

the global extrapolations varied from 8 Gg/yr (assuming 

salt marsh areas included low-producing mudflats) to 24 

Gg/yr (assuming surface areas were entirely vegetated) 

(Manley et al., 2006). However, much smaller CH3Br net 

emission rates were observed from coastal salt marshes in 

Scotland (56°N) (Drewer et al., 2006) and Tasmania, Aus-
1.27

ODSs and Related Chemicals

tralia (41°S) (Cox et al., 2004), which suggested global 

emission rates of 1 (0.5–3.0) Gg/yr and 0.6 (0.2–1.0) Gg/

yr, respectively. Because emissions are strongly related 

to plant species and climatic conditions, the quantification 

of this source requires a more detailed understanding of 

salt marsh distributions and vegetation types. The updated 

value in Table 1-2 (7 Gg/yr) represents the mid-range of 

globally extrapolated fluxes, with the full range representing the various study mean values.

Since the previous Assessment report, several additional natural CH3Br sources have been identified (see Table 1-2), although the addition of these sources does not yet 

resolve the budget imbalance between sources and sinks. 

Mangroves (Manley et al., 2007) and fungus cultivated by 

leaf-cutter ants (Mead et al., 2008b) are newly identified 

sources, although they are estimated to be relatively minor 

sources globally. Measurements from tropical trees (Blei 

et al., 2010) and ferns (Saito and Yokouchi, 2006) in SE 

Asia suggest that these may be large sources, up to 18 Gg/

yr if results from these studies are globally representative. 

Aircraft measurements over a South American rainforest 

revealed no significant net emissions, however, and suggest an upper limit of 17 Gg/yr for the global tropical forest 

flux (Gebhardt et al., 2008). Because of this disparity, this 

source is not included in Table 1-2. In addition to the previously identified mechanism to produce methyl halides 

abiotically from the degradation of organic matter (Keppler et al., 2000), the abiotic production of methyl bromide 

in plant material has also been shown (Wishkerman et al., 

2008); it is not clear how important these mechanisms are 

in relation to biotic production rates.

Natural terrestrial ecosystems can be both sources 

(e.g., from fungi, litter decomposition, certain plant species) and sinks (biological degradation in soils) for CH3Br. 

Both emissions and uptake have been observed in temperate 

grasslands (Cox et al., 2004; Rhew and Abel, 2007; Teh et 

al., 2008), temperate forest (Dimmer et al., 2001; Varner 

et al., 2003), temperate shrubland (Rhew et al., 2001), and 

Arctic tundra (Rhew et al., 2007; Teh et al., 2009). Temperate woodland soils in Scotland are a net source for CH3Br 

(Drewer et al., 2008) while oak-savanna woodland soils in 

California are a net sink (Rhew et al., 2010). To better understand the highly variable net fluxes found in many of 

these ecosystems, measurements of gross production and 

consumption rates have been derived by clearing vegetation 

from some sites (White et al., 2005; Drewer et al., 2006), 

using a stable isotope tracer method (Rhew and Abel, 2007; 

Teh et al., 2008; Teh et al., 2009; Rhew et al., 2010), or 

modeling soil uptake separately (Varner et al., 2003).

Known sinks of CH3Br (oceans, OH, photolysis, 

and soil microbial uptake) have uptake rates that scale to 

tropospheric concentrations and, therefore, have declined 

in the current decade, as indicated in Table 1-2 (YvonLewis et al., 2009). The partial lifetime of atmospheric 

CH3Br with respect to ocean loss has been updated to 

2.2–2.4 years based on new model results (Yvon-Lewis 

et al., 2009). The partial lifetime with respect to loss by 

oxidation by OH and photolysis remains unchanged at 

1.7 years. The partial lifetime with respect to soil loss 

has increased from 3.1 years to 3.3 to 3.4 years, as described below. The best estimate lifetime for atmospheric CH3Br has therefore increased slightly from 0.7 years 

to 0.8 years.

In the last two Assessment reports (Clerbaux and 

Cunnold et al., 2007; Montzka and Fraser et al., 2003), 

the soil sink was estimated to be 47 Gg/yr based on 

earlier studies that assumed Northern Hemisphere tropospheric concentrations of 11 ppt. The soil sink was 

recalculated for 1996–1998 and 2008 using background 

Northern Hemisphere concentrations of 10.3 ppt and 8.2 

ppt, respectively, to yield uptake rates of 44 ± 15 and 

35 ± 12 Gg/yr, respectively. The addition of the tundra 

sink (0.31 ± 0.06 Gg/yr) does not significantly change 

these results (Rhew et al., 2007; Hardacre et al., 2009). 

However, recent field studies in temperate grasslands in 

California (Rhew and Abel, 2007) and Colorado (Teh 

et al., 2008) show gross uptake rates that are one-fourth 

and one-eighth of the previous temperate grassland uptake rates. The average growing season uptake rate incorporating these new studies is half of the previously 

reported flux for temperate grasslands, which implies a 

15% reduction of the soil sink to 40 ± 16 Gg/yr for 1996–

1998 and 32 ± 13 Gg/yr for 2008. While still within the 

range of errors, the best estimate for partial atmospheric 

lifetime for CH3Br with respect to the soil sink would 

be increased from 3.0–3.1 years to 3.3–3.4 years. New 

results from a temperate woodland also suggest lower 

soil uptake rates (Rhew et al., 2010), but these were not 

included in the revised soil uptake rate here.

1.2.1.7 Methyl Chloride (CH3Cl)

Methyl chloride is the most abundant chlorinecontaining organic compound in the atmosphere and contributes 16% to the total chlorine from long-lived gases in 

the troposphere (see, for example, Table 1-12). Although 

it is not a controlled substance, CH3Cl has several natural sources and sinks in common with CH3Br, which is a 

controlled substance. Thus, atmospheric changes in CH3Cl 

and an updated knowledge of its global budgets can provide 

a context for understanding a large amount of atmospheric 

Cl that is not controlled by the Montreal Protocol, as well 

as insights into changes in the natural CH3Br budget.

Global mixing ratios of CH3Cl increased by small 

amounts during 2004–2008 (2.3 to 2.7 ± 1.2 ppt/yr, or 

0.4–0.5 ± 0.2%/yr), though these changes follow larger decreases reported for 1998–2001 of −1.5%/yr (Simmonds et 

al., 2004). Changes observed from 2007 to 2008 are small, 
1.28

Chapter 1

though both increases and decreases are reported by different measurements networks. The cause of these differences is not well documented, but the differences may be 

insignificant relative to measurement errors, or may reflect 

regional variations in rates of change for this compound.

Incorporating CH3Cl observations from global and 

regional networks (AGAGE, NOAA, System for Observation of halogenated Greenhouse gases in Europe (SOGE), 

National Institute for Environmental Studies (NIES)) into 

a 3-D global chemical transport model, Xiao et al. (2010a) 

estimated global emissions for 2000–2004 to be 4.10 ± 

0.47 Tg/yr, which is comparable to emissions reported in 

the previous Assessment (4.1 to 4.4 Tg/yr, Clerbaux and 

Cunnold et al., 2007). Model results indicate that about 

55%, or 2.3 ± 0.3 Tg/yr, of CH3Cl comes from tropical terrestrial sources that vary with global temperature changes 

(Xiao et al., 2010a). This is also comparable to prior model study estimates of 2.4 to 2.9 Tg/yr from tropical ecosystems, as reported in the previous Assessment.

Since the previous Assessment, several studies on 

South American and Asian rainforests have reinforced the 

importance of the tropical terrestrial CH3Cl source. Air 

samples obtained by aircraft over South American tropical 

rainforests (Suriname, French Guyana) in 2005 indicated 

net emissions of CH3Cl (Gebhardt et al., 2008). A survey of 197 plants from the subtropical island of Iriomote 

(Japan) found that 18% produced significant amounts of 

CH3Cl (Yokouchi et al., 2007), while a survey of 119 tree 

species from tropical Malaysia found that 21% produced 

significant amounts of CH3Cl (Saito et al., 2008). A separate study in Malaysian Borneo showed large emissions 

from live plants, though much smaller emissions were 

found from leaf litter (Blei et al., 2010). Global tropical 

source estimates based on extrapolations of the above studies ranged from 0.9 to 2.5 Tg/yr (Blei et al., 2010; Gebhardt et al., 2008; Saito et al., 2008; Yokouchi et al., 2007). 

These values are on the lower side of estimates of tropical 

emission magnitudes provided in the previous Assessment. 

Additional minor tropical sources of CH3Cl have been 

identified, including mangroves (~0.011 Tg/yr) (Manley et 

al., 2007; Yokouchi et al., 2007) and fungus cultivated by 

leaf cutter ants (<0.001 Tg/yr) (Mead et al., 2008b).

Emissions of CH3Cl can originate from biogenic 

production by vascular plants (Yokouchi et al., 2002), abiotic release from dead or senescent plant material (Hamilton et al., 2003), and emissions from tropical wood-rot 

fungi (Moore et al., 2005). Mass balance approaches have 

been used to estimate emission magnitudes from sources that have distinct stable carbon isotope ratios (δ13C). 

Methyl chloride produced from the abiotic methylation of 

chloride has an exceptionally depleted δ13C value (Keppler et al., 2004), and a mass balance approach suggests 

that this mechanism is a dominant source of atmospheric 

CH3Cl in terrestrial tropical and subtropical ecosystems at 

1.8 to 2.5 Tg/yr (Keppler et al., 2005). Saito and Yokouchi (2008) also use an isotopic mass balance approach, incorporating more recent isotopic measurements of CH3Cl 

from tropical plants, and estimate the total tropical source 

to be 2.9 to 4.2 Tg/yr, with 1.5 to 3.0 Tg/yr coming from 

live vegetation and the remainder from abiotic production 

from senescent leaves.

Biomass burning, oceans, and anthropogenic activities are major sources of atmospheric CH3Cl. A mechanism to produce CH3Cl through the photochemical reaction of dissolved organic matter in saline waters has been 

reported (Moore, 2008). Supersaturations of CH3Cl have 

been observed in coastal waters off of the United States 

(Hu et al., 2010) and China (Lu et al., 2010), and global 

coastal emissions may increase the global oceanic emission estimate by 0.02 to 0.10 Tg/yr (Hu et al., 2010). The 

inversion analysis of Xiao et al. (2010a) suggests emissions of 0.43 ± 0.1 Tg/yr from the global ocean, consistent with the estimate of 0.38–0.51 Tg/yr in the previous 

Assessment. Though there are no updates to the biomass 

burning and anthropogenic source terms, the Xiao et al. 

(2010a) inversion analysis suggests global biomass burning emissions of 0.92 ± 0.2 Tg/yr.

Coastal salt marshes were previously estimated 

to be a large CH3Cl source, up to 0.17 (0.07 to 0.44) Tg/

yr, based on measurements in southern California (Rhew 

et al., 2000). Recent studies, however, show large geographic variability in this source and high sensitivity to 

methods of extrapolation. Measurements at a different 

southern California salt marsh scale up to 0.05 to 0.16 Tg/

yr, depending on whether or not the areal averaging incorporates low-producing mudflats along with vegetated 

areas (Manley et al., 2006). Much smaller emission rates 

observed at a salt marsh in Tasmania scale to <0.01 Tg/yr 

(Cox et al., 2004). A salt marsh in China appeared to be 

a net sink during the growing season, though these results 

were derived under unusually high ambient concentrations 

(1–60 parts per billion (ppb)) (Wang et al., 2006).

The major sinks of tropospheric CH3Cl include 

oxidation by hydroxyl radicals, loss to the stratosphere, 

reaction with chlorine radicals, loss to polar ocean waters, and uptake by soils. These sink estimates remain 

largely unmodified since the previous Assessment so the 

CH3Cl lifetime remains unchanged (Table 1-3), although 

a study using stable isotope ratios suggests that the soil 

sink may be much larger than previously estimated (Keppler et al., 2005). The Arctic tundra appears to be a net 

sink for CH3Cl, with uptake rates increasing under drier 

conditions (Rhew et al., 2007; Teh et al., 2009). However, 

measurements at a sub-Arctic wetland show a small net 

source for CH3Cl (Hardacre et al., 2009). The causes of 

such geographic differences are not currently understood, 

although there are significant temperature, vegetation, and 

hydrological differences between the two regions.
1.29

ODSs and Related Chemicals

Table 1-3. Global trace gas lifetimes for selected halocarbons with lifetimes longer than 0.5 years. 

(See Table 1-4 for estimates of local lifetimes for short-lived gases).

Industrial

Designation

or Common Name

Chemical

Formula

 Total Lifetime 

 from Previous 

 Assessments a

 (years)

New 

Total 

Lifetime a

(years)

OH 

Lifetime d

(years)

Stratospheric 

Lifetime

(years)

Notes

Halogenated Methanes

HFC-41 CH3F 2.4 2.8 2.9 64 2, 5, 6

HFC-32 CH2F2 4.9 5.2 5.5 89 1, 4

HFC-23 CHF3 270 222 245 2347 2, 4

PFC-14 (Carbon 

tetrafluoride)

CF4 50,000 > 50,000 > 50,000 7

Methyl chloride CH3Cl 1.0 1.0 1.5 1, 8

Carbon tetrachloride CCl4 26 26 35 see text 

and 

Box 1-2

HCFC-31 CH2ClF 1.3 1.3 1.3 38 1, 5, 9

HCFC-22 CHClF2 12.0 11.9 12.8 186 1, 4

HCFC-21 CHCl2F 1.7 1.7 1.8 33 1, 5, 6, 9

CFC-11 CCl3F 45 45 45 8

CFC-12 CCl2F2 100 100 100 8

CFC-13 CClF3 640 640 640 7

Methyl bromide CH3Br 0.7 0.8 1.9 1, see text 

Bromodifluoromethane CHBrF2 5.8 5.2 6.0 39 1, 5, 9

Halon-1301 CBrF3 65 65 65 8

Halon-1211 CBrClF2 16 16 8, 10

Halon-1202 CBr2F2 2.9 2.9 8, 11

Halogenated Ethanes

HFC-152a CH3CHF2 1.4 1.5 1.6 45.4 2, 4, 6

HFC-143 CH2FCHF2 3.5 3.5 3.7 73 1, 5

HFC-143a CH3CF3 52 47.1 55 327 1, 4

HFC-134 CHF2CHF2 9.6 9.7 10.5 134 1, 5

HFC-134a CH2FCF3 14.0 13.4 14.3 232 1, 4, 6

HFC-125 CHF2CF3 29 28.2 32 246 1, 4, 6

PFC-116 

(Perfluoroethane)

CF3CF3 10,000 > 10,000 > 10,000 7

Methyl chloroform CH3CCl3 5.0 5.0 6.1 e 39 1, 8

HCFC-141b CH3CCl2F 9.3 9.2 10.7 64.9 1, 4, 6

HCFC-142b CH3CClF2 17.9 17.2 19.3 160 1, 4, 6

HCFC-133a CH2ClCF3 4.3 4.5 72 2, 5, 9

HCFC-123 CHCl2CF3 1.3 1.3 1.4 35.6 1, 4, 6

HCFC-123a CHClFCF2Cl 4.0 4.3 63 1, 5, 9
1.30

Chapter 1

Industrial

Designation

or Common Name

Chemical

Formula

 Total Lifetime 

 from Previous 

 Assessments a

 (years)

New 

Total 

Lifetime a

(years)

OH 

Lifetime d

(years)

Stratospheric 

Lifetime

(years)

Notes

HCFC-123b CHF2CCl2F 6.2 ~7 52 5, 9, 12

HCFC-124 CHClFCF3 5.8 5.9 6.3 111 1, 4, 6

HCFC-124a CHF2CClF2 9.1 ~10 107 5, 9, 12

CFC-113 CCl2FCClF2 85 85 85 8

CFC-113a CCl3CF3 ~45 ~45 13

CFC-114 CClF2CClF2 300 190 190 34

CFC-114a CCl2FCF3 ~100 ~100 14

CFC-115 CClF2CF3 1,700 1020 1020 34

Halon-2311 

(Halothane)

CHBrClCF3 1.0 1.1 11 1, 5, 9

Halon-2402 CBrF2CBrF2 20 20 8, 10

Halogenated Propanes

HFC-263fb CH3CH2CF3 1.6 1.2 1.2 38 2, 5

HFC-245ca CH2FCF2CHF2 6.2 6.5 6.9 105 1, 5

HFC-245ea CHF2CHFCHF2 4.0 3.2 3.4 70 2, 5

HFC-245eb CH2FCHFCF3 4.0 3.1 3.3 69 2, 5

HFC-245fa CHF2CH2CF3 7.6 7.7 8.2 116 1, 5

HFC-236cb CH2FCF2CF3 13.6 13.1 14.3 160 2, 5

HFC-236ea CHF2CHFCF3 10.7 11.0 11.9 144 1, 5

HFC-236fa CF3CH2CF3 240 242 253 5676 1, 4

HFC-227ea CF3CHFCF3 34.2 38.9 44.5 310 2, 5

PFC-218 

(Perfluoropropane)

CF3CF2CF3 2,600 2,600 2,600 8

PFC-c216 

(Perfluorocyclopropane)

c-C3F6 > 1,000 ~3,000 ~3,000 15

HCFC-243cc CH3CF2CCl2F 26.4 19.5 27.1 70 1, 5, 9

HCFC-234fb CF3CH2CCl2F 49 117 84 2, 5, 9

HCFC-225ca CHCl2CF2CF3 1.9 1.9 2.0 43.7 1, 4

HCFC-225cb CHClFCF2CClF2 5.8 5.9 6.3 101 1, 4

Halogenated Higher Alkanes

HFC-365mfc CH3CF2CH2CF3 8.6 8.7 9.3 125 1, 5

HFC-356mcf CH2FCH2CF2CF3 1.2 1.3 1.3 40 1, 5

HFC-356mff CF3CH2CH2CF3 8.1 8.3 8.9 122 1, 5

HFC-338pcc CHF2CF2CF2CHF2 12.3 12.9 14.0 159 1, 5

HFC-329p CHF2CF2CF2CF3 28.4 32 256 5, 16

PFC-C318 

(Perfluorocyclobutane)

c-C4F8 3,200 3,200 3,200 7

Table 1-3, continued.
1.31

ODSs and Related Chemicals

Industrial

Designation

or Common Name

Chemical

Formula

 Total Lifetime 

 from Previous 

 Assessments a

 (years)

New 

Total 

Lifetime a

(years)

OH 

Lifetime d

(years)

Stratospheric 

Lifetime

(years)

Notes

PFC-31-10 

(Perfluorobutane)

C4F10 2,600 2,600 2,600 7

HFC-43-10mee CF3CHFCHFCF2CF3 15.9 16.1 17.9 157 1, 4

HFC-458mfcf CF3CH2CF2CH2CF3 23.2 22.9 25.5 224 1, 5

PFC-41-12 

(Perfluoropentane)

C5F12 4,100 4,100 4,100 7

HFC-55-10mcff CF3CF2CH2CH2CF2CF3 7.7 7.5 8.0 115 1, 5

HFC-52-13p CHF2CF2CF2CF2CF2CF3 32.2 36.4 275 3, 5

PFC-51-14 

(Perfluorohexane)

C6F14 3,200 3,100 3,100 7

PFC-61-16 

(Perfluoroheptane)

C7F16 ~3,000 ~3,000 17

Perfluorodecalin C10F18 2,000 ~2,000 ~2,000 18

Fluorinated Alcohols

1,1,1,3,3,3-

hexafluoroisopropanol

(CF3)2CHOH 2.0 1.9 2.0 51 1, 5

Halogenated Ethers

HFE-143a CH3OCF3 4.3 4.8 5.1 88 1, 5

HFE-134 CHF2OCHF2 26 24.4 28.4 240 1, 5

HFE-125 CHF2OCF3 136 119 147 620 1, 5

HFE-227ea CF3OCHFCF3 11 b 51.6 60 370 2, 5

HCFE-235da2 

(Isoflurane)

CHF2OCHClCF3 2.6 3.5 3.7 65 2, 5, 9

HFE-236ea2 

(Desflurane)

CHF2OCHFCF3 5.8 b 10.8 11.7 143 2, 5

HFE-236fa CF3OCH2CF3 3.7 b 7.5 ~8 115 5, 19

HFE-245fa1 CF3OCH2CHF2 2.2 b 6.6 ~7 106 5, 20

HFE-245fa2 CHF2OCH2CF3 4.9 5.5 5.8 95 2, 5

HFE-245cb2 CH3OCF2CF3 5.1 4.9 5.2 89 1, 5

HFE-254cb2 CH3OCF2CHF2 2.6 2.5 2.6 60 1, 5

HFE-236ca CHF2OCF2CHF2 26.5 20.8 23.1 212 2, 5

HCFE-235ca2 

(Enflurane)

CHF2OCF2CHFCl 4.3 4.6 62 5, 9, 21

CF3CF2OCF2CHF2 6.8 b 22.5 20–30 222 5, 22

CF3CF2OCH2CF3 4.3 b 7.5 ~8 115 5, 23

CH3OCF2CF2CF3 5.2 5.0 5.3 90 1, 5

CF3CF2OCH2CHF2 2.8 b 6.6 ~7 106 5, 24

CHF2OCH2CF2CF3 5.9 5.7 6.0 97 1, 5

Table 1-3, continued.
1.32

Chapter 1

Industrial

Designation

or Common Name

Chemical

Formula

 Total Lifetime 

 from Previous 

 Assessments a

 (years)

New 

Total 

Lifetime a

(years)

OH 

Lifetime d

(years)

Stratospheric 

Lifetime

(years)

Notes

CF3CH2OCF2CHF2 7.1 6.0 6.4 100 5, 25

CH3OCF2CHFCF3 0.94 b ~3 ~3 77 5, 26

CH3OCF2CF2CHF2 0.93 b ~3 ~3 77 5, 26

CHF2CH2OCF2CHF2 2.0 b 5.7 ~6 97 5, 27

CHF2OCH2CF2CHF2 3.6 3.5 3.7 73 1, 5

HFE-347 isomer

(Sevoflurane)

(CF3)2CHOCH2F 2.2 2.3 56 5, 28

HFE-338 isomer (CF3)2CHOCHF2 3.1 b 21.2 23.5 214 2, 5

(CF3)2CFOCH3 3.4 3.7 3.9 75 2, 5

CH3O(CF2)3CF3 5 4.7 5.0 87 2, 5

HFE-54-11 isomer CF3CHFCF2OCH2CF2CF3 9.1 8.8 9.5 127 5, 29

CH3CH2O(CF2)3CF3 0.77 0.8 0.8 30 5, 30

CHF2OCF2OCHF2 12.1 c 25.0 28.0 237 2, 5

CHF2OCF2CF2OCHF2 6.2 c 12.9 14.0 159 2, 5

CHF2OCF2OCF2CF2OCHF2 6.3 c 13.5 14.7 163 2, 5

CF3OC(O)H 3.6 < 3.5 3.7 73 5, 31, 33

C2F5OC(O)H 3.6 < 3.5 3.7 73 5, 32, 33

n-C3F7OC(O)H 2.6 < 2.6 2.7 61 5, 32, 33

Other Fluorinated Compounds

TrifluoromethylsulfurpentafluorideSF5CF3 650–950 650–950 650–950 8

Sulfur hexafluoride SF6 3,200 3,200 3,200 7

Nitrogen trifluoride NF3 740 500 500 34

Sulfuryl fluoride SO2F2 36 > 300 630 35

a Includes OH reaction, ocean loss, and stratospheric loss (reactive and photolytic) as appropriate.

b Lifetime estimated from theoretically calculated reaction rate constants.

c Lifetime estimated using room temperature data.

d Lifetime with respect to reaction with 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).

e The value of τOH of 6.1 years for methyl chloroform was derived from its measured overall lifetime of 5.0 years (Clerbaux and Cunnold et al., 2007; Prinn et al., 2005) taking into account an ocean loss of 89 years and stratospheric loss of 

39 years.

_____________

Notes:

1. OH rate constants taken from JPL 06-2.

2. OH rate constants taken from JPL 10-6. JPL 10-6 is cited here whenever there is a change in a rate constant recommendation or the accompanying note. It does not necessarily mean that a major change was recommended for a rate 

constant. Nevertheless, updates in JPL 10-6 reflect improved kinetic understanding.

Table 1-3, continued.
1.33

ODSs and Related Chemicals

3. OH rate constants taken from Atkinson et al. (2008).

4. Stratospheric lifetime taken from Naik et al. (2000).

5. Stratospheric reaction lifetime estimated from an empirical correlation between the tropospheric and stratospheric 

lifetimes that were reported by Naik et al. (2000) for HFCs for which OH and O(1D) reaction rate constants were available.

6. Total lifetime includes a very minor contribution for ocean loss derived in Yvon-Lewis and Butler (2002).

7. Total lifetime from Ravishankara et al. (1993). The lifetimes for CF4, CF3CF3, and C4F10 were reported as lower limits.

8. Total lifetime taken from Table 1-4 in Clerbaux and Cunnold et al. (2007).

9. Stratospheric lifetime includes a photolysis lifetime estimated from that of a fully halogenated compound containing 

equivalent Cl and/or Br groupings.

10. Lifetime is due to a combination of tropospheric and stratospheric photolysis.

11. Lifetime is due primarily to tropospheric photolysis with a smaller contribution associated with the stratospheric 

photolysis.

12. OH rate constant estimated relative to CHF2CF3 using available data on the effects of Cl substitution.

13. Total lifetime assumed to be the same as CFC-11.

14. Total lifetime assumed to be the same as CFC-12.

15. Total lifetime assumed to be comparable to those for perfluoropropane, perfluorobutane, and perfluorocyclobutane.

16. OH rate constants taken from Young et al. (2009).

17. Total lifetime assumed to be the same as perfluorohexane.

18. Total lifetime estimated by Shine et al. (2005).

19. OH lifetime estimated by adjusting the lifetime of CF3CH2OCF2CHF2 for the contribution of -CF2CHF2.

20. OH lifetime estimated relative to CF3OCH3 and CHF2CH2CF3.

21. OH rate constants taken from Tokuhashi et al. (1999).

22. OH lifetime estimated as greater than CHF2CF2OCHF2 and less than CHF2CF2CF2CF3.

23. OH lifetime assumed to be the same as CF3OCH2CF3.

24. OH lifetime assumed to be the same as CHF2CH2OCF3.

25. OH rate constants taken from Wilson et al. (2007).

26. OH and total lifetimes estimated as approximately the same as for CH3OCF2CHF2.

27. OH lifetime estimated relative to CF3CF2OCF2CHF2 and CF3CF2OCH2CHF2.

28. OH rate constant taken from Langbein et al. (1999) using an estimated temperature dependence.

29. OH rate constants taken from Chen et al. (2005a).

30. OH rate constants taken from Christensen et al. (1998) using an estimated temperature dependence.

31. OH rate constants taken from Chen et al. (2004a).

32. OH rate constants taken from Chen et al. (2004b).

33. Ocean loss for perfluoro esters has been estimated from hydrolysis and solubility data for non-fluorinated and partially 

fluorinated esters by Kutsuna et al. (2005). These authors suggest that the ocean sink can be comparable to the tropospheric reaction sink for perfluoro ethers, thereby reducing the total lifetimes given in this table by as much as a factor 

of 2.

34. From Prather and Hsu (2008) with the lifetimes recalculated using the JPL 06-2 recommended rate constants for the 

O(1D) reactions corrected for the deactivation channel (see also Section 1.2.2).

35. From Papadimitriou et al. (2008b) and Mühle et al. (2009). The total lifetime is primarily due to ocean uptake.

Table 1-3, continued (notes).
1.34

Chapter 1

Box 1-3. Atmospheric Lifetimes and Removal Processes

The total atmospheric lifetime or turnover time (τtotal) of a trace gas is the time required to remove or chemically transform approximately 63% (i.e., 1−1/e) of its global atmospheric burden (Cglobal). τtotal can be defined as the 

ratio of the burden to the total removal rate (Ltotal) from the atmosphere:

τtotal = Cglobal / Ltotal

Ltotal can be broken down into the sum of loss rates associated with various removal processes

Ltotal = Latm + Lsoil + Locean + ...LX

where Latm is the loss rate for processes occurring in the atmosphere, Lsoil is the loss rate for soil uptake, Locean is the 

rate for loss to the oceans, and LX is the loss rate for any other removal process, X. Each of these loss rates is associated with a characteristic lifetime such that

Li = Cglobal / τi

Therefore,

Ltotal = Cglobal / τtotal = Cglobal / τatm + Cglobal / τsoil + Cglobal / τocean +... Cglobal / τX

and τtotal can be expressed as

(τtotal)

−1

 = (τatm)

−1

 + (τsoil)

−1

 + (τocean)

−1

 + ...(τX)

−1

The atmospheric loss rate is typically broken down into separate loss rates associated with processes occurring in the 

troposphere and in the stratosphere

Latm = Ltrop + Lstrat = Cglobal / τtrop + Cglobal / τstrat

such that

(τatm)

−1

 = (τtrop)

−1

 + (τstrat)

−1

The tropospheric lifetimes for most trace gases are primarily controlled by reactions with hydroxyl radicals (OH), 

although reactions with ozone (O3), nitrate radical (NO3), and chlorine radicals (Cl) can also play limited roles. UV 

photolysis in the troposphere, designated by the characteristic lifetime τJ, is the dominant sink for iodine-containing 

compounds and can also be important for some bromine-containing compounds. Hydrolysis-related removal processes (such as wet deposition, rainout, or washout), characterized collectively by τhydrol, can also contribute. Thus,

(τtrop)

−1

 = (τOH)

−1

 + (τO3)

−1

 + (τCl)

−1

 + (τNO3)

−1

 + (τhydrol)

−1

 + (τJ)

−1

Compounds whose tropospheric lifetimes are longer than the characteristic mixing time in the troposphere are designated as “long lived” in this chapter. For such well-mixed compounds, lifetimes due to reaction with tropospheric 

hydroxyl radicals can be estimated relative to the corresponding lifetime of methyl chloroform (Spivakovsky et al., 

2000)

where is the partial OH lifetime for the compound of interest (RH), kRH(272 K) and kMCF(272 K) are the rate 

constants for the reactions of OH with RH and with methyl chloroform (MCF) at T = 272 K, and = 6.1 years 

(see Table 1-3, footnote “e”).

Stratospheric loss processes, such as reactions with OH, O3, and excited state oxygen atoms (O(1D)), and UV 

RH

OH OH

MCF

(272 K)

kMCF (272 K)

kRH

τ = τ

RH τ OH

OH

MCF

τ
1.35

ODSs and Related Chemicals

1.2.2 Loss Processes

Halocarbons are removed from the atmosphere by 

a number of processes including oxidation by hydroxyl 

radicals, photolysis, uptake by the ocean, and uptake by 

terrestrial ecosystems. The combined influence of these 

processes can be expressed as a lifetime (see Box 1-3). 

Lifetimes for ODSs are generally based on a combination 

of modeling studies and analysis of observations, particularly stratospheric observations. Updates to our understanding of the persistence, expressed as lifetimes, of halocarbons in the environment are based on the consideration 

of updated kinetic information (see Table 1-3). Uncertainties associated with these lifetimes stem from uncertain hydroxyl radical concentrations and distributions, and 

uncertain reaction rate constants. Modeling uncertainties 

related to absorption cross sections and stratospheric air 

transport rates also affect the calculation of lifetimes. This 

is particularly important for CFCs, for example, as this 

represents their primary sink.

New results from models that more accurately 

simulate air transport rates through the stratosphere than 

older models suggest a steady-state lifetime for CFC-11 

of 56–64 years (Douglass et al., 2008), notably longer 

than the 45 years used in recent past Assessments. This 

finding has the potential to affect lifetimes, ODPs, and 

emissions derived from global measured trends for some 

other ODSs. For example, while the CFC-12 lifetime calculated in this same study was not substantially different 

from its current, best-estimate lifetime of 100 years, the 

stratospheric lifetime of CCl4 is calculated as ~50 years 

in this same model (Rontu Carlon et al., 2010), which is 

notably longer than the current best estimate of 35 years. 

A revision to the CFC-11 or CCl4 lifetime is not recommended in this Assessment, pending a comprehensive 

analysis by the atmospheric observation and modeling 

communities.

New modeling calculations by Prather and Hsu 

(2008) have yielded significantly shorter lifetimes for 

CFC-114 and CFC-115. The O(1D) (excited-state oxygen atoms) reaction rate coefficients used in their calculation, however, included both the chemical oxidation of the 

CFCs as well as the nonreactive, physical quenching of 

O(1D). This resulted in an overestimation of loss for these 

CFCs from O(1D). When the O(1D) rates for only reactive channels are considered (Sander et al., 2006, hereafter 

referred to as JPL 06-2), lifetimes of 190 years for CFC114 and 1020 years for CFC-115 are calculated (Table 

1-3; Prather and Hsu, 2010), which are still significantly 

shorter than considered in past Assessments (300 years 

for CFC-114 and 1700 years for CFC-115; Clerbaux and 

Cunnold et al., 2007).

A significantly shorter lifetime for nitrogen trifluoride (NF3) was also calculated by Prather and Hsu (2008) 

using an older O(1D) reaction rate constant (Sorokin et 

al., 1998). The NF3 lifetime given in this Assessment 

(500 years compared to 740 years in Clerbaux and Cunnold et al., 2007) is based on the Prather and Hsu (2008) 

results revised using the O(1D) + NF3 rate constant recommended in JPL 06-2. A new study of this rate constant 

has been published by Zhao et al. (2010). Use of the rate 

constant reported by these authors would reduce the lifephotolysis can also play important roles in dictating the lifetime of a long-lived compound. The contribution from 

photolysis at Lyman-α (121.6 nm), which is important only for very long-lived compounds, can be included here, although it occurs at altitudes above the stratosphere (≥70 km). UV and Lyman-α photolysis are the most important loss 

processes for nonreactive compounds:

(τstrat)

−1

 = (τOH)

−1

 + (τO3)

−1

 + (τO(1D))

−1

 + (τJ)

−1

 + (τLyman-α)

−1

For long-lived compounds, τtotal can be considered to be a global lifetime that is representative of the compound’s 

persistence in Earth’s atmosphere.

In the context of this chapter, very short-lived substances (VSLS) are compounds with lifetimes less than or 

equal to 0.5 years (i.e., shorter than the characteristic time of mixing processes in the troposphere) (see Section 1.3). 

The actual atmospheric lifetimes of VSLS depend on the emission location and season as well as on local atmospheric 

conditions. Consistent with past Assessments, for VSLS was estimated with hydroxyl radical concentrations 

[OH] = 1 × 106 molecule/cm3 and an OH reaction rate constant at T = 275 K:

 = 1/(kRH(275 K) × [OH])

The use of a mean global tropospheric OH concentration in such calculations provides only approximate “local lifetimes.” The concept of a single global lifetime, Ozone Depletion Potential (ODP), or Global Warming Potential 

(GWP) is inappropriate for such short-lived gases (see Section 1.3).

RH τ OH

RH τ OH
1.36

Chapter 1

time for NF3 by about 7% from the 500 years reported in 

this chapter.

The principal differences between these new (revised) CFC-114 and CFC-115 lifetimes and those recommended in earlier Assessments (based on Ravishankara et 

al., 1993) are primarily due to improved photolysis models 

(see Chapter 6 in SPARC CCMVal, 2010) that include updated solar fluxes in the Herzberg continuum (200–220 

nanometers (nm)) and in the Schumann-Runge bands 

(<205 nm). Note that the more rapidly photolyzed CFCs 

(-11, -12, and -113) are destroyed in the lower stratosphere 

at wavelengths whose opacities are dominated by the molecular oxygen (O2) Herzberg continuum, while the much 

longer-lived gases (CFC-114, -115, and NF3) are lost primarily between altitudes of 32 to 40 km and in the wavelength region where the opacity is controlled by the O2

Schumann-Runge bands. This is why the CFC-11, CFC12, and CFC-113 lifetimes are not appreciably influenced 

by use of the improved photolysis model.

The Prather and Hsu (2008) model did not include 

losses above 70 km, where photolysis at Lyman-alpha 

(121.6 nm) wavelengths occurs. This loss is likely important for some very long-lived chemicals, particularly 

CFC-115. The lifetime for CFC-13 derived by Ravishankara et al. (1993) remains unchanged here (640 years) despite these considerations because atmospheric loss of this 

chemical is primarily associated with reactive removal by 

O(1D) atoms in the stratosphere. For NF3, the Prather and 

Hsu (2008) calculations revise earlier lifetime estimates 

that were based on a one-dimensional (1-D) photolysis 

calculation (not including O(1D) reactions) referenced as a 

private communication in Molina et al. (1995).

Other considerations also affect the usefulness of 

the stratospheric lifetimes calculated in Table 1-3. These 

lifetimes are derived for steady-state conditions, yet global 

atmospheric mixing ratios of most ODSs are not at steady 

state. For CFCs and other compounds destroyed primarily 

by photolysis in the stratosphere, steady-state lifetimes are 

generally longer than lifetimes in the absence of significant emissions (e.g., Prather, 1997). Finally, many models suggest that future changes in stratospheric circulation 

expected from enhanced greenhouse gas concentrations 

will enhance photolytic loss rates for the long-lived halocarbons. This influence is predicted to result in shorter 

stratospheric lifetimes than those listed in Table 1-3.

Since the previous Assessment, considerably more 

kinetic data have become available for the reactions of OH 

with a wide variety of CFC and halon substitutes, especially fluorinated alkenes, ethers, and alcohols for which 

limited or no data previously existed. A comprehensive 

analysis of available kinetic data since the last Assessment 

has revealed a number of compounds for which our understanding of losses has improved.

Most notably, lifetimes of three hydrofluoroethers 

(HFEs) are now based on experimentally measured OH 

reaction rate constants (HFE-227ea, HFE-236ea2, and 

(CF3)2CHOCHF2) (Sander et al., 2010; hereafter referred 

to as JPL 10-6). Previous estimates were based on theoretically calculated OH reaction rate constants that significantly overestimated the loss rates for these gases, leading 

to unrealistically short lifetimes.

Rate constants for eight other HFEs (HFE-236fa, 

HFE-245fa1, CF3CF2OCF2CHF2, CF3CF2OCH2CF3, 

CF3CF2OCH2CHF2, CH3OCF2CHFCF3, CH3OCF2CF2CHF2, 

CHF2CH2OCF2CHF2) that had also been derived from 

theoretical considerations can now be better estimated 

from the kinetic data on fluoroethers that are now available (Table 1-3). The revised lifetimes for these HFEs 

are considerably longer than recommended in the previous 

Assessment.

Updated lifetimes for three additional HFEs 

(CHF2OCF2OCF2CF2OCHF2, CHF2OCF2OCHF2, and 

CHF2OCF2CF2OCHF2) have been derived using the temperature dependencies of the OH reaction rate constants 

estimated in JPL 10-6 (which are based on existing data 

for other fluoroethers) and evaluated at 272 K (see Box 

1-3). Previously assessed lifetime estimates for these three 

HFEs were based on rate constants at 298 K (Ramaswamy 

et al., 2001), which are ~80% faster.

Significant revisions have also been made to the 

lifetimes of three HFCs:

• The updated total lifetime for HFC-23 (CHF3) of 222 

years is ~20% shorter than the 270 years recommended in the previous Assessment (Clerbaux and Cunnold 

et al., 2007). This change in total lifetime results from 

a shorter tropospheric lifetime due to the OH rate constant recommended in JPL 10-6 being ~15% larger 

than given in JPL 06-2 and a stratospheric lifetime 

of 2347 years based on the model calculation of Naik 

et al. (2000).

• The new total lifetime of HFC-227ea (CF3CHFCF3) 

is longer than given in the previous Assessment 

(38.9 years vs. 34.2 years) primarily because the updated rate constant (JPL 10-6) is smaller than previously recommended due to the inclusion of new 

temperature-dependent kinetic data (Tokuhashi et al., 

2004).

• The revised total lifetime of HFC-41 (CH3F) is 2.8 

years. The revision is based on the current recommendation for the OH reaction rate constant (JPL 

10-6) and an estimation of stratospheric loss (64 years 

derived from Naik et al. (2000); Table 1-3).
1.37

ODSs and Related Chemicals

1.3 VERY SHORT-LIVED HALOGENATED 

SUBSTANCES (VSLS)

While the longer-lived ODSs account for the majority of the halogen loading in the present-day stratosphere, there is evidence suggesting that halogenated 

very short-lived substances (VSLS) with lifetimes of 0.5 

years or less and their degradation products contribute to 

halogen in the stratosphere (Ko and Poulet et al., 2003; 

Law and Sturges et al., 2007). In Table 1-4, some of the 

halogenated VSLS considered in this section are listed together with their lifetimes. The abbreviations for chemical and meteorological parameters used in this section are 

explained in Box 1-4).

While some VSLS originate from anthropogenic 

sources, others are of partly or wholly natural origins. 

We include discussions of naturally emitted VSLS here 

for two reasons. First, much of our current understanding 

of the mode and rate of delivery of the VSLS—whether 

natural or anthropogenic—to the stratosphere arises from 

studies of naturally occurring halocarbons. Second, the 

emission and the stratospheric injection of natural halocarbons might not be constant in a changing climate.

1.3.1 Emissions, Atmospheric 

Distributions, and Abundance 

Trends of Very Short-Lived 

Source Gases

Since the 2006 Assessment (Law and Sturges et al., 

2007) new findings on emissions of anthropogenic and natural VSLS have been made as well as on trends of anthropogenic chlorinated VSLS. New data are also available 

on very short-lived (VSL) source gases (SGs) in the upper 

troposphere and the importance of product gases (PGs) is 

assessed.

1.3.1.1 Chlorine-Containing Very

Short-Lived Source Gases

A number of chlorinated VSLS from anthropogenic 

and natural sources have been measured in the background 

atmosphere. Industrial sources of these gases are discussed 

in McCulloch et al. (1999), Simmonds et al. (2006), Worton 

et al. (2006), and on the web site of the U.S. Environmental 

Protection Agency (www.epa.gov; Toxic Air Pollutants). 

Dichloromethane (methylene chloride, CH2Cl2) is principally used as a paint remover, and also in foam production 

and foam blowing applications, as a solvent and degreaser, 

and in fumigation. Trichloromethane (chloroform, CHCl3) 

is used almost exclusively in the production of HCFC-22 

and fluoropolymers, but is also released as a by-product 

from chlorine bleaching in the paper and pulp industry, 

and from water chlorination. Tetrachloroethene (perchloroethylene, CCl2CCl2) is used in the textile industry, in 

dry-cleaning applications, and in vapor degreasing of metals, as is trichloroethene (trichloroethylene, CHClCCl2). 

1,2-Dichloroethane (CH2ClCH2Cl) is principally used in 

the production of polymers and rubbers, and is also used as 

a solvent, as a fumigant, and was historically used widely 

as an additive to leaded gasoline. Chloroethane (C2H5Cl) 

is mostly used in the manufacture of ethyl cellulose, dyes, 

and pharmaceuticals and was formerly used in the manufacture of tetraalkyl lead additives for gasoline. Other related compounds include a number of mixed-halogenated 

compounds, notably the bromochloromethanes. These latter compounds are believed to originate largely from natural sources and are generally present at lower abundances 

than the aforementioned chlorinated compounds. They are 

discussed in Section 1.3.1.2 in the context of brominated 

source gases.

Estimates of total global and industrial emissions of 

chlorinated VSLS are shown in Table 1-5 for those gases 

for which such estimates exist. From this table it can be 

seen that total CCl2CCl2 and CH2Cl2 emissions derived 

from a “top-down” analysis of atmospheric observations 

are similar to independently estimated industrial emissions derived from inventories. Although not all of the 

estimates coincide in time (e.g., industry inventories are 

mostly from the early or mid-1990s, whereas the atmospheric measurements used in the emissions models are 

mostly from several years later), the results suggest that 

most emissions of CCl2CCl2 and CH2Cl2 arise from anthropogenic activity. A model analysis of the atmospheric 

data (Simmonds et al., 2006) also indicates that more than 

90% of the emissions for these two compounds emanate 

from the Northern Hemisphere (NH), further indicative of 

their predominantly industrial origins. The industrial inventories, however, indicate that only 1% of industrial emissions are in the Southern Hemisphere (SH) (McCulloch et 

al., 1999), suggesting that other sources such as biomass 

burning might contribute. An earlier estimate of the biomass burning contribution to global atmospheric CH2Cl2

(7% according to Keene et al., 1999) is supported by Xiao 

(2008), but both are based largely on the same underlying 

data. These high values have been questioned by Simmonds et al. (2006) on the basis of the lower (by two orders of magnitude) biomass burning emission factors they 

observed from Australian wild fires.

Industrial emissions of chloroform were estimated to contribute 25 to 29% of total emissions in 2001 

by modeling atmospheric measurements (Worton et al., 

2006). This is less than that modeled in earlier years due 

to a rapid decline in emissions attributed to the paper and 

pulp industry. The absolute emissions estimate from this 

study for 2001 is, however, higher than an earlier industry 

inventory figure given by McCulloch et al. (2003) (see 
1.38

Chapter 1

Table 1-4. Lifetimes for very short-lived halogenated source gases.

Compound

Local Lifetime 

from Previous 

Assessments 

(τlocal), days

OH Lifetime1

(τOH), days

Photolysis 

Lifetime from 

Previous 

Assessments 

(τlocal), days

New Local 

Lifetime,

(τlocal), days

Notes

Chlorocarbons

CH2Cl2 140 144 > 15000 144 2, 8

CHCl3 150 149 > 15000 149 2, 8

CH3CH2Cl 30 39 39 2

CH2ClCH2Cl 70 65 65 4

CH3CH2CH2Cl 14 14 5

CHClCCl2 4.9 > 15000 4.9 3, 8

CCl2CCl2 99 90 90 3

CH3CHClCH3 18 18 5

Bromocarbons

CH2Br2 120 123 5000 123 2, 8

CHBr3 26 76 36 24 2, 8

CH2BrCl 150 137 15000 137 2, 8

CHBrCl2 78 121 222 78 6, 8

CHBr2Cl 69 94 161 59 7, 8

CH3CH2Br 34 41 41 2

CH2BrCH2Br 55 70 70 4

n-C3H7Br 13 12.8 > 1200 12.8 3, 8

iso-C3H7Br 16.7 16.7 3

Iodocarbons

CH3I 7 158 7 (4–12) 7 4, 8

CF3I 4 860 not determined 4 2

CH2ClI 0.1 0.1 0.1 8

CH2BrI 0.04 0.04 0.04 8

CH2I2 0.003 0.003 0.003 8

CH3CH2I 4 17.5 5 (2–8) 4 4, 8

CH3CH2CH2I 0.5 7.7 0.6 (0.5–1.5) 0.5 4, 8

CH3CHICH3 1.2 9.7 1.4 (1–3) 1.2 4, 8

CF3CF2CF2I < 2 < 2 9

Notes:

1. These local OH lifetimes are calculated using an average tropospheric OH concentration of 1×106 molecule/cm3 and the OH reaction rate constant at 

T = 275 K. Local lifetimes quoted here are not meant to be estimates of global lifetimes, which for short-lived gases depend on the emission location 

and season as well as local atmospheric conditions. The concept of a single global lifetime, ODP, or GWP is inappropriate for such short-lived gases.

2. OH reaction rate constant taken from JPL 06-2.

3. OH reaction rate constant taken from JPL 10-6. JPL 10-6 is cited here whenever there is a change in a rate constant recommendation or the accompanying note. It does not necessarily mean that a major change was recommended for a rate constant. Nevertheless, updates in JPL 10-6 reflect 

improved kinetic understanding.

4. OH reaction rate constant taken from Atkinson et al. (2008).

5. OH reaction rate constant taken from Yujing and Mellouki (2001).

6. Room temperature OH reaction rate constant taken from Bilde et al. (1998). The temperature dependence of the OH reaction rate constant was 

estimated from the reactivity of CHCl3, CHCl2Br, and CHBr3.

7. OH reaction rate constant estimated from the OH reactivity of CHCl3, CHCl2Br, and CHBr3.

8. Photolysis lifetime taken from Table 2-4 in Ko and Poulet et al. (2003).

9. Photolysis lifetime estimated from comparison of UV spectra and lifetimes for CF3I and CF3CF2CF2I.
1.39

ODSs and Related Chemicals

Table 1-5). Biomass burning is thought to contribute 1% 

or less to emissions of CHCl3, CCl2CCl2, and CHClCCl2

(Table 1-5).

Annual mean mole fractions and trends of some of 

these chlorinated VSL source gases at Earth’s surface during recent years are given in Table 1-6. CH2Cl2 has increased significantly in recent years in both hemispheres, 

having previously fallen from 1995 (when records began) 

to about 2000 (updated from Simmonds et al., 2006). In 

the NH, mean surface mixing ratios in 2008 were similar to those previously observed in the mid-1990s. CHCl3

concentrations have remained approximately constant 

since 2000 suggesting little change in the fraction arising 

from industrial emissions. CCl2CCl2 has decreased almost 

monotonically since the late 1980s (updated from Simpson et al., 2004). Data in Table 1-6 indicate that this trend 

may have reversed in the most recently reported year, but 

this may be within the scatter of the interannual variability 

of the measurements. Global long-term measurements are 

available for CHClCCl2 for the first time (Table 1-6). The 

large standard deviation reflects the very large seasonal 

cycle of this very reactive gas. The short record and large 

annual variability do not yet allow a trend to be established 

for this compound.

There are very few reported lower tropospheric 

measurements of the other chlorinated VSLS since the 

last Assessment, and likewise no industrial emission estimates are available for these species. In the last Assessment (Law and Sturges et al., 2007) one citation (Low et 

al., 2003) gave mean NH and SH C2H5Cl mixing ratios of 

2.6 and 1.6 ppt, respectively, which, given the short lifetime of this gas (39 days; Table 1-4), requires a significant, 

presumably nonindustrial, source to sustain the observed 

SH abundances. Higher values were noted in the tropical 

Box 1-4. Definition of Acronyms Related to Short-Lived Gases

Chemical Parameters

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 halogenated organic and inorganic degradation products and 

their injection into the stratosphere

Bry total inorganic stratospheric bromine (e.g., HBr, BrO, 2×Br2) resulting from degradation of brominecontaining organic source gases (halons, methyl bromide, VSLS), and natural inorganic bromine sources 

(e.g., volcanoes, sea salt, and other aerosols)

Cly total inorganic stratospheric chlorine (e.g., HCl, ClO) resulting from degradation of chlorine-containing 

source gases (CFCs, HCFCs, organic VSLS), and natural inorganic chlorine sources (e.g., sea salt and 

other aerosols)

Iy total inorganic stratospheric iodine (e.g., IO, OIO, HOI) resulting from degradation of iodine-containing 

source gases (VSLS), and natural inorganic sources (e.g., sea salt and other aerosols)

Bry

VSLS the component of stratospheric Bry from the degradation of organic brominated VSL SGs and tropospheric 

inorganic bromine sources (also called “additional” stratospheric Bry)

Meteorological Parameters

BDC Brewer-Dobson circulation

LMS lowermost stratosphere

TTL tropical tropopause layer—a layer exhibiting properties of both the stratosphere and troposphere. In this 

Assessment we follow the definition of the TTL as used in Law and Sturges et al. (2007). The bottom 

of TTL is taken as the region of maximum convective outflow (about 12 km altitude, or 345K potential 

temperature) and the upper end is identical to the tropical cold point tropopause (about 17 km or 380 K 

potential temperature).

z0 altitude of clear-sky zero radiative heating. As in the previous Assessment (Law and Sturges et al., 2007) 

we define the level at about 15 km or 360K, where there is a transition from clear-sky radiative cooling to 

clear-sky radiative heating. In general, air masses above this level are expected to enter the stratosphere.

CPT cold point tropopause—defined by the minimum atmospheric temperature and gives the position of the 

tropical tropopause. At this altitude air parcels enter the stratospheric overworld.

Washout the removal of trace gases from the gas phase, including subsequent falling of rain or ice particles
1.40

Chapter 1

marine boundary layer, also indicative of a natural origin. The mixing ratios given in the last two Assessments 

for CH2ClCH2Cl (20–40 ppt in the NH and 5–7 ppt in the 

SH) appear to have been incorrectly cited from an original 

paper (Class and Ballschmiter, 1987). The values given in 

this latter paper are <1 to 24 ppt for the NH and <1 ppt 

for the SH. We do not know of any globally distributed 

or long-term measurements of this species since this date, 

however, the large interhemispheric ratio suggests predominantly anthropogenic uses at that time. There are no emission estimates or inventories for this compound, making it 

a significant uncertainty in the chlorinated VSLS budget.

It is likely that significant calibration differences 

exist between measurements of chlorinated VSLS by different laboratories. This is particularly evident in the case 

of CCl2CCl2 (Table 1-6) but is equally true for all of the 

chlorocarbons discussed here (e.g., Simmonds et al., 2006; 

Butler et al., 2010). Simmonds et al. (2006) mention a 

10% difference between NOAA and AGAGE measurements for CCl2CCl2, but no intercomparison exercises 

Table 1-6. Measured mole fractions and growth rates of chlorinated very short-lived source gases.

Chemical 

Formula

Common or 

Industrial 

Name

Annual Mean Mole Fraction (ppt) Growth (2007–2008)

Laboratory a

2004 2007 2008 (ppt/yr) (%/yr)

CH2Cl2 dichloromethane 17.3 ± 5.7 20.3 ± 6.7 22.1 ± 6.7 1.8 ± 0.5 8.1 ± 2.3 AGAGE, in situ 

(Global)

CHCl3 chloroform 6.8 ± 1.2 7.0 ± 1.5 7.0 ± 1.5 −0.02 ± 0.09 −0.3 ± 1.3 AGAGE, in situ 

(Global)

CCl2CCl2 tetrachloroethene 1.9 ± 0.9 1.7 ± 0.7 1.7 ± 0.7 0.01 ± 0.00 0.6 ± 0.1 AGAGE, in situ 

(Global)

3.0 2.5 2.7 0.2 7 UCI, flask 

(Global)

CHClCCl2

b trichloroethene - 0.19 ± 0.19 0.28 ± 0.28 - - AGAGE, in situ 

(Global)

a AGAGE global mixing ratios updated from Simmonds et al. (2006). UCI global mixing ratios updated from Simpson et al. (2004). b Because of the short lifetime (4.9 days) the global mole fraction of CHClCCl2 should be seen rather as an approximation than a defined global value.

Table 1-5. Annual emissions of chlorinated VSLS.

Compound

Fraction of 

Global

Emissions in

the NH 

(%)

“Top-Down” 

Average

Annual Global 

Emissions

(Gg/yr)

Estimated

Industrial 

Emissions

(Gg/yr)

Estimated 

Biomass Burning 

Emissions 

(Gg/yr)

Origin

CH2Cl2 93

-

515 ± 22

629 ± 44

519 ± 32

430 ± 12

59

75 ± 18

1

2

CHCl3 -

62

-

-

-

370 ± 120

315–373

-

-

-

79–108

66 ± 23

2

2 ± 2

-

-

1

2

3

4 

CCl2CCl2 94 250–205 278 ± 20 0 1

CHClCCl2 - - 246 ± 13 0 1

Notes:

1. Simmonds et al. (2006) global and hemispheric emissions 1999–2003; McCulloch et al. (1999) industrial emission estimates (using the factors given 

in McCulloch et al. (1999) to estimate values appropriate for 1995); Keene et al. (1999) biomass burning emissions.

2. Xiao (2008) (values shown are for 2000–2004).

3. Worton et al. (2006) (values shown are for 2001).

4. McCulloch et al. (2003) (reference year not given).
1.41

ODSs and Related Chemicals

have yet been published for CCl2CCl2 or any other chlorinated VSLS.

The anthropogenic contribution to the tropospheric 

abundances of chlorine from VSLS can be estimated by 

assuming that the anthropogenic fraction (industrial plus 

biomass burning) of individual gases is 90% for CH2Cl2

(average of 100% and 80% from modeled global emissions from Table 1-5); 100% for CCl2CCl2, CHClCCl2, 

and CH2ClCH2Cl; and 25% for CHCl3. All other chlorinated gases are assumed to be of solely natural origin. 

Taking the averaged 2008 molar ratios in Table 1-6 and an 

average value of approximately 6 ± 6 ppt for CH2ClCH2Cl 

(see above), yields a global average surface abundance of 

Cl from anthropogenic VSLS in 2008 of 67 ± 30 ppt.

In Table 1-7, observations of a range of various 

VSLS in the marine boundary layer, the upper troposphere, 

and the tropical tropopause layer are compiled from a number of measurement campaigns (see also Section 1.3.3.1). 

These measurement campaigns cover a range of dates and 

locations and, therefore, are not directly comparable with 

one another or the surface values in Table 1-6. Nevertheless, mixing ratios in the upper troposphere are on average 

less than those measured in the marine boundary layer, as 

would be expected for measurements of VSLS at some 

distance from their sources.

Updated mixing ratios of organic chlorine from 

VSLS of 59 (36–70) ppt in the tropical upper troposphere 

(10–12 km) (Table 1-7) are similar to the value of 55 (52–

60) ppt at 10 km reported in the last Assessment, although 

with a wider spread. The main contributors to organic 

chlorine from VSLS at this altitude are CH2Cl2 and CHCl3.

Abundances at about 15 km in the tropics can be 

considered to represent the lower limit of source gas injection into the stratosphere (see Section 1.3.3.1). Taking this 

observed total of 55 (38–80) ppt Cl from all chlorinated 

VSLS at this height and the same anthropogenic fractions 

as discussed above, a minimum stratospheric injection of 

39 (27–59) ppt Cl from anthropogenic VSLS is estimated 

(Table 1-7).

These anthropogenic totals do not include the product gases hydrogen chloride (HCl) and phosgene (COCl2), 

which can be formed during atmospheric decomposition 

from both VSL SGs (primarily CHCl3 and CCl2CCl2) and 

long-lived SGs (e.g., CCl4 and CH3CCl3) and together 

could contribute an additional 0–50 ppt Cl to the stratosphere (see Section 1.3.3.3).

1.3.1.2 Bromine-Containing Very

Short-Lived Source Gases

In contrast to chlorinated VSLS, anthropogenic 

sources of brominated VSLS are small in comparison to 

natural sources. The most notable anthropogenic brominated VSLS are 1-bromopropane (n-propyl bromide, 

C3H7Br), which is used as a substitute for CFC-113 

and CH3CCl3 in metal and electronic part cleaning; and 

1,2-dibromoethane (ethylene dibromide, CH2BrCH2Br), 

which is used as a fumigant and chemical intermediate 

and was formerly widely used as an additive to leaded 

gasoline (U.S. EPA web site at www.epa.gov; Toxic Air 

Pollutants). Small emissions of trihalomethanes, notably 

CHBr3, occur from chlorination of drinking water and 

power plant cooling water (Worton et al., 2006). There 

have been very few studies of long-term trends in atmospheric abundances of brominated VSLS. In the last Assessment one study was reported indicating small increases 

in brominated trihalomethanes since the early 20th century 

from Arctic firn air, but no significant trends in dibromomethane (CH2Br2) and bromochloromethane (CH2BrCl) 

(Worton et al., 2006), which leads to the conclusion that 

the latter are of entirely natural origin. Atmospheric concentrations of brominated VSLS are highly variable due to 

the heterogeneous distribution and time-varying nature of 

sources; notably natural sources. Natural sources include 

marine phytoplankton, coastal macrophytes (seaweeds), 

and possibly terrestrial vegetation such as rice (Law and 

Sturges et al., 2007). Surface mixing ratios (predominantly from the marine boundary layer) reported in the last 

Assessment are summarized in Table 1-7. Mixing ratios 

of other bromocarbons tend to be lower than any of the 

aforementioned substances.

Most progress since the last Assessment has been 

with regard to estimating natural oceanic emissions of bromocarbons (see Table 1-8), notably bromoform (CHBr3) 

and CH2Br2. These range from 430–1400 Gg Br/yr for 

CHBr3, and 57–280 Gg Br/yr for CH2Br2 and are mostly 

larger than the estimates given in the previous Assessment 

report (Table 2-3 in Law and Sturges et al., 2007). There 

are large uncertainties associated with these estimates, 

notably due to the large variability in emissions from 

coastal zones (Butler et al., 2007). Some of the studies 

included in Table 1-8 emphasize the potential importance 

of coastal oceans (e.g., Butler et al., 2007; Carpenter et al., 

2009; Liang et al., 2010). In contrast, Palmer and Reason 

(2009) considered tropical coastal sources to be unimportant based on a supposition that seaweeds are largely absent from the tropics. Yokouchi et al. (2005), however, 

reported very high mixing ratios of CHBr3, dibromochloromethane (CHBr2Cl), and CH2Br2 associated with algalcolonized tropical shores. While numerous studies have 

identified coastal macroalgae as sources of halocarbons, 

and CHBr3 in particular (Law and Sturges et al., 2007), 

there have been relatively few studies in the tropics compared to the extratropics. Furthermore, it has not been unequivocally proven that macroalgae account for all such 

coastal emissions. Close to shore in Cape Verde in the 

tropical Atlantic, O’Brien et al. (2009) detected elevated 

mixing ratios of CHBr3 and CH2Br2 without evidence of 
1.42

Chapter 1

Table 1-7. Summary of available observations of VSLS source gases from the marine boundary layer 

(MBL) to the tropical tropopause layer (TTL). Abundances measured in the MBL are taken from the last 

Assessment (Law and Sturges et al., 2007). Data in and above the upper troposphere have been compiled 

from observations during the PEM-West A and B, TC4, Pre-AVE, and CR-AVE aircraft campaigns (see Schauffler 

et al., 1999) and from Teresina balloon observations (Laube et al., 2008 and updates). All table entries are 

mixing ratios, with units of parts per trillion (ppt).

Marine

Boundary Layer

(MBL) a

Upper

Troposphere

(UT) b Lower TTLc LZRH (z0)

c,d Upper TTLc Tropical

Tropopause c

Height

Range 10–12 km 12–14 km 14.5–15.5 km 15.5–16.5 km 16.5–17 km

Potential

Temperature

Range 

340–355 K 355–365 K 365–375 K 375–385 K

Median

(ppt)

Range e

(ppt)

Mean

(ppt)

Range e

(ppt)

Mean

(ppt)

Range e

(ppt)

Mean

(ppt)

Range e

(ppt)

Mean

(ppt)

Range e

(ppt)

Mean

(ppt) 

Range e

(ppt)

CH2Cl2 17.5 9-39 13.9 8.6–20.9 14.9 11.7–18.4 14.3 10.8–20.6 13.2 9.8–21.2 12.6 7.2–22.5

CHCl3 7.8 5.2–13.1 5.8 3.9–7.5 7.1 5.9–9.2 5.7 3.5–7.9 4.8 3.5–6.6 4.9 3.3–6.4

CH2ClCH2Cl 3.7 0.7–14.5 3.7 f 1.9–5.4 2.9 g 1.9–4.1 2.7 g 1.6–4.9 2.2 g 1.2–4.0 2.0 0.6–4.3

CHClCCl2 0.5 0.05–2 0.36 f 0–2.02 0.05 g 0.00–0.16 0.03 g 0.00–0.17 0.02 g 0.00–0.05 0.03 0.00–0.17

CCl2CCl2 1.8 1.2–3.8 1.3 0.7–1.8 1.1 0.8–1.5 0.9 0.4–1.3 0.6 0.3–0.9 0.5 0.1–1.0

CH2Br2 1.1 0.7–1.5 0.86 0.63–1.21 0.92 0.77–1.15 0.74 0.59–0.99 0.66 0.43–0.83 0.51 0.3–0.86

CHBr3 1.6 0.5–2.4 0.50 0.12–1.21 0.61 0.3–1.11 0.22 0.00–0.63 0.14 0.01–0.29 0.09 0.00–0.31

CH2BrCl 0.5 0.4–0.6 0.09 h 0.03–0.16 0.14 i 0.13–0.16 0.10 i 0.08–0.13 0.11 i 0.1–0.12 0.08 0.05–0.11

CHBr2Cl 0.3 0.1–0.8 0.11 0.01–0.36 0.10 i 0.06–0.15 0.06 i 0.03–0.11 0.05 i 0.01–0.11 0.03 0.00–0.14

CHBrCl2 0.3 0.1–0.9 0.11 0.02–0.28 0.20 0.18–0.22 0.10 0.12–0.18 0.12 0.11–0.14 0.06 0.03–0.12

Other brominated SG j < 0.2 j < 0.2 j < 0.2 j

CH3I 0.80 0.3–1.9 0.13 0.03–0.42 0.12 k 0.00–0.23 0.04 k 0.00–0.10 0.00 k 0.00–0.01 0.01 0.00–0.06

Total Cl 76 41–171 59 36–70 62 48–78 55 38–80 48 38–74 46 26–77

Anthrop. Cl l 55 27–136 43 26–51 43 32–55 39 27–59 34 25–55 32 17–58

Total Br 8.40 3.6–13.3 3.5 1.7–7.4 4.3 2.8–6.5 2.7 1.4–4.6 2.0 1.1–3.2 1.5 0.7–3.4

Total I 0.80 0.3–1.9 0.13 0.03–0.42 0.12 0.00–0.23 0.04 0.00–0.10 0.00 0.00–0.01 0.01 0.00–0.06

PEM-West = Pacific Exploratory Mission-West; TC4 = Tropical Composition, Cloud, and Climate Coupling mission; Pre-AVE = PreAura Validation Experiment; CR-AVE = Costa Rica Aura Validation Experiment. _________

Notes:

a Marine boundary layer mixing ratios are identical to those in Table 2-2 of the previous Assessment (Law and Sturges et al., 2007). b UT data are from DC-8 aircraft observations during PEM-West A, PEM-West B, and TC4. c TTL and tropopause data are from WB-57 aircraft observations during TC4, Pre-AVE, and CR-AVE (see Aschmann et al., 2009, 

Hossaini et al., 2010 and Liang et al., 2010) and balloon observations from Teresina, Brazil, in June 2005 (Laube et al.2008) and June 

2008 (update of Laube et al., 2008). d LZRH(z0) corresponds to the level of zero clear-sky radiative heating (z0) (see Box 1-4). As in the previous assessment (Law and 

Sturges et al., 2007), we define this level at about 15 km or 360K, where there is a transition from clear-sky radiative cooling to clearsky radiative heating. In general, air masses above this level are expected to enter the stratosphere. e The stated observed range represents the smallest mean minus 1 standard deviation and the largest mean plus 1 standard deviation 
1.43

ODSs and Related Chemicals

significant macroalgae on the local shore. Peak concentrations were observed at solar noon (2–44 ppt CHBr3 and 

1–9 ppt CH2Br2), rather than low tide as would be expected for emissions from intertidal macroalgae.

In a case study off the West coast of Africa, Quack 

et al. (2007) suggest that CH2Br2 might originate partly from biologically mediated reductive hydrogenolysis of CHBr3. Some studies (e.g., Palmer and Reason, 

2009; Kerkweg et al., 2008) suggest significant regional 

“hotspots” in the tropics and subtropics especially in the 

west Pacific Ocean near Indonesia, but with large seasonal and temporal variations. This point is also in accord with Butler et al. (2007), who estimated that 64% 

of bromoform emissions originated from the Pacific 

(47% from the tropical Pacific, and 70% in total from the 

tropics), compared with just 3% from the Atlantic. For 

CH2Br2 the corresponding figures were 40% (24%, 40%) 

and 20%, respectively. In the last Assessment (Law and 

Sturges et al., 2007) it was proposed that regions of oceanic upwelling could be important sources of bromocarbon emissions. Since then Quack et al. (2007) and Carpenter et al. (2009) have shown, based on measurements 

in the Mauritanian upwelling, that such regions may be 

collectively relatively minor contributors to global emissions (e.g., around 1% of CHBr3 and CH2Br2 global 

emissions).

It should also be noted that discrepancies between 

observation-based studies arise from the “snapshot” picture they supply of a highly heterogeneous and variable 

system, as well as possible substantial differences between 

calibration scales (Butler et al., 2010) and methods of estimating fluxes and extrapolating them to larger areas (e.g., 

Carpenter et al., 2009 and O’Brien et al., 2009). The picture of global bromocarbon emissions is therefore highly 

complex with substantial uncertainties. Understanding the 

regional and seasonal variability of emissions is, however, 

among all measurement campaigns. f CH2ClCH2Cl (dichloroethane) and CHClCCl2 (trichloroethene) in the UT only from TC4. g CH2ClCH2Cl and CHClCCl2 in the TTL, including the LZRH, only from TC4, Pre-AVE, and CRAVE. h CH2BrCl in the UT only from PEM-West B. i CH2BrCl and CHBrCl2 in the TTL, including the LZRH, only from TC4 and Teresina balloon measurements.

j Estimated maximum contribution from species like C2H5Br, C2H4Br2, C3H7Br. k CH3I in the TTL, including the LZRH, only from TC4 and Pre-AVE. l The anthropogenic fraction of chlorinated VSLS (Anthrop. Cl) has been calculated by adding 90% of CH2Cl2, 25% of CHCl3, and 

100% of CCl2CCl2, CHClCCl2, and CH2ClCH2Cl (see also Section 1.3.1.1).

Table 1-8. Fluxes of bromine from bromoform (CHBr3) and dibromomethane (CH2Br2) in Gg Br/yr, and 

iodine from methyl iodide (CH3I) in Gg I/yr.

Reference

CHBr3 Flux (Gg Br/yr) CH2Br2 Flux (Gg Br/yr) CH3I Flux (Gg I/yr)

Global Open 

Ocean Coastal Global Open Ocean Coastal Global Open Ocean 

Coastal 

Butler et al. (2007) 800 150 650 280 50 230 550 270 280

Carpenter et al. (2009) 200 

Liang et al. (2010) e 430 260 170 57 34 23

O’Brien et al. (2009) 820 a

1400 b

 

Palmer and Reason (2009) 120 c 

Yokouchi et al. (2005) 820 a 

Warwick et al. (2006) 560 d 280 d 280 d 100 

a Scaled to CH2Br2 emissions from Ko and Poulet et al. (2003), based on global loss rates and an estimated global burden. b Scaled to CH2Br2 emissions from Warwick et al. (2006). c Tropical ocean only. d Modeling study: “Scenario 5”: 70% of emissions in the tropics; August/September. e Emissions from Liang et al. (2010) are from modeling of airborne measurements in the Pacific and North American troposphere and lower stratosphere.

Table 1-7, continued (notes).
1.44

Chapter 1

vital in assessing their likely impact on the stratosphere, 

since it is the coincidence in both space and time of VSL 

SG emissions with surface to tropical tropopause layer 

(TTL) transport that will largely dictate the efficiency with 

which these emissions reach the stratosphere.

Updated mixing ratios of organic bromine from 

VSLS of 3.5 (1.7–7.4) ppt in the tropical upper troposphere 

(10–12 km) (Table 1-7) confirm reported figures from the 

last Assessment of 3.5 (3.1–4.0) ppt at 10 km. This organic bromine from VSLS consists mostly of CH2Br2 and 

CHBr3 with smaller amounts of bromochloromethanes 

and is discussed further in Section 1.3.3.1.

1.3.1.3 Iodine-Containing Very

Short-Lived Source Gases

The occurrence and chemistry of iodinated compounds, all of which are VSLS, were discussed extensively in previous Assessments due to their potentially 

very high ODPs should they reach the stratosphere (the 

“alpha” factor for iodine in the stratosphere is higher than 

bromine). Iodinated gases are predominantly of natural 

origin, but have been considered for use as fire suppressants as a halon replacement on aircraft (e.g., CF3I) or to 

reduce flammability in refrigerant mixtures. CH3I is also 

currently being licensed in the U.S. and Japan as a replacement for CH3Br as a fumigant (UNEP/TEAP, 2010). Domestic biomass burning in Africa reportedly contributes a 

small flux of CH3I to the atmosphere (Mead et al., 2008c).

The dominant source of iodinated VSLS is from the 

ocean. Butler et al. (2007) estimated approximately equal 

global fluxes of CH3I from the open ocean and coastal regions (Table 1-8) and a higher total flux than given in the 

previous Assessment by about a factor of two. Yokouchi 

et al. (2008) observed an increase in CH3I with sea surface 

temperature at midlatitudes, which might be of significance 

in a warming climate, and a corresponding seasonality 

with higher concentrations in summer. An intriguing 

observation is that of Williams et al. (2007), who presented 

evidence for enhanced emissions of CH3I from the ocean, 

apparently from an abiotic chemical mechanism stimulated 

by dust deposition. A corollary of this is that a similar 

mechanism might operate in atmospheric aerosols. Sive 

et al. (2007) conclude that terrestrial biomes contribute a 

small but significant flux to the atmosphere.

Several recent studies have measured a range of 

iodinated species in the marine boundary layer (MBL) 

and considered the active iodine-mediated chemistry 

therein. It is unlikely, however, that either the shortestlived iodinated species (those with lifetimes of minutes or 

hours), or inorganic iodinated gases (e.g., I2, IO, and OIO; 

Read et al., 2008), would escape the MBL in appreciable 

quantities (see also Section 1.4.3).

1.3.1.4 Halogen-Containing Aerosols

In addition to organic VSLS, inorganic sources, notably halogen-containing aerosols, may be precursors of 

reactive halogens in the atmosphere. Although aerosols 

are subject to gravitational and wet depositional processes, 

they can have significant atmospheric residence times under the dry conditions prevalent in the free troposphere 

(e.g., several days for sea salt aerosol; Yang et al., 2005). 

Yang et al. (2005) suggested that 0.1–1.0 ppt of bromine 

monoxide (BrO) in the troposphere arises from a combination of Br release from sea salt and decomposition of bromomethanes. Support for such a mechanism comes from 

the model described by O’Brien et al. (2009) where additional release of halogens from halide-containing aerosol 

(over and above that from degradation of bromocarbons) 

was required to explain BrO abundances measured within 

the tropical marine boundary layer (Read et al., 2008). 

Pyle et al. (2007) note that sea salt production rates are 

highly uncertain and that using two separate formulations 

results in a factor of two difference in sea salt aerosol production rates. Bromine and iodine have been detected in 

aerosols of tropospheric origin in the stratosphere several 

kilometers above the local tropopause height (Murphy et 

al., 2007). The influence of these aerosols on lower stratospheric halogen abundances is likely small but is poorly 

quantified. We conclude that the contribution from sea 

salt aerosol to stratospheric Bry is likely to be substantially 

less than 1 ppt.

1.3.2 Transport of Very Short-Lived 

Substances into the Stratosphere

The efficiency of VSLS transport into the stratosphere results from the competition between fast vertical 

transport and chemical destruction or removal via washout of the source gases (SGs) and product gases (PGs). 

Conversion of SG to PG reduces the efficiency of halogen transport to the stratosphere if the PG is taken up by 

aerosols and becomes washed out before reaching the 

stratosphere. The probability of washout decreases with 

increasing height within the troposphere and becomes 

small in the TTL. In general, SG and PG reach the “tropical stratosphere” above the cold point tropopause (CPT), 

which is climatologically located at 380 K (~17 km altitude). Therefore, timescales for transport through the TTL 

are especially important for determining the efficiency of 

VSLS transport through this layer. For a detailed TTL 

schematic see Figure 1-9 (adapted from Fueglistaler et al. 

(2009a) and modified to provide definitions of the TTL 

consistent with past Assessments).
1.45

ODSs and Related Chemicals

1.3.2.1 VSLS Transport from the Surface

in the Tropics to the Tropical 

Tropopause Layer (TTL)

While shorter-lived VSLS with lifetimes of only 

hours to days are unlikely to reach the TTL unless emitted directly into the active cell of tropical deep convection, VSLS with lifetimes of weeks to months can reach 

the TTL via a range of transport pathways described in 

Figure 1-9.

Additional evidence has become available since the 

previous Assessment supporting the tropical Western Pacific as a region where VSLS emissions can be efficiently 

transported to the TTL (Aschmann et al., 2009) (Figure 

1-10). New results also suggest that the efficiency of this 

TTL transport varies with season. Higher VSLS concentrations in the TTL are predicted in NH winter compared 

to the rest of the year (Aschmann et al., 2009; Gettelman et 

al., 2009) due to a combination of higher convective cloud 

tops reaching the TTL and higher vertical velocities within 

the TTL (Gettelman et al., 2009).

In NH summer, air masses can cross the TTL and 

enter the stratosphere, but with a reduced and more zonally uniform distributed flux compared to NH winter (FuegFigure 1-9. Schematic showing principal dynamical pathways in the tropics transporting VSL source gases 

(SG) and product gases (PGs) into the stratosphere (modified from Fueglistaler, 2009a). The position of the 

tropopause is designated by the bold black line.

Ci: cirrus cloud

CPT: cold point tropopause

FT: free troposphere

LMS: lowermost stratosphere

BL: boundary layer

MCO: maximum convective outflow

TTL: tropical tropopause layer

z0: level of clear-sky zero radiative heating

Relative Contribution

Altitude [km]

Figure 1-10. The relative contribution (%) of individual source regions 

to the total amount of an idealized 

CHBr3 tracer (with a 20-day lifetime) 

between 8 and 20 km in the tropics, 

averaged over 2000–2005. Australia 

includes Maritime Continent and Africa includes Arabian Peninsula. The 

relative contribution from the Mediterranean region was also diagnosed 

but is too small to be visible in this 

figure (modified from Aschmann et 

al., 2009).
1.46

Chapter 1

listaler et al., 2005). However, the northward shift of the 

deepest convection with the Indian summer monsoon 

linked with the anticyclonic circulation in the upper troposphere/lower stratosphere could lead to more efficient 

transport of surface emissions in certain regions of the 

world (e.g., Bay of Bengal and Sea of China) to the TTL 

and potentially to the stratospheric overworld (e.g., James 

et al., 2008; Randel et al., 2010). For instance, Stohl et 

al. (2002) have estimated that in NH summer, Asian surface emissions can be transported into the TTL with average timescales of a week or less instead of three weeks in 

NH winter. Consistent with these findings, Donner et al. 

(2007) simulate enhanced methyl iodide (CH3I) concentrations over the Tibetan plateau and the Himalaya during 

NH summer in the free troposphere and above the coldpoint tropopause due to overshooting convection. Thus 

tropical overshooting convection can directly inject surface air into the upper TTL and to the stratosphere as is observed (e.g., Ricaud et al., 2007) and simulated by different models (Donner et al., 2007; James et al., 2008). Due 

to the very rare and localized occurrence of such events 

(Liu and Zipser, 2005; Liu et al., 2007), the amount of 

VSLS halogen reaching the stratosphere via these events 

is expected to be small on a global basis, as has been discussed analogously for the stratospheric water vapor budget (Fueglistaler et al., 2009b; Schiller et al., 2009; Homan 

et al., 2010). Besides, Hossaini et al. (2010) suggest that 

especially for longer-lived VSLS (e.g., CH2Br2), largescale transport and mixing of planetary boundary layer 

(PBL) air to the TTL could also be important transport 

processes in addition to overshooting convection.

Because observational evidence for direct VSLS 

transport from the surface to the TTL and to the stratosphere is scarce, the relative roles of boundary layer mixing, convection, and overshooting convection are still a 

matter of debate. In addition to the uncertainties in convective parameterizations in models discussed in the last 

Assessment (Law and Sturges et al., 2007), washout of 

VSLS influences the amount of VSLS actually reaching 

the stratosphere (Section 1.3.3).

1.3.2.2 VSLS Transport from the TTL

to the Stratosphere

Transport through the upper TTL primarily takes 

place as large-scale horizontal and slow vertical motions (weeks to ascend through this region of the TTL) 

accompanied by infrequent localized and rapid (hours) 

overshooting convection into the stratosphere (Law and 

Sturges et al., 2007; Fueglistaler et al., 2009a; Figure 1-9). 

Furthermore, the transport of VSLS through the TTL is 

also dependent on the region and season of VSLS emissions and varies interannually (Aschmann et al., 2009; 

Gettelman et al., 2009; Krüger et al., 2009).

In the last Assessment a large uncertainty was reported for the residence time in the TTL, with a range of 

20 to 80 days in the upper TTL. The understanding of 

residence times in the TTL has since improved to some extent. In particular, the residence time depends strongly on 

vertical velocities, on the particular layer being considered 

within the TTL, and on the temporal and horizontal variability used for calculating the transport. New approaches 

and improved data assimilation have become available. 

This has led to a reduced uncertainty range of 25 to 45 

days for transport through the upper TTL (Krüger et al., 

2009; Ploeger et al., 2010).

Of particular relevance for the VSLS fraction 

reaching the stratosphere is the slow, radiatively driven 

transport through the upper TTL (from the level of clearsky zero radiative heating (z0) at 360 K to the CPT at 380 

K) (see Figure 1-9). The annual cycle of air mass transport 

from the TTL to the stratosphere peaks over the tropical 

Western Pacific during NH winter and over the Indian 

monsoon regions during NH summer (Fueglistaler et al., 

2005; Law and Sturges et al., 2007). In addition, the TTL 

residence time close to the CPT is lower by 25% in NH 

winter compared to NH summer (Kremser et al., 2009; 

Ploeger et al., 2009) and this also influences the efficiency 

of VSLS transport to the stratosphere.

At the CPT and above, large-scale vertical motion dominates the air mass transport (Krüger et al., 2008; 

Fueglistaler et al., 2009a; Gettelman et al., 2009) and readily brings air to regions of the stratosphere where ozone 

is depleted. Above the tropical Western Pacific, Krüger 

et al. (2009) found that most air parcels are transported in 

less than 25 days through the 360 K to 380 K layers during NH winters 1962–2001, which is at the lower end of 

the averaged residence time of 25–45 days stated above. 

They derived a long-term interannual variability of the 

averaged residence time for air parcels in the upper TTL 

during NH winter of up to 20% (±2 sigma); with shorter 

residence times being significantly correlated with periods 

of enhanced extratropical and subtropical wave driving. 

General changes in tropical upwelling rates and of the 

Brewer-Dobson circulation (BDC) are discussed in more 

detail in later Chapters (Sections 3.2.4 and 4.2.2). These 

considerations lead to a higher potential of halogens from 

VSLS entering the stratosphere within these regions and 

seasons during periods of enhanced wave driving and/or 

convective overshoot.

1.3.2.3 VSLS Transport from the Surface

to the Extratropical Stratosphere

The extratropical lowermost stratosphere (LMS) is 

the region between the extratropical tropopause and approximately the 380 K isentropic surface (see Figure 1-9). While 

the contribution of VSLS to halogen in the extratropical 
1.47

ODSs and Related Chemicals

stratosphere is mainly controlled by upwelling into the tropical stratosphere and quasi-horizontal isentropic transport 

into the extratropics, it has been perceived in recent years 

that vertical transport by convection into the extratropical 

stratosphere can also be of importance (Section 2.4.2 of 

Law and Sturges et al., 2007). Furthermore, new evidence 

suggests that a larger fraction of tropical air masses in the 

TTL are transported into the extratropical stratosphere than 

into the stratospheric overworld, as detailed below.

Levine et al. (2008) simulate that up to 5 times more 

air mass is transported from the bottom of the TTL into the 

LMS than from the bottom of the TTL into the tropical 

stratospheric overworld throughout the year. However, 

the amount of transport into the extratropical LMS is largest toward the SH and shows, for both transport pathways, 

a seasonal dependence with a maximum during NH winter 

and a minimum during NH summer.

In contrast, the studies by Bönisch et al. (2009) 

and Berthet et al. (2007), which concentrated on the NH 

only, analyzed a stronger transport from the tropical troposphere to the extratropical LMS during NH summer. 

Bönisch et al. (2009) estimate a strong seasonality of the 

origin of tropical air masses in the LMS of the NH, with 

a maximum tropospheric influence during October. The 

authors conclude that the seasonality is influenced by the 

transit times from the tropical troposphere into the LMS, 

which are shortest (<0.3 years) during summer and longest (>0.8 years) during late spring. From this it can be 

concluded that the strength of the subtropical jet and the 

lower branch of the BDC strongly influence the transport 

of VSLS in the LMS (i.e., the subtropical jet is weaker 

from summer to mid-autumn, and the residual transport 

timescales within the lower branch of the BDC are shorter 

in summer than in winter and dominate the tropospheric 

influence in the LMS). Berthet et al. (2007) also find that 

during NH summer, the extratropical LMS layer between 

370 K and 410 K is well ventilated with tropospheric 

air from the tropics compared to the layer below 340 K, 

which is mainly influenced by extratropical air. During 

NH winter this distinction holds but with weaker tropical troposphere-to-stratosphere transport (TST) due to a 

more impermeable barrier.

Convective overshooting of air across the tropopause into the LMS is quite frequent in the extratropics. 

These injections, together with transport from the TTL, 

form a mixed layer extending about 25 K above the tropopause (Hoor et al., 2004). Sometimes, injections even 

deeper into the stratosphere can occur and influence the 

chemical composition of the lower stratosphere (Hegglin 

et al., 2004). The deepest injections are produced by the 

high instability and breaking of gravity waves excited by 

strong convection inside super-cell thunderstorms (Wang, 

2007; Wang et al., 2009). While water vapor enrichments 

have been documented by this mechanism, no studies 

related to VSLS transport via this mechanism have been 

published. Another possibility for surface air to be injected directly into the extratropical stratosphere (even into 

the stratospheric overworld) is by pyro-convection associated with strong forest fires. Single pyro-convection 

events can double the zonally averaged aerosol optical 

depth in the LMS and these aerosol layers can persist in 

the stratosphere for months (Fromm et al., 2008). While 

this process is not captured by global models, special highresolution models have recently shown skill in simulating 

pyro-convection (Trentmann et al., 2006; Luderer et al., 

2007). So far, no observations of halogenated VSLS in 

pyro-convective outflows have been reported.

1.3.3 VSLS and Inorganic Halogen Input 

to the Stratosphere

In the previous Assessment (Law and Sturges et 

al., 2007) the contribution of VSLS species to the stratospheric halogen loading was discussed in detail. Since 

this last Assessment new observations of VSLS SGs in the 

tropical upper troposphere and TTL have become available, as well as modeling studies of the possible input of 

product gases (PGs) from these VSLS. However, for brominated PGs still no observations are available from the 

upper tropical troposphere and from the TTL that would 

enable us to identify the contribution of PGs to stratospheric bromine loading.

1.3.3.1 Source Gas Injection (SGI)

Source gases deliver halogen into the stratosphere 

in the same form as they are emitted at the surface. Some 

recent modeling studies (Donner et al., 2007; Aschmann 

et al., 2009; Liang et al., 2010; Hossaini et al., 2010) have 

used observational data of the most important bromineand iodine-containing VSLS species (CH2Br2, CHBr3, 

and CH3I) to compare with their modeling results. These 

new observational data on VSLS species in the TTL have 

become available from the Aura Validation Experiment 

(AVE; http://espoarchive.nasa.gov/archive/arcs/pre_ave/ 

and http://www.espo.nasa.gov/ave-costarica2/) and TC4 

(Tropical Composition, Cloud and Climate Coupling, 

http://www.espo.nasa.gov/tc4/) missions, and from balloonborne observations (Laube et al., 2008). In addition, 

available data from the Pacific Exploratory Mission (PEM) 

West A and B campaigns have been used. New balloonborne observations of chlorinated VSLS and of additional 

brominated VSLS have also been reported by Laube et 

al. (2008). The averages and the ranges of these observations are compiled in Table 1-7 for the upper tropical 

troposphere (10 km altitude range), the bottom of the TTL 

(340–355 K), z0 (taken as about 355–365 K), the upper 
1.48

Chapter 1

TTL (365–375 K), and the tropical cold point tropopause 

(375–385 K) (Figure 1-9). In the following discussion of 

uncertainties, we do not take into account uncertainties in 

absolute calibrations (e.g., Butler et al., 2010). Rather we 

use data from different sources and take the ranges of the 

reported mixing ratios. Uncertainties in the absolute calibrations will thus be reflected in the range of uncertainties 

due to the use of data from different sources. Furthermore, 

the data coverage is still sparse (e.g., not all possible input 

areas and seasons have been sampled) and it is unclear 

how far the available observations reflect the complete 

range of possible input values.

SGI from Chlorinated VSLS

Law and Sturges et al. (2007) concluded on the 

basis of multiple airborne campaigns that the chlorine content from VSL SGs in the tropical upper troposphere was 

about 50 ppt, representing about 1 to 2% of the inorganic 

stratospheric chlorine (Cly) input from long-lived source 

gases. Recent airborne and balloonborne observations in 

the tropics have confirmed and refined these conclusions 

(Laube et al., 2008) (see Table 1-7). The major loss process 

for most chlorine-containing VSLS is the reaction with 

OH. As the rate constants for these reactions decrease at 

the low temperatures in the upper tropical troposphere and 

TTL, vertical mixing ratio gradients are rather small. The 

total chlorine contained in VSLS source gases decreases 

from about 60 ppt in the upper tropical troposphere and 

lower TTL to about 50 ppt near the tropical tropopause. 

C2H5Cl was not quantified, but was estimated to be on the 

order of 1.5 ppt in the upper tropical troposphere (Law and 

Sturges et al., 2007).

Our best estimate of chlorine injected as SG from 

VSLS into the stratosphere is 55 (38–80) ppt and largely 

confirms the estimate of 50 (±10) ppt given in the previous 

Assessment (Law and Sturges et al., 2007).

SGI from Brominated VSLS

The two main brominated VSLS (CHBr3 and 

CH2Br2) show different vertical gradients in the transition 

region between the troposphere and the stratosphere (see 

Table 1-7). Whereas mixing ratios of CH2Br2 decrease 

from about 0.9 ppt in the upper tropical troposphere and the 

lower TTL to about 0.5 ppt near the tropical tropopause, 

mixing ratios of CHBr3 decrease more substantially in this 

region, declining from 0.5 ppt in the upper troposphere and 

lower TTL to 0.1 ppt near the tropical tropopause.

These observed differences are explained by the 

different local lifetimes of these gases within the TTL. 

While photolysis dominates the destruction of CHBr3 and 

leads to a modeled local lifetime that is similar to transport 

times for air through the TTL, the lifetime of CH2Br2 (set 

mainly by [OH]) is much longer than transport timescales 

(Figure 1-11).

Lifetime differences are also noted in the troposphere for these gases, and this influences the efficiency 

of surface emissions reaching the upper troposphere. The 

tropospheric local lifetime of 23 days for CHBr3 is mainly 

determined by photolysis (36-day partial lifetime) and to 

a lesser degree by its reaction with the OH radical (76-day 

Figure 1-11. The modeled 

vertical distribution of the annual mean lifetime (days) of 

CHBr3 and CH2Br2 in the tropics (±20° latitude) with respect 

to the two main loss processes: photolysis (solid red line) 

and reaction with OH (solid 

blue line), as well as the overall lifetime (dashed black line). 

The region between the base 

of the TTL and the cold-point 

tropopause (CPT) is indicated 

as the gray bar. Modified from 

Hossaini et al. (2010).
1.49

ODSs and Related Chemicals

partial lifetime). In contrast, the considerably longer tropospheric local lifetime of CH2Br2 of 123 days is nearly 

completely governed by its reaction with the OH radical 

(Table 1-4 and Figure 1-11).

As a result of these influences, a 1-D model calculation suggests that 70% and 28% of the specified surface 

concentrations of CH2Br2 and CHBr3, respectively, reach 

18 km altitude (top of the TTL), when operational European 

Centre for Medium-Range Weather Forecasts (ECMWF) 

velocity fields, a convective source, and a VSLS sink are 

taken into account (Gettelman et al., 2009). Hossaini et al. 

(2010) also discuss the breakdown of CH2Br2 and CHBr3

in the atmosphere and compare their three-dimensional 

(3-D) model results with observational data (Figure 1-12). 

In the upper tropical troposphere and the TTL, Hossaini et 

al. (2010) achieved an overall better agreement of simulated and observed CH2Br2 and CHBr3 profiles when using 

the slower (climatological) diabatic heating rates in the 

Single Layer Isentropic Model of Chemistry and Transport (SLIMCAT) model compared to using the faster analyzed vertical velocities in the Toulouse Off-line Model 

of Chemistry and Transport (TOMCAT)-chemical transport model. Assuming uniform mixing ratios of 1.2 ppt 

of CH2Br2 and CHBr3 in the tropical lower troposphere, 

they derive a SGI of about 2 ppt Br into the stratosphere. 

Additionally, the sum of CH2BrCl, CHBr2Cl, and CHBrCl2

contributes about 0.5 ppt of bromine to the upper troposphere and lower TTL. Near the tropical tropopause the 

total contribution of these source gases declines to about 

0.2 ppt. Additional bromine from species not quantified, 

like C2H5Br and n-propyl bromide (C3H7Br), is expected to 

be presently less than 0.2 ppt in the TTL. In summary, we 

estimate VSLS bromine source gases, including unmeasured species, to account for about 2.7 (1.4–4.6) ppt of bromine at about 15 km altitude in the tropics and about 1.5 

(0.7–3.4) ppt at the tropical CPT at around 17 km altitude.

SGI from Iodinated VSLS

Newly available data for CH3I in the TTL from 

the Pre-AVE and TC4 experiments (see Table 1-7) (Aschmann et al., 2009) as well as data from Bell et al. (2002), 

which are compared to modeling results in Donner et al. 

(2007), indicate that the mixing ratios of this most important iodine source gas in the TTL at 15 km altitude are 

generally below 0.1 ppt and drop below 0.05 ppt at the 

cold point tropopause. This is consistent with the estimate 

of 0.08 ppt from the previous Assessment for CH3I in the 

upper tropical troposphere (Law and Sturges et al., 2007). 

Three recent studies have also investigated the transport of 

methyl iodide into the TTL (Donner et al., 2007; Gettelman et al., 2009; Aschmann et al., 2009). These studies 

confirm that due to its short lifetime, only a very small 

fraction of CH3I from the boundary layer reaches the TTL 

and the stratosphere.

1.3.3.2 Product Gas Injection (PGI)

In addition to the input of halogen from VSLS 

source gases, their chemical degradation products (socalled product gases or PGs) are also likely to contribute 

halogens to the stratosphere. Due to their solubility in water 

Figure 1-12. A comparison 

between observed and modeled vertical distributions of 

CHBr3 (left panel) and CH2Br2

(right panel) mixing ratios (ppt) 

at tropical latitudes (Hossaini et 

al., 2010). The red and black 

lines are from TOMCAT model 

runs using ECMWF vertical 

velocities, without and with parameterized convection, respectively, while the SLIMCAT 

model run (blue lines) is based 

on calculated diabatic heating 

rates. Black dots are measurements from the CR-AVE campaign. Modified from Hossaini 

et al. (2010).
1.50

Chapter 1

and the significant uptake coefficients on ice of many of 

these PGs (see Law and Sturges et al., 2007, for a detailed 

discussion) their fate is less clear than that of source gases 

and depends critically on the physical and chemical conditions in the area where they are formed. This leads to 

a strong coupling between chemistry, removal processes, 

and atmospheric transport processes relevant for VSLS 

species. While observations of source gases discussed in 

Section 1.3.3.1 have become available, observations of 

PGs are sparse for chlorine and no observations of PGs 

above detection limits of the respective instruments have 

been reported for bromine and iodine in the upper tropical 

troposphere and the TTL. An assessment of plausible values of PGI for these two halogens therefore relies largely 

on modeling results that are not currently constrained by 

observations.

PGI from Chlorinated VSLS

In the previous Assessment there were large uncertainties regarding the contribution of chlorinated 

PGs, notably COCl2 (phosgene) and HCl, arising from 

degradation of chlorinated VSLS (and longer-lived 

SGs) to Cly. The picture is complicated due to possible 

“double counting” of chlorinated PGs recirculated from 

the stratosphere and thus already taken into account in the 

budget of source gases entering the stratosphere. Fu et al. 

(2007) reported satellite measurements of COCl2, albeit 

with large uncertainty limits, between February 2004 and 

May 2006. Chlorine from COCl2 between 15 and 20°N 

was 43 ± 27 ppt Cl at 8.5 km (upper troposphere), 31 ± 22

ppt Cl at 14.5 km, and 36 ± 26 ppt Cl at 17.5 km, i.e., approximately 32 ± 22 ppt of chlorine at z0 (~15 km). These 

values are at the lower end of the range (40 – 50 ppt Cl)

estimated in the previous Assessment (Law and Sturges 

et al., 2007). As it is unclear how much of the chlorine in 

COCl2 is already included in the budgets from the SGs, we 

follow the same approach as in the previous Assessment 

by considering a range of values, or 0–32 ppt of additional 

chlorine. A best estimate of 16 ppt results as possible additional chlorine input to the stratosphere in the form of 

COCl2, which is not accounted for in the very short-lived 

source gas budgets.

Marcy et al. (2007) reported HCl to be below their 

detection limit of 5 ppt in upper tropospheric air over the 

tropical east Pacific. In the lower TTL over the east Pacific 

in January 2004, however, Marcy et al. (2007) reported HCl 

to be mostly non-zero with values ranging to almost 40 ppt 

around 15 km altitude, and increasing to about 20–80 ppt 

in the upper TTL, with the increase being mainly attributed to in-mixing of lowermost extratropical stratospheric 

air. Mébarki et al. (2010), using a less sensitive long-path 

instrument, placed an upper limit of 30 and 20 ppt HCl in 

the upper TTL (~15–17 km) from two flights over Brazil 

in June 2005 and June 2008, respectively. Marcy et al. 

(2007) estimated that at most only 3–8 ppt of the HCl that 

they observed in the TTL could have originated from in

situ degradation of CH3Cl. Park et al. (2008) report that 

profiles of HCl concentration reached minimum values of 

about 10 to 30 ppt at around 13 km above the Asian continent, consistent with the measurements reported by Marcy 

et al. (2007). In summary, we conclude that HCl values 

in the TTL are typically on the order of 20 (0–40) ppt. 

As it is unclear how much of this chlorine will eventually 

reach the stratosphere or has already been accounted for 

by source gases, we estimate that somewhere between 0 

and 100% of this HCl (i.e., 0–20 ppt of chlorine) could 

represent an additional source of chlorine from VSLS to 

the stratosphere, with a best estimate value being the midpoint (10 ppt).

PGI from Brominated VSLS

The only bromine PG that has been measured in the 

TTL is bromine monoxide (BrO) (Dorf et al., 2008; see 

also Section 1.4). The observed values were reported to be 

below detection (i.e., below 1 ppt BrO) in the TTL, though 

a value of 2 ± 1.5 ppt BrO was derived around the local tropopause at about 17 km. Because this sparse observational 

information does not tightly constrain the average input of 

brominated PG into the stratosphere, a range of plausible 

values is derived here based on modeling studies.

Most modeling studies rely on the assumption that 

bromine from photo-oxidized VSLS is immediately transformed to inorganic bromine (mainly as hydrogen bromide 

(HBr), hypobromous acid (HOBr), and BrO). This is supported by Hossaini et al. (2010), who concluded that all 

intermediate organic products of the oxidation of bromocarbons are short lived, with COBr2 being the longest-lived 

organic intermediate with an atmospheric lifetime of 7 days.

The assumption that the degradation of organic 

bromine-containing VSLS yields inorganic bromine is 

thus fairly well established. The subsequent fate, however, of this inorganic bromine (Bry) is less clear. Most 

inorganic bromine gases, especially HOBr and HBr, are 

highly soluble in water and have high uptake coefficients 

on ice (see Section 2.3.4 of Law and Sturges et al., 2007). 

The removal of Bry from the gas phase and the subsequent 

washout is parameterized in models. This parameterization has a strong influence on the fraction of product gases 

reaching the stratosphere. Furthermore, as already emphasized in the last Assessment, heterogeneous reactions of 

HOBr and HBr with sulfuric acid producing photolabile 

dibromine monoxide (Br2O) and molecular bromine (Br2) 

further complicate the modeling of this process.

Model studies presented in the previous Assessment used prescribed removal timescales for Bry in the 

TTL (e.g., Warwick et al., 2006; Sinnhuber and Folkins, 
1.51

ODSs and Related Chemicals

2006), i.e., the removal processes (wet deposition, uptake 

on ice, and dehydration) have not been calculated explicitly. Hossaini et al. (2010) follow a similar approach and 

test the sensitivity of the amount of bromine reaching the 

stratosphere on different washout times for Bry, ranging from 10 days to infinity. Depending on the assumed 

time constants for washout times, the PGI bromine 

from CH2Br2 and CHBr3 varies between 0.4 ppt (10 days 

washout) and 1.7 ppt (40 days washout time) and can be 

as high as 3.9 ppt of bromine if an infinite washout time 

is assumed. This range of values is consistent with the 

ranges presented in the previous Assessment (Law and 

Sturges et al., 2007). Two other recent studies (Aschmann 

et al., 2009; Liang et al., 2010) have attempted to explicitly calculate the removal processes, based on dehydration 

processes in the model. Aschmann et al. (2009) used an 

idealized tracer that has properties similar to CHBr3. The 

inorganic bromine is either completely removed if dehydration occurs in a grid box or it is not removed at all. 

They derive a bromine input from this tracer that is between 1.6 and 3 ppt, of which 0.3 ppt is in the form of the 

SGs. The range in PGI from bromoform is thus between 

1.3 and 2.7 ppt. Therefore, even with the assumption of 

complete solubility and instantaneous washout, only about 

50% of the PGs are removed in the TTL in this study. 

Using the Goddard Earth Observing System (GEOS) 

chemistry-climate model (version 2) with a detailed wet 

deposition scheme, Liang et al. (2010) suggest that washout is mainly effective below 500 hectoPascals (hPa) and 

that virtually all inorganic bromine that is released from 

the SGs above that level reaches the stratosphere.

These large differences in the effect of washout 

predicted by models show the high degree of uncertainty 

in the physical understanding of the removal of inorganic 

bromine during the transport into the tropical stratosphere. 

Due to this high level of uncertainty and due to the lack of 

observations of bromine PG in the TTL, it is not possible 

to derive a best estimate of the amount of bromine entering 

the stratosphere in the form of PG. Based on the studies 

presented above and in previous Assessments, the range of 

PGI from CHBr3 and CH2Br2 is estimated to be between 

0.4 and 3.9 ppt. Based on the vertical profiles of other 

minor bromine source gases, an additional product gases 

injection from these could range between 0 and 0.3 ppt. 

Total PGI could thus be somewhere in the range between 

0.4 and 4.2 ppt Br.

PGI from Iodinated VSLS

Consistent with the observations presented in the 

previous Assessment (Law and Sturges et al., 2007), balloonborne measurements using the Differential Optical 

Absorption Spectroscopy (DOAS) technique have revealed no measurable amount of iodine monoxide (IO) or 

iodine dioxide (OIO) in the TTL above detection limits of 

0.1 ppt (Butz et al., 2009).

1.3.3.3 Total Halogen Input into the

Stratosphere from VSLS and Their

Degradation Products

As discussed above, both PGI and SGI can contribute halogen to the stratosphere. Here we consider only 

amounts transported into the stratosphere above the CPT 

as estimated from the mixing ratios of SGs and PGs at 

that level and at 15 km (z0). Since the previous Assessment, additional observations of bromine VSL SGs in the 

TTL have become available. The halogen contribution 

from PGs, however, remains much more uncertain. We 

summarize the available estimates of SGI and PGI for the 

individual halogens (chlorine, bromine, iodine) in Table 

1-9. PG observations are available for chlorine but not 

for bromine. Therefore, a best estimate of total chlorine 

input to the stratosphere is derived, based on SG values 

observed at z0 and typical values of chlorine PG observed 

in the TTL. In contrast only a range of possible values is 

derived for bromine, based on observations at the CPT and 

modeled PGI. Only upper limits are available for iodine 

from observations.

Total Input from Chlorinated VSLS

Adding up the entire range of SG and PG observations discussed in Sections 1.3.3.1 and 1.3.3.2 yields 

a range between 25 and 170 ppt of chlorine from VSLS. 

For a best estimate of chlorine input into the stratosphere 

from this range, we use SG observations in the middle of 

the TTL (at z0), where VSLS chlorine is 55 (38–80) ppt. 

It is difficult to assess if chlorine from the PGs HCl and 

COCl2 represents an additional source or is already taken 

into account in the source gas budgets. The best estimate 

is 16 (0–32) ppt of chlorine from COCl2 and 10 (0–20) 

ppt of HCl (see Section 1.3.3.2). The best estimate values 

for total VSLS chlorine is thus 80 (40–130) ppt, largely in 

agreement with the estimate given in the previous Assessment of about 50 ppt of chlorine from SGs and about 40 

to 50 ppt chlorine from PGs (taken then to include COCl2

only).

Total Input from Brominated VSLS

In order to avoid double counting, two approaches 

are used to derive total bromine input from VSLS: (i) the 

range of possible values of VSLS bromine entering the 

stratosphere is derived from the range of SG observations 

at the CPT and the range of modeled PG injections, and 

(ii) a probable input is derived based on SG observations 

at z0, and assuming that PGs produced above that level 
1.52

Chapter 1

are transported into the stratosphere, while PGs produced 

lower in the TTL and upper tropical troposphere have been 

washed out. Following the first approach, we estimate that 

a lower limit of 1.1 (0.7 SG and 0.4 PG) ppt of bromine 

from VSLS is transported into the stratosphere, while our 

upper limit estimate is 7.6 (3.4 ppt SG and 4.2 ppt PG). 

Based on the SG observations at z0 (15 km) (the second 

method of estimation), it is probable that 2.7 (1.4–4.6) ppt 

reaches the stratosphere. In summary, by taking into account PGI and SGI the total contribution is estimated to 

be between 1 and 8 ppt (values rounded). This estimated 

range is slightly lower than given in the previous Assessment (Law and Sturges et al., 2007), where a contribution 

of 5 ppt (range 3 to 8 ppt) was estimated, but is well within 

the respective ranges of uncertainty.

Total Input from Iodinated VSLS

Both SG and PG observations suggest that no significant amount of iodine enters the stratosphere. Butz et 

Table 1-9. Summary of source gas (SG) and product gas (PG) observations and modeling results to 

constrain input of halogens from VSLS into the stratosphere. Note that only observations of chlorinecontaining PGs exist; estimates of PG amounts for bromine are based solely on modeling studies and only 

upper limits of iodine-containing PGs are available. For bromine and iodine only ranges can be estimated from 

this. Details on the way that these numbers have been derived can be found in the Sections 1.3.3.1 (SG), 

1.3.3.2 (PG), and 1.3.3.3 (total). All values are given in ppt.

Halogen or Compound

Measured TTL to

CPT Abundance

(ppt Cl, Br, or I)

“Best Estimate” 

TTL Abundance 

(ppt Cl, Br, or I)

“Best Estimate”

Contribution from VSLS 

(ppt Cl, Br, or I)

Chlorine

VSL SGs 26–80 a, b 55 (38–80) c 55 (38–80)

HCl PG 0–40 d 20 (0–40) 10 (0–20)

COCl2 PG 31 ± 22 to 36 ± 26 e 32 (± 22) 16 (0–32)

Total chlorine 25–170 f 80 (40–130) f

Bromine

VSL SGs 0.7–6.5 a, b 2.7 (1.4–4.6) a, c 0.7–3.4 a, g 

PG sum 0.4–4.2 h

Total bromine 1–8 i, f

Iodine

CH3I SG < 0.05 a < 0.05 a < 0.05 a

IO, OIO PG < 0.1 j < 0.1 j < 0.1 j

Total iodine < 0.15 k

a Based on observations compiled in Table 1-7. b Entire range of values observed in TTL and at the CPT, based on Table 1-7. c Value and range at the level of zero clear-sky heating (z0). d Range of observed values from Marcy et al. (2007), Mebarki et al. (2010), Park et al. (2008). e Fu et al. (2007). f Rounded from the range of sums of PG and SG. g For the best estimate of bromine, only the SG range at the CPT has been used, as modeled PGI would include contributions from all SG broken 

down at lower altitudes. h Based on modeling work of Hossaini et al. (2010), Aschmann et al. (2009), and Liang et al. (2010); see Section 1.3.3.2. i Calculated from the range of SG observations at the tropical tropopause and the PG modeling range. j Upper limits reported by Butz et al. (2009) for IO and OIO, respectively. k Rounded from Butz et al. (2009), based on observed upper limits of IO and OIO in the lower stratosphere and photochemical modeling.
1.53

ODSs and Related Chemicals

al. (2009) derived an upper limit of 0.09–0.16 ppt of total 

iodine in the stratosphere from photochemical modeling 

and the upper limits of IO and OIO from their observations in the tropical lower stratosphere. The total amount 

of inorganic stratospheric iodine (Iy) is thus estimated to 

be below 0.15 ppt (Table 1-9).

1.3.4 Potential Influence of VSLS on Ozone

In the previous Assessment (Law and Sturges et 

al., 2007) a modeling study (Salawitch et al., 2005) was 

discussed that suggested that 4 to 8 ppt of “additional” 

stratospheric bromine from VSLS caused enhanced ozone 

depletion primarily during periods of elevated aerosol 

loading. Halogens are released efficiently from VSL SGs 

in the lowermost stratosphere, leading to ozone loss at 

these altitudes by BrO + ClO catalytic cycles. This process is enhanced in the presence of co-located high aerosol 

loading of volcanic origin. Law and Sturges et al. (2007) 

also indicated that iodine-mediated ozone loss due to 0.1 

ppt of Iy could only account for at most a few percent of 

ozone loss and only below about 17 km, depending on the 

efficiency of catalytic cycles. The associated chemistry 

remains a matter of significant uncertainty.

More recently Feng et al. (2007) and Sinnhuber et 

al. (2009) have reported qualitative agreement with the 

above studies in the sense that additional bromine from 

VSLS decreased modeled column ozone, but only by a 

small amount, except during volcanic aerosol events. The 

two studies, however, produced divergent quantitative column ozone amounts that do not fit all features of the trends 

at all times, nor do they reproduce absolute column ozone 

amounts simultaneously in both hemispheres. A detailed 

discussion of possible reasons for these discrepancies in 

terms of the respective model setups is beyond the scope 

of this Assessment, but serves to highlight the uncertainties of such modeling studies.

Sinnhuber et al. (2009) showed better agreement 

between their modeled trends in column ozone and satellite observations (1980–2005) in the midlatitude SH 

than in the NH using a two-dimensional (2-D) transport, 

chemistry, and radiation model. Without additional Br 

from VSLS, the model follows the larger features of the 

ozone trend in NH midlatitudes but generally underpredicts the observed declines in column ozone, particularly 

between 1991 and 1998, which includes a period of high 

aerosol loading from the 1991 eruption of Mt. Pinatubo. 

Modeled and measured midlatitude NH ozone depletion 

for this period of time are, however, in closer agreement 

when an additional 4 ppt of stratospheric bromine from 

VSLS are included in the calculations. The addition of 4 

ppt of bromine from VSLS enhances modeled ozone loss 

above that due to all other ODSs by as much as 40% at 

altitudes below 20 km during the period of peak aerosol 

loading following the eruption of Mt. Pinatubo (see Figure 

1-13). The effect of Bry from VSLS (Bry

VSLS) on calculated ozone during times of background aerosol loading 

is quite small, in agreement with the results of Salawitch 

et al. (2005).

Sinnhuber et al. (2009) also discussed altitudinal 

profiles of the modeled ozone changes. Figure 1-13 

compares calculated ozone abundances during NH spring 

during a year with high stratospheric aerosol (1993 with 

1980 as a reference year). The significant modeled impact of bromine from VSLS is evident and is confined 

between the tropopause and about 20 km, which is in 

qualitative agreement with the calculations by Salawitch 

et al. (2005).

We conclude from these various studies that the addition of 4–6 ppt of inorganic bromine to the stratosphere 

(i.e., 240–360 ppt of equivalent chlorine using α = 60) 

from VSLS results in a considerable increase in midlatitude column ozone loss during times of elevated aerosol 

Ozone change 1993 – 1980 [1012 molec/cm3]

Altitude [km]

Figure 1-13. Modeled differences between annual 

mean ozone concentration (molecules per cubic 

centimeter) calculated for March 1980 and March 

1993 at 47°N, illustrating the effect of additional 

bromine from very short-lived source gases during 

a period of enhanced stratospheric aerosol (i.e., influenced by the eruption of Mt. Pinatubo in 1991) 

(Sinnhuber et al., 2009). The base model run (blue 

line) includes contributions to stratospheric Bry from 

CH3Br, halon-1301, and halon-1211; the solid red 

model line includes 4 ppt of additional bromine from 

CH2Br2 and CHBr3; and the dashed red line results 

from the same model run as shown by the solid red 

line, but with total bromine loading fixed at 1959 levels. Modified from Sinnhuber et al. (2009).
1.54

Chapter 1

loading, with the largest effect in the lowermost stratosphere (altitudes below ~20 km). To date no analogous 

study has been undertaken to quantify the effects on ozone 

of additional chlorine from VSLS, which may amplify the 

effect on ozone of bromine from VSLS.

1.3.5 The Potential for Changes in 

Stratospheric Halogen from 

Naturally Emitted VSLS

As discussed in Section 1.3.4, the current emissions of natural halogenated VSLS have a small but nonnegligible impact on the halogen loading in the stratosphere. This situation could change in the future, resulting 

in a larger role for VSLS in response to factors connected 

to climate change.

Natural oceanic emissions of VSLS might change 

in the future in response to changes in climate, as might 

the efficiency with which VSLS either as SGs or as PGs 

are transported to the stratosphere. Future oceanic emissions of VSLS will potentially be influenced by changes 

in seawater temperature, pH, wind speed, mixed layer 

depth, light penetration, nutrient supply, and depth of biotic production/degradation of trace gases in a potentially 

more stratified ocean (e.g., Kloster et al., 2007; Schmittner 

et al., 2008). Hense and Quack (2009) modeled higher 

CHBr3 emissions in NH winter for the tropical Atlantic 

Ocean concurrent with a deepening of the mixed layer 

depth (MLD), enhancing the connection between the zone 

of subsurface net production and ocean-air transition. For 

nutrients, diverging effects have been found for different bromocarbons around the Mauritanian Upwelling of 

West Africa (Quack et al., 2007). Measurements in this 

region showed that while CH2Br2 concentrations were 

enhanced in colder, nitrogen-enriched and deeper waters, 

CHBr3 productivity was higher in warmer nutrient-poor 

waters, with little relationship to chlorophyll a abundance. 

Palmer and Reason (2009), however, found a correlation 

of elevated CHBr3 concentrations with chlorophyll a from 

the tropical oceans and more surface CHBr3 from waters 

with a larger MLD. Zhou et al. (2008) found a linear relationship between bromocarbon emissions and wind speeds 

above 4 meters per second over the ocean, and strong enhancements in air masses recently affected by a tropical 

storm. Taken together, seasonal MLD deepening and 

greater wind-driven surface evasion may account for an 

order of magnitude higher emissions of CHBr3 in winter 

than summer in tropical oceans (Hense and Quack, 2009; 

Palmer and Reason, 2009). This leads to the conclusion 

that if climate change would result in greater amplitudes 

of MLDs and stronger average wind velocities above the 

oceans, then future emissions of CH2Br2 and CHBr3 could 

possibly be enhanced, although other factors could either 

mitigate or enhance such effects (such as deepening of the 

subsurface production zone).

When assessing the fate of naturally emitted halocarbons in a future atmosphere, two concurring effects 

have been discussed to be relevant (i.e., changes in transport and in atmospheric degradation). Pyle et al. (2007) 

and Dessens et al. (2009) modeled higher Bry in the tropical troposphere and stratosphere by 2100 due to enhanced 

convection resulting from increased greenhouse gases and 

sea surface temperatures. This effect, which is mainly 

attributed to CH2Br2 (>90%), and secondarily to CHBr3

was, however, only obtained if VSLS were allowed to be 

emitted in the tropics (20°N–20°S). Placing the emissions 

outside the tropics resulted in no enhancement of stratospheric Bry, as the VSLS species were largely oxidized 

before injection into the stratosphere. This underlines the 

crucial influence of the geographic location of sources of 

VSLS on their impact on the stratosphere.

These projections for atmospheric transport in a 

warmer climate are, however, still under discussion, as 

other models simulate a weakening of the tropospheric circulation and convective updrafts in the tropics in the future 

(e.g., Vecchi and Soden, 2007). Furthermore, Zeng et al. 

(2008) indicate OH increases in a simulation of 2100 conditions, with the largest increases of 20–30% in the tropics 

between 3 and 8 km (although some decreases occurred 

in the boundary layer and upper troposphere due to decreased ozone in these regions). This would likely affect 

those VSLS species with predominantly OH-oxidation 

sinks (including CH2Br2) but would have little effect on 

those with predominantly photolysis-limited lifetimes 

(i.e., CHBr3).

1.3.6 Environmental Impacts of 

Anthropogenic VSLS, 

Substitutes for Long-Lived 

ODSs, and HFCs

Anthropogenic emissions of VSLS can either occur 

by intensifying natural processes (e.g., agriculture, biomass burning) or by release after industrial production of 

VSLS. This section covers the potential future influence 

of both types of emissions. Furthermore, environmental 

impacts of non-ozone depleting VSLS being considered 

as substitutes for long-lived HFCs are also discussed in 

this section.

1.3.6.1 Evaluation of the Impact of Intensified

Natural Processes on Stratospheric

Ozone

There are a number of ways in which human activity might intensify emissions of otherwise naturally 
1.55

ODSs and Related Chemicals

occurring VSLS. An area of high potential impact is 

changes in agriculture practices. It is noted in Section 

1.2.1.5 that some crops are capable of producing methyl 

halides. For VSLS, anticipated increases of future rice 

production in the tropics (Bhatia et al., 2010) could be 

important, as rice cultivation is known to emit considerable amounts of CH3I (e.g., Lee-Taylor and Redeker, 

2005). Furthermore, Gan et al. (1998) and Rhew et al. 

(2003) reported that emissions of methyl halides from 

halogen-methylating plants are strongly linked to the halide content of the growth substrate. With increasing salinization of soils through irrigation, coastal inundation, 

or deliberate cultivation of poorer soils, emissions might 

increase in the future. This issue is emphasized by the 

finding of Sive et al. (2007) that fluxes of CH3I from terrestrial biomes over the eastern U.S. are presently comparable with oceanic sources over the Atlantic. Another 

possible impact from future agriculture would be the 

aquaculture of marine algae for food, chemicals, biofuel, 

and for sequestering atmospheric carbon (Packer, 2009; 

Aizawa et al., 2007). Furthermore, water chlorination, 

both of drinking water and domestic wastewater, and 

also seawater coolant in coastal power plants may contribute small amounts of halogen to the stratosphere and 

this contribution could change in the future, but no new 

information has become available since the previous Assessment concerning these potential contributions.

1.3.6.2 Very Short-Lived New ODSs and Their

Potential Influence on Stratospheric

Halogen

Chlorocarbons such as CCl2CCl2 (perchloroethene) and CHClCCl2 (trichloroethene) have been used for 

many decades in large quantities as industrial solvents. 

Additionally, new very short-lived chlorocarbons, bromocarbons, and iodocarbons (with lifetimes <0.5 years) 

have recently been introduced or are proposed as potential substitutes for long-lived ODSs. The pathways and 

reactions by which halogens from these anthropogenic 

VSLS reach the stratosphere are similar to those of natural compounds, such as CH2Br2 and CHBr3, as discussed 

in Sections 1.3.2–1.3.3. These anthropogenic VSLS are 

considered separately from the long-lived ODS, as the 

traditional concept of a single, geographical-independent 

Ozone Depletion Potential (ODP) does not apply and their 

impact depends on the location and season of emissions.

The issue of ODPs for short-lived ODSs has already been discussed for n-propyl bromide and CF3I in 

the previous two Assessments (Ko and Poulet et al., 2003; 

Law and Sturges et al., 2007). For CHCl3, CCl2CCl2, and 

CHClCCl2, Kindler et al. (1995) have derived mean, semiempirical ODPs of 0.008–0.01, 0.006–0.007, and 0.0005–

0.0007, respectively, based on a 2-D modeling analysis 

and source gas atmospheric distributions that are similar 

to those observed. Furthermore, in Law and Sturges et al. 

(2007) it was estimated that 1% of uniformly distributed 

land mass emissions of a theoretical VSLS with a tropospheric lifetime of 25 days would reach the stratosphere. 

Consequently the ODP of such a compound with one 

chlorine atom and similar molecular weight to CFC-11 

has been calculated as being about 0.003. This approach, 

however, neglects the issue of the ODP being sensitive to 

the location and seasonality of emission.

Wuebbles et al. (2009) updated an earlier 2-D 

model approach of Li et al. (2006) to calculate the ODPs 

for CF3I by using a 3-D global chemical transport model 

(CTM). An ODP of 0.016 and 0.008 is calculated for CF3I 

surface emissions in the tropics and in the midlatitudes, 

respectively, which compares well with the earlier ODP 

values from Li et al. (2006) of 0.018 and 0.011 for CF3I released in the tropics and in the midlatitudes, respectively. 

These new figures confirm the higher probability of VSLS 

released in the tropics reaching the stratosphere, which results in higher ODPs for VSLS released near the equator. 

In the last Assessment (Law and Sturges et al., 2007) it 

was also noted that ODPs are dependent upon the altitude 

of emission, which would be relevant if CF3I were used on 

aircraft as a replacement for halons. However, no update 

on this is available.

1.3.6.3 Evaluation of Potential and In-Use

Substitutes for Long-Lived ODSs

A wide range of chemicals with zero or near-zero 

ODPs and low GWPs are currently being considered as 

substitutes for ODSs and long-lived HFCs. In Table 1-10, 

atmospheric lifetimes of substitutes that are already in use 

or potentially could be used in applications such as refrigeration, foam blowing, fire control, propellants, and solvents are listed. The list is not restricted to VSLS because 

many longer-lived HFCs are in use or are being considered as ODS replacements.

In Table 1-11, substitutes that are already in use or 

have been identified as most likely to be used in the near 

future are listed together with their applications, potential 

feedstocks, and atmospheric lifetime. Furthermore, metrics of environmental effects such as GWP, Photochemical Ozone Creation Potential (POCP), and trifluoroacetic 

acid (TFA) formation are specified. For reasons of completeness and to link with the issue of global warming and 

climate change, fluorinated substances are also listed that 

are not ODS substitutes in the strictest sense (i.e., NF3, 

hexafluorobutadiene, COF2).

Hydrocarbons receive use as ODS substitutes 

in refrigeration and foam-blowing applications. They 

do not deplete stratospheric ozone and they have very 

small GWPs. However, the oxidation of hydrocarbons 
1.56

Chapter 1

Table 1-10. Local and partial lifetimes of in-use and potential replacement compounds for long-lived 

ODSs (long-lived substitutes in italics).

Compound

Local Lifetime

from Previous 

Assessment 

(days)

OH Lifetime 

(days)

Photolysis 

Lifetime 

(days)

New Local 

Lifetime

(days)

Notes

Hydrocarbons

CH2=CHCH3 (propene) 0.37 0.35 0.35 1, 5

(CH3)2C=CH2 (isobutene) 0.20 0.20 1, 5

CH3CH2CH3 (propane, R-290) 12.5 12.5 1, 3

(CH3)2CHCH3 (isobutane, R-600a) 6.0 6.0 1, 5

CH3CH2CH2CH2CH3 (n-pentane) 3.7 3.4 3.4 1, 6

c-CH2CH2CH2CH2CH2

 (cyclopentane)

2.9 2.7 2.7 1, 6

(CH3)2CHCH2CH3 (isopentane) 3.7 3.4 3.4 1, 6

CH3OCHO (methyl formate) 58 72 72 1, 7

(CH3)2CHOH (isopropanol) 4.8 2.0 2.0 1, 3

CH3OCH2OCH3 (methylal) 2.2 2.2 1, 8

Hydrofluorocarbons

CH3CH2F (HFC-161) 77 66 66 1, 3

CH2FCH2F (HFC-152) 219 146 146 1, 4

CH3CHFCH3 (HFC-281ea) 22 23 23 1, 3

CHF2CH2CF3 (HFC-245fa) 7.6 years 8.2 years 116 years 7.7 years 2, 3, 9

CH3CF2CH2CF3 (HFC-365mfc) 8.6 years 9.3 years 125 years 8.7 years 2, 3, 9

CHF2CHF2 (HFC-134) 9.6 years 10.5 years 134 years 9.7 years 2, 3, 9

CF3CHFCF3 (HFC-227ea) 34.2 years 44.5 years 310 years 38.9 years 2, 4, 9

CF3CH2CF3 (HFC-236fa) 240 years 253 years 5676 years 242 years 2, 3, 9

Unsaturated Fluorocarbons

CH2=CHF 2.1 2.1 1, 4

CH2=CF2 4.0 4.0 1, 4

CF2=CF2 1.1 1.1 1, 4

CH2=CHCH2F 0.7 0.7 1, 4

CH2=CHCF3 7.6 7.6 1, 4

CH2=CFCF3 10.5 10.5 1, 4

E-CF3CH=CHF 16.4 16.4 1, 4

E-CF3CF=CHF 4.9 4.9 1, 4

Z-CF3CF=CHF 8.5 8.5 1, 4

CF2=CFCF3 4.9 4.9 1, 4

CH2=CHCF2CF3 7.9 7.9 1, 4

CF2=CFCF=CF2 1.1 1.1 1, 10

Unsaturated Chlorocarbons

E-CF3CH=CHCl 26 26 1, 11

CH2=CHCl 1.5 1.5 1, 4
1.57

ODSs and Related Chemicals

Table 1-10, continued.

Compound

Local Lifetime

from Previous 

Assessment 

(days)

OH Lifetime 

(days)

Photolysis 

Lifetime 

(days)

New Local 

Lifetime

(days)

Notes

CH2=CCl2 0.9 0.9 1, 4 

CHCl=CCl2 4.6 4.9 > 15000 4.9 1, 4, 21

CCl2=CCl2 99 90 90 1, 4

CF2=CFCl 1.4 1.4 1, 12

CF2=CFCF2Cl ~ 5 ~ 5 1, 13 

CF2=CFCF2CFCl2 ~ 5 ~ 5 1, 13

Unsaturated Bromocarbons

CFBr=CF2 1.4 1.4 1, 14

CHBr=CF2 2.3 2.3 1, 14

CH2=CBrCF3 2.7 2.7 1, 14

CH2=CBrCF2CF3 3.1 3.1 1, 14

CH2=CHCF2CF2Br 6.5 6.5 1, 14

Fluorinated Ethers, HFE

CH3OCH2CF3 (HFE-263fb2) 37 23 23 1, 15, 17

CH3OCHFCF3 (HFE-254eb2) 88 88 1, 4

CH3OCH2CF2CF3 40 21 21 1, 16, 17

CH3CH2OCF2CHF2 1826 (error) 64 64 1, 4

CF3CH2OCH2CF3 146 105 105 1, 4

CH3OCH(CF3)2 61 61 1, 18

Fluorinated Ketones

CF3CF2C(O)CF(CF3)2 (FK-5-1-12) < 14 > 63 years 7–14 7–14 19

Fluorinated Alcohols

CH2FCH2OH 15 12.9 12.9 1, 4

CHF2CH2OH 51 51 1, 4

CF3CH2OH 150 142 142 1, 4

C2F5CH2OH 142 143 143 1, 4

C4F9CH2OH 164 142 142 1, 4

CF3CHFCF2CH2OH 124 112 112 1, 5

Special Compounds

CF3CF2CF2I

 (1-iodo-heptafluoropropane)

< 2 < 2 20

CH3I (methyl iodide) 7 158 7 (4–12) 7 1, 5, 21, 22

COF2 (carbonyl fluoride) 5–10 27

PBr3 < 0.01 0.14 < 0.01 22, 23 

NH3 Few days 93 Few days 1, 3, 24

SO2F2 (sulfuryl fluoride) >300 years 630 years 36 years 9, 25

NF3 740 years 500 years 26

CH3CH2Br (bromoethane) 34 41 41 1, 3
1.58

Chapter 1

Notes:

1. These local OH lifetimes are calculated using an average tropospheric OH concentration of 1×106 molecule/cm3 and 

the OH reaction rate constant at T = 275 K. Local lifetimes quoted here are not meant to be estimates of global lifetimes, which, for short-lived gases depend on the emission location and season as well as local atmospheric conditions. 

The concept of a single global lifetime, ODP, or GWP is inappropriate for such short-lived gases.

2. Lifetimes for long-lived compounds with respect to reaction with 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).

3. OH reaction rate constant taken from JPL 06-2.

4. OH reaction rate constant taken from JPL 10-6. JPL 10-6 is cited here whenever there is a change in a rate constant 

recommendation or the accompanying note. It does not necessarily mean that a major change was recommended for 

a rate constant. Nevertheless, updates in JPL 10-6 reflect improved kinetic understanding.

5. OH reaction rate constant taken from Atkinson et al. (2008).

6. OH reaction rate constant taken from Calvert et al. (2008).

7. OH reaction rate constant taken from Le Calvé et al. (1997).

8. OH reaction rate constant taken from Porter et al. (1997).

9. Lifetime shown in “Photolysis Lifetime” column corresponds to the overall stratospheric lifetime. See Table 1-3.

10. OH reaction rate constant taken from Acerboni et al. (2001).

11. OH reaction rate constant taken from Sulbaek Andersen et al. (2008).

12. OH reaction rate constant taken from Abbatt and Anderson (1991).

13. Local lifetime estimated as similar to that of CF3CF=CF2.

14. OH reaction rate constant taken from Orkin et al. (2002).

15. OH reaction rate constant taken from Oyaro et al. (2005).

16. OH reaction rate constant taken from Oyaro et al. (2004).

17. Room temperature data only; OH reaction lifetime calculated assuming a temperature dependence (E/R) of 500 K.

18. OH reaction rate constant from Chen et al. (2005b).

19. Lifetimes taken from Taniguchi et al. (2003).

20. Photolysis lifetime estimated from comparison of UV spectra and lifetimes for CF3I and CF3CF2CF2I.

21. Photolysis lifetime taken from Table 2-4 in Ko and Poulet et al. (2003).

22. Local lifetime taken from Table 2-1 in Law and Sturges et al. (2007).

23. OH reaction rate constant taken from Jourdain et al. (1982). Local lifetime is probably dictated by photolysis.

24. Local lifetime taken from IPCC/TEAP (2005); it is dictated by washout rates (see Box 1-4).

25. From Papadimitriou et al. (2008b) and Mühle et al. (2009). The total lifetime is primarily due to ocean uptake.

26. From Prather and Hsu (2008) with the lifetimes recalculated using the JPL 06-2 recommended rate constants for the 

O(1D) reactions corrected for the deactivation channel (see also Section 1.2.2).

27. Local lifetime taken from Wallington et al. (1994).

Table 1-10, continued (notes).
1.59

ODSs and Related Chemicals

Table 1-11. Applications, feedstocks, and environmental impact of in-use and likely potential future 

ODS-substitutes (long-lived substitutes in italics).

Compound Application Feedstock Lifetime 1 POCP 2,3,4 GWP100 5

TFA

molar

yield

Hydrocarbons

propene, CH3CH=CH2

(HC1270) 

refrigerant petroleum 0.35 d 117 2 5 6,7 0%

isopentane,

(CH3)2CHCH2CH3

foams (replacing 

CFC-11 and

HCFC-141b)

petroleum/

natural gas

3.4 d 34 2 0%

cyclopentane,

c-CH2CH2C2CH2CH2

foams (replacing 

CFC-11 and

HCFC-141b)

natural gas 2.7 d 51 3 0%

n-pentane,

CH3CH2CH2CH2CH3

foams (replacing 

CFC-11 and

HCFC-141b)

petroleum/

natural gas

3.4 d 40 2 0%

isobutene,

(CH3)2C=CH2 

foams (replacing 

CFC-12)

petroleum 0.2 d 75 2 0%

methylal, 

CH3OCH2OCH3

foams (replacing

CFC-11 and

HCFC-141b) 

methanol 2.2 d 32 3 0%

methyl formate,

CH3OCHO

foams (replacing 

HCFCs)

methanol/

formic acid

72 d 3 2 0%

Hydrofluorocarbons 

HFC-245fa,

CHF2CH2CF3

foams (replacing 

HCFCs)

HF + CCl4 + 

CH2=CHCl

7.7 y 0.1 3 1050 <10% 8,9

HFC-365mfc,

CH3CF2CH2CF3

foams (replacing 

HCFCs)

HF + CCl4 + 

CH2=CClCH3

8.7 y 0.1 3 842 <10% 9,10

HFC-227ea,

CF3CHFCF3

propellant for CFC-12 

and firefighting for 

halon-1301

HF + CHCl3 →

CHClF2 → 

CF3CF=CF2

39 y <0.1 3 3580 100% 11

HFC-236fa,

CF3CH2CF3

firefighting for 

halon-1211/1301 (civil 

aviation); refrigerant

HF + CCl4 + 

CH2=CCl2

242 y <0.1 3 9820 <10% 12

Unsaturated Fluorocarbons 

trans-HFC-1234ze, 

trans-CF3CH=CHF 

foams (for CFC-12) a) CHCl3 + HF →

CHClF2 →

CF3CF=CF2 →

CF3CHFCH2F

or

b) HFC-245fa 

16.4 d 6.4 4 7 6,13 < 10% 14,9

HFC-1234yf, 

CF3CF=CH2

refrigerant potentially 

for replacing 

HFC-134a

CHCl3 + HF →

CHClF2 →

CF3CF=CF2 →

CF3CHFCH2F 

10.5 d 7.0 4 4 6,15 100% 16,17

Fluorinated Ketones 

FK-5-1-12,

CF3CF2C(O)CF(CF3)2

halon-1301 

replacement 

C3F6

C2F5COF

7-14 d < 10% 18
1.60

Chapter 1

Compound Application Feedstock Lifetime 1 POCP 2,3,4 GWP100 5

TFA

molar

yield

Special Compounds 

ammonia, NH3 CFC replacement in 

refrigeration

N2 + H2 few days NA 0%

1-iodoheptafluoropropane, 

CF3CF2CF2I 

halon replacement (in 

Russia: -2402 and/or 

-1301) 

CF3I + C2F4 < 2 d NA < 10% 18

bromoethane, 

CH3CH2Br 

solvent and part of 

halon-2402 blend

a) HBr + C2H4

or

b) HBr + 

CH3CH2OH

41 d 4.2 3 0%

2-bromo-3,3,3-

trifluoroprop-1-ene,

CH2=CBrCF3

halon-1211 

replacement 

CH2=CHCF3 2.7 d 9.3 3 < 10% 19, 9

methyl iodide, CH3I CH3Br replacement CH3OH / I2 7 d 1.0 3 0%

sulfuryl fluoride, 

SO2F2

CH3Br replacement KF + SO2F2 36 y 20 NA 4740 21 0%

NF3 etchant to replace 

C2F6

a) NH3 + F2

or

b) (NH4)HF2

500 y 22 NA 17,500 0%

hexafluorobutadiene etchant 1.1 d 10 3 0%

carbonyl fluoride, 

COF2

etchant a) CO2 + F2

or

b) Oxidation of 

C2F4

 5–10 d NA 0%

Notes:

1. See Table 1-10 for notes on lifetime derivations.

2. Derwent et al. (2007).

3. No value for the POCP (Photochemical Ozone Creation Potential) of this substance has been published by full trajectory analyses. This value was 

calculated based on information from Derwent et al. (1998) and Jenkin (1998), using the method that Wallington et al. (2010) used for unsaturated 

HFCs. The derived POCP is an approximation and is related to the molecular structure of the compound and its OH-reactivity. The calculations were 

only conducted for species lost mainly through OH-radical reactions; others are declared as NA (not available). .

4. Wallington et al. (2010).

5. GWPs from Chapter 5 Appendix (Table 5A.1) if not specified differently.

6. GWPs are generally calculated assuming a uniform global distribution of the trace gas. Hence, such methods are in principle not accurate for shortlived gases (i.e., gases with atmospheric lifetime shorter than about 0.5 years) because the atmospheric distribution of short-lived gases will likely be 

non-uniform and will also depend upon the location and season of emission. Thus, GWPs calculated in this way for short-lived gases provide only a 

very rough measure of the time-integrated radiative effect of these gases.

7. This value represents an indirect GWP (100-year), derived with a global three-dimensional Lagrangian model (Collins et al., 2002) and an emission 

distribution related to fossil-fuel use. The indirect GWP includes influences on the abundance of methane (through changes in OH), tropospheric 

ozone, and on CO2.

8. The oxidation of HFC-245fa gives CF3CHO as a major product (Chen et al., 1997).

9. Photolysis is the main fate of CF3CHO (lifetime of approximately 19 hours in lower troposphere for overhead Sun; Calvert et al., 2008) and this does 

not give TFA. Reactions with water (Sulbaek Andersen et al., 2006) or OH radicals are minor losses of CF3CHO and can lead to TFA (Hurley et al., 

2006). The yield of TFA is estimated to be <10%.

10. The oxidation of HFC-365mfc gives CF3CHO and COF2 as major products (Inoue et al., 2008).

11. Zellner et al. (1994).

12. The oxidation of HFC-236fa gives CF3C(O)CF3 (Møgelberg et al., 1995). Photolysis is the main fate of CF3C(O)CF3 (lifetime of approximately 6.2 

days in lower troposphere for overhead Sun; Calvert et al., 2008). The photolysis of CF3C(O)CF3 will give CF3CO radicals, a small fraction of which 

Table 1-11, continued.
1.61

ODSs and Related Chemicals

in the presence of elevated nitrogen oxides (NOx) leads 

to the formation of carbon dioxide (CO2) (e.g., Collins 

et al., 2002) and can contribute to the production of tropospheric ozone, which indirectly affects atmospheric 

levels of methane (CH4). POCPs are used to assess 

the relative efficiency of hydrocarbon species for forming tropospheric ozone. POCPs, expressed relatively 

to ethene (POCP = 100), are shown in Table 1-11 and 

have either been taken from Derwent et al. (2007) or 

Wallington et al. (2010) or have been calculated using 

the method of Derwent et al. (1998) and Jenkin (1998) 

(specified for unsaturated HFCs by Wallington et al., 

2010). While this methodology provides a rough estimate of POCPs for trace gases that are predominately 

destroyed by atmospheric oxidation, its application to 

species for which photolysis and hydrolysis are dominant loss processes is not valid.

POCPs are lower for saturated hydrocarbons, 

typically used as refrigerants and for foam blowing (i.e., 

propane, isobutane, pentanes), but higher for the alkenes 

(propene, isobutene). Although releases associated with 

refrigeration are minor in comparison to hydrocarbon 

emissions from traffic and solvent usage, the combined 

impact of these hydrocarbons could be non-negligible, 

especially in already polluted regions. For the HFCs and 

other compounds discussed as substitutes in Table 1-11, 

the POCP values are generally smaller than 10 and hence 

their relevance for the formation of tropospheric ozone 

will be very small.

The unsaturated halocarbons (halogenated alkenes) contain a C=C double bond that makes them readily susceptible to degradation by OH (and ozone (O3)). 

As a result, their atmospheric lifetimes are typically 

days to weeks. This group of compounds can be further divided into unsaturated HFCs (sometimes referred 

to as hydrofluoro-olefins, HFOs) and chloro- and bromoalkenes. Unsaturated HFCs are being considered as 

potential replacements for ODSs and HFCs in mobile air 

conditioners and for foam blowing, as they typically have 

smaller GWPs than saturated (long-lived) HFCs. These 

smaller GWPs are partially offset for these unsaturated 

compounds by the production of tropospheric ozone, a 

strong greenhouse gas.

Chlorinated and brominated alkenes also have 

short lifetimes but still can have non-negligible ODPs 

(e.g., Kindler et al., 1995). Chlorinated alkenes such as 

CCl2CCl2 (perchloroethene) and CHClCCl2 (trichloroethene) have been used in large quantities for decades as 

industrial solvents and a certain amount of this chlorine 

reaches the stratosphere as a result of use of these gases 

(see Section 1.3.3.1).

Oxygenated fluorocarbons are another group of 

ODS replacement compounds with lifetimes typically 

spanning days to months (Table 1-10 and Table 1-11). 

In fact a number of fluorinated ethers (e.g., HFE-7100, 

HFE-7200) have been already used as first-generation 

ODS substitutes in niche applications such as refrigerants, as solvents, and as heat transfer fluids, but 

no information about their atmospheric abundance is 

available.

The atmospheric degradation of (saturated and 

unsaturated) HFCs and oxygenated fluorocarbons 

is initiated by their reaction with OH radicals or O3, 

(important only for unsaturated HFCs). The oxygenated products from these reactions are removed from 

the atmosphere via deposition and washout processes 

and may accumulate in oceans, lakes, and other reservoirs. A potential toxic by-product of the atmospheric 

degradation of fluorocarbons with CF3 groups is trifluoroacetic acid (TFA; CF3C(O)OH). Yields of TFA 

from the degradation of different HFCs are included 

in Table 1-11. TFA is removed from the atmosphere 

by wet deposition but is known to accumulate in certain ecosystems. However, much uncertainty remains 

Table 1-11, continued.

can react further to give TFA (Hurley et al., 2006). The molar yield of TFA from the atmospheric oxidation of HFC-236fa is estimated to be <10%.

13. This direct GWP (Orkin et al., 2010) was calculated using a semi-empirical approach with the revised lifetime of 16.4 days given in Table 1-10 but 

without accounting for inhomogeneous emission and atmospheric mixing ratio distributions.

14. Søndergaard et al. (2007); Javadi et al. (2008).

15. This direct GWP is from Papadimitriou et al. (2008a) and Orkin et al. (2010) and was calculated with the revised lifetime of 10.5 days given in Table 

1-10 but without accounting for inhomogeneous emission and atmospheric mixing ratio distributions.

16. Nielsen et al., 2007.

17. Hurley et al., 2008.

18. Photolysis will lead to the formation of C2F5 radicals. Approximately 1–10% of CxF2x+1 radicals are converted in the atmosphere into Cx-1F2x-1C(O)

F (Wallington et al., 2006). The sole atmospheric fate of the 1–10% CF3C(O)F produced is hydrolysis to give TFA.

19. The atmospheric oxidation of CF3CBr=CH2 is believed to give CF3C(OH)=CH2 and/or CF3C(O)CH2OH (Orkin et al., 2002; Sulbaek Andersen et al., 

2009a). Further oxidation will likely give CF3CHO.

20. Mühle et al. (2009).

21. Derived from Papadimitriou et al. (2008b) with a radiative efficiency rounded to two decimal places (0.22 W/m2ppb versus 0.222 W/m2ppb) used in 

Papadimitriou et al. (2008b) (see Chapter 5).

22. See Table 1-10, footnote 26.
1.62

Chapter 1

in understanding the processes involved in maintaining 

measured abundances of TFA in today’s aquatic environments. Whereas average TFA concentrations of 

200 nanograms per liter in deep ocean waters suggest 

natural sources in the ocean (Frank et al., 2002), TFA 

levels were below detection in old ice core and groundwater samples (Nielsen et al., 2001).

At present HFC-134a (CF3CH2F) is the HFC that 

contributes the most TFA to the environment among 

ODS replacement gases. With global emissions of 149 ± 

27 Gg/yr in 2008 (Section 1.5) and a TFA yield of 

7–20% (Figure 1-14; Wallington, 1996), a current yearly 

global input of 9–35 Gg/yr of TFA can be derived from 

HFC-134a. The unsaturated HFC-1234yf (CF3CF=CH2) 

is being considered as a replacement for HFC-134a in 

mobile air conditioners (Hurley et al., 2008). The TFA 

production yield of the atmospheric degradation of HFC1234yf is 100% (Hurley et al., 2008; Figure 1-14; Table 

1-11). The use of HFC-1234yf has the potential to influence TFA concentrations near the source regions to a 

greater extent than HFC-134a, owing to its shorter lifetime and its higher efficiency to produce TFA. A recent 

modeling study has shown that deposition of TFA could 

average 0.16–0.24 kg/km2 by 2017 in the continental 

U.S. from the full adoption of HFC-1234yf in mobile air 

conditioners, and that concentrations of TFA in Eastern 

U.S. rainfall would more than double as a result in comparison to today’s values (Luecken et al., 2010).

Apart from TFA, monofluoroacetic acid (MFA) 

and difluoroacetic acid (DFA) could potentially be produced in the degradation of HFCs. MFA and DFA have 

been measured in the atmosphere and their environmental

occurrence has been discussed, for example by Römpp et 

al. (2001) and Martin et al. (2003).

Replacements for halons (firefighting) and CH3Br 

(agriculture) are also being sought. Whereas for halons 

various substitutes are being considered for different applications, SO2F2 is specifically considered as an important replacement for CH3Br (see also Section 1.5.2.2). 

This compound is, however, toxic to humans and has a 

longer atmospheric lifetime (36 years) than CH3Br (Papadimitriou et al., 2008b).

Ammonia (NH3) used as a replacement in refrigeration could potentially influence the tropospheric particle 

loading via its reaction with nitric acid (HNO3) to form 

ammonium nitrate (NH4NO3). However, emissions from 

in-use systems are negligible compared to NH3 emitted 

from agriculture and natural sources.

Nitrogen trifluoride (NF3), hexafluorobutadiene, 

and carbonyl fluoride (COF2) can be used as substitutes 

for the extremely long-lived perfluorocarbons (PFCs) as 

etchants in the electronics industry. In fact, NF3 has already become an important chemical for silicon etching, 

e.g., in the production of flat computer and television 

screens (see also Section 1.5), though it has a long lifetime 

of 500 years and a high GWP100 of 17,500 (Chapter 5).

Finally, information on feedstock for the production of many substitutes is provided in Table 1-11. As 

different production methods exist for specific molecules, 

this list is not exhaustive, but it is meant to serve as a 

starting point for a discussion of emission that could potentially arise during production of substitute chemicals. 

Although emissions to the atmosphere during production 

are believed to be generally small, a possible pathway of 

ODSs to the atmosphere could be the production of HFCs, 

which can include the reaction of carbon tetrachloride 

(CCl4) with chlorinated alkenes and subsequent replacement of the Cl-atoms with fluorine (see also the discussion 

on CCl4 in Section 1.2.1.3).

Figure 1-14. Degradation scheme of HFC-1234yf 

(yielding 100% TFA; CF3C(O)OH) and of HFC-134a 

(yielding 7–20% TFA) (Wallington et al., 1996; Hurley 

et al., 2008).
1.63

ODSs and Related Chemicals

1.4 CHANGES IN ATMOSPHERIC HALOGEN

1.4.1 Chlorine in the Troposphere 

and Stratosphere

1.4.1.1 Tropospheric Chlorine Changes

Total organic chlorine from long-lived gases 

reached its peak in the troposphere at 3660 ± 23 ppt between 1992 and 1994 and has since declined. While this 

peak value includes 80 ppt from VSLS, the quoted error 

encompasses only the difference in total Cl determined 

from two different sampling networks and not uncertainties related to the contribution of Cl from VSLS (see Section 1.3). By mid-2008, tropospheric organic Cl from 

long-lived gases had declined by 8.4% from its peak to a 

value of 3352 ± 4 ppt (Table 1-12 and Figure 1-15). The 

main drivers behind this decline have changed over time. 

Methyl chloroform (CH3CCl3) now contributes much less 

to the decline in total Cl than it did in the mid-1990s and 

early 2000s. This is because the global tropospheric mixing ratio of CH3CCl3 has declined substantially: by 2008 it 

accounted for 32 ± 1 ppt of Cl, or only 1%, of tropospheric 

Cl (Table 1-12). As a result, declines in Cl from CH3CCl3

have become less important to total Cl changes and this 

influence will continue to diminish in the future.

As the influence of CH3CCl3 on tropospheric Cl 

changes has diminished, declines in tropospheric Cl have 

slowed. Tropospheric chlorine from long-lived gases 

changed at a rate of −14 ppt Cl/yr (−0.42%/yr) during 

2007–2008, compared to −21 ppt/yr during 2003–2004, 

and was slower than the change of −23 ppt/yr projected 

for 2007–2008 in the A1 (most likely, or baseline) scenario of the previous Assessment. The decline observed 

during 2007–2008 was also about half as rapid as measured in 1995–1996 (−28 ppt Cl/yr or −0.8%/yr; Clerbaux 

and Cunnold et al., 2007) (Table 1-12). For reference, the 

mean annual rate of change in tropospheric Cl from longlived gases averaged over 1996–2008 was −0.6(± 0.05)%/

yr in data from the NOAA or AGAGE global sampling 

networks.

The decline in tropospheric Cl has also slowed recently because of increases observed in the accumulation 

rate of hydrochlorofluorocarbons (HCFCs) since 2004 

(Table 1-12). By 2008, Cl from HCFCs was increasing 

at a faster rate (10.6 ppt Cl/yr) than had been observed in 

2000 and 2004 and Cl from HCFCs accounted for 251 ± 3 

ppt Cl, or 7.5% of total tropospheric chlorine.

3300

3400

3500

3600

3700

AGAGE

NOAA

Scenario Ab ('02)

Scenario A1('06)

-2.0%

-1.0%

0.0%

1.0%

1990 1995 2000 2005 2010

Year

Tropospheric Cl (ppt) Cl rate of change (per yr)

Figure 1-15. Top panel: The tropospheric abundance (ppt) of organic chlorine (CCly) from the NOAA (gray) and 

AGAGE (black) global measurement networks (updates of Montzka et al., 2003, 

and O’Doherty et al., 2004). Quantities 

are based upon independently measured mixing ratios of CFC-11, CFC12, CFC-113, HCFC-22, HCFC-141b, 

HCFC-142b, methyl chloroform, carbon 

tetrachloride, and halon-1211. Results 

for CFC-114 and CFC-115 from Prinn et 

al. (2000) are used in both aggregations. 

An additional constant 550 ppt was added for CH3Cl and 80 ppt was added for 

short-lived gases such as CH2Cl2, CHCl3, 

CCl2CCl2, and COCl2 (consistent with 

40–130 ppt discussed in Section 1.3). 

Bottom panel: Annual rates of change 

(% per year) determined from 12-month 

differences. In both panels, observations 

are compared with the baseline scenario 

(Ab) from WMO 2002 (green line; Montzka and Fraser et al., 2003) and the baseline scenario A1 from WMO 2006 (red 

line; Daniel and Velders et al., 2007).
1.64

Chapter 1

But while increases in Cl from HCFCs have accelerated, declines in Cl from CFCs have become more rapid. 

The 2004–2008 period was the first time that declines derived for tropospheric Cl from long-lived CFCs as a class 

were larger than the decline from CH3CCl3 (Table 1-12). 

CFCs accounted for 2076 ± 3 ppt Cl during 2008. Chlorine from CFCs (as a class) declined by just over 13 ppt 

(0.6%) from 2007–2008 and this Cl decrease was comparable to the summed influence of other decreasing compounds (CCl4, CH3CCl3, and halon-1211). The fraction of 

total organic chlorine attributable to CFCs has remained 

fairly constant at about 62% over the past decade even as 

Cl from CFCs has decreased by 4% from its 1998 peak.

As mentioned above, the decline in tropospheric Cl 

since 2004 was slower than anticipated in scenario A1 of 

the 2006 Ozone Assessment (Daniel and Velders et al., 

2007) (Figure 1-15). Measured tropospheric Cl from 

long-lived gases in 2008 was ~40 ppt higher than projected in the A1 scenario. This slower decline in tropospheric 

Cl resulted mostly from CFC-11 and CFC-12. Although 

global mixing ratios of CFC-11 and CFC-12 declined 

more rapidly during 2005–2008 than in earlier years, Cl 

from these CFCs in 2008 was ~28 ppt Cl higher than had 

been projected. HCFCs also played a role in the slower 

than anticipated Cl decline. By 2008, mixing ratios observed for HCFC-22 and HCFC-142b contributed ~8 ppt 

more Cl than had been projected (see Figure 1-1). Some 

of the discrepancy for CFC abundances arises because the 

calibration of CFC-12 measurements by the NOAA group 

has been revised with improved standards since the last 

Assessment. Observed changes for other Cl-containing 

compounds during 2005–2008 were similar to those projected in the A1 scenario.

1.4.1.2 Stratospheric Chlorine Changes

As discussed in previous Assessments, the stratospheric burden of chlorine is controlled by input from the 

troposphere and by transport and mixing processes. Timescales for air to be transported to higher altitudes in the 

midlatitude and polar stratosphere after crossing the tropopause range from 3 to 6 years. As a result, changes in 

stratospheric chlorine and bromine abundances lag behind 

tropospheric changes and, furthermore, are smoothed and 

flattened because of mixing (Waugh and Hall, 2002). 

Changes in total stratospheric Cl calculated from tropospheric observations allow this lag to be visualized for 

the midlatitude stratosphere (Figure 1-16) (update of 

Engel et al., 2002). They also show how the magnitude 

of the decline in stratospheric Cl is expected to be altitude 

dependent.

Most chlorine enters the stratosphere chemically 

bound to carbon (organic Cl) in long-lived source gases 

and undergoes photochemical oxidation to inorganic forms 

as air is transported to higher altitudes within the stratosphere. Long-term changes in stratospheric inorganic Cl 

abundances are derived from ground-based and satellitebased instruments measuring total column abundances or 

Table 1-12. Contributions of halocarbons to total chlorine in the troposphere.

Total Cl * (ppt Cl) Contribution to Total Cl (%) Rate of Change in Total Cl **

(ppt Cl / yr)

Mid2000Mid2004 Mid-2008 Mid2000Mid2004 Mid-2008 2000 2004 2008

All CFCs 2154 2123 2076 (3) 61% 62% 62% −1.9 (1.9) −9.4 (2.0) −13.2 (0.8)

CCl4 392 377 359 (6) 11.2% 11.0% 10.7% −4.0 (0.7) −4.0 (0.6) −5.1 (0.7)

HCFCs 182 214 251 (3) 5.2% 6.3% 7.5% 9.2 (0.8) 6.0 (1.3) 10.6 (0.5)

CH3CCl3 136 66 32 (1) 3.9% 1.9% 1.0% −26.8 (1.0) −13.4 (2.1) −6.2 (0.3)

Halon12114.02 4.26 4.17 (0.1) 0.1% 0.1% 0.1% 0.1 (0.00) 0.0 (0.08) −0.1 (0.01)

Total Cl 3499 3414 3352 (4) −23 (2.4) −21 (3.2) −14 (1.2)

−0.67% −0.61% −0.42%

* An average of AGAGE and NOAA/ESRL global means was used to derive these mid-year mixing ratios. Uncertainties are given 

in parentheses and represent 1 standard deviation of the results reported by different global networks. They do not include an 

additional amount for potential systematic calibration errors, errors associated with deriving tropospheric means from surface 

measurements, or uncertainties in Cl contributed from VSLS.

** Total Cl changes and relative rates of change are calculated assuming a constant 550 ppt Cl from CH3Cl and a constant 80 ppt from 

VSLS (see Section 1.3). Rates of change were calculated as the difference between the indicated year and the previous year and 

relative rates were normalized by the mean mixing ratio over these two-year periods. Numbers for past years differ slightly from 

previous Assessments because of updated calibration information (see text).
1.65

ODSs and Related Chemicals

mixing ratio profiles of hydrogen chloride (HCl) and chlorine nitrate (ClONO2), the two main inorganic reservoir 

chemicals for Cl in the stratosphere. Total column measurements from ground-based Fourier transform infrared 

(FTIR) instrumentation above the Jungfraujoch and Lauder show changes in inorganic Cl compound abundances 

that are fairly consistent with the amounts and changes 

in Cl being delivered to the stratosphere from long-lived 

source gases (Figure 1-17, Table 1-13) (update to Mahieu 

et al., 2004 and Rinsland et al., 2003). These findings are 

consistent with only a fairly small contribution of VSLS 

to stratospheric Cl. The mean fraction of stratospheric Cl 

attributable to VSLS, 80 (40–130) ppt (see Section 1.3), 

is approximately 1.5–3%. The relative contribution of 

VSLS to reactive Cl, however, can be substantially higher 

in the lower stratosphere for VSLS that have local lifetimes shorter than the longer-lived ODSs.

Average rates of change observed by ground-based 

FTIR measurements of total stratospheric chlorine 

(HCl + ClONO2) since 1996 were −0.9 ± 0.2%/yr to 

−1.5 ± 0.2%/yr (Table 1-13). The faster declines are 

apparent in the column measurements above Kiruna, but 

this may reflect greater meteorological variability at polar sites than above lower latitude sites. Comparable or 

slightly slower rates of change (−0.6 ± 0.1 to −0.9 ± 0.2%/

yr) have been measured since 2004 for upper stratospheric 

HCl measured over much of the globe from ACE-FTS and 

the Aura Microwave Limb Sounder (MLS) instruments 

(see Figure 1-18 and Table 1-13). These observed trends 

are reasonably consistent with the observed changes in 

total tropospheric Cl since 1996, though uncertainties 

related to chemical partitioning, mixing processes, and 

time lags associated with transport make a direct comparison between these different measures difficult.

Satellite data for upper stratospheric and lower mesospheric HCl from two different instruments (MLS and 

the ACE-FTS) agree to within ~0.15 ppb in recent years. 

This small difference is not significant given the systematic uncertainties on the ACE-FTS and MLS HCl measurements of 0.15 to 0.3 ppb (Figure 1-18) (Froidevaux et al., 

2008; Mahieu et al., 2008). These satellite results agree 

to within ± 0.3 ppb (± 8%) with total Cl amounts derived 

from long-lived chlorinated source gases, once mixing and 

lag times associated with transport are considered (Figure 

1-18). The uncertainty in this measured stratospheric Cl 

burden is large relative to the expected chlorine contributions from shorter-lived source gases and product gases 

(80, 40–130 ppt), and so these stratospheric data do not 

provide additional constraints to Cl contributions from 

chlorinated VSLS gases. HCl is estimated to account for 

>95% of total stratospheric Cl at altitudes above ~50 km.

Though the Cl contained in HCl and ClONO2 is 

not directly involved in the chemical reactions depleting 

stratospheric ozone, their summed abundance in the upper 

stratosphere is approximately proportional to the amount 

of reactive Cl present in the stratosphere. Chlorine monoxide (ClO), however, is directly involved in reactions that 

deplete stratospheric ozone. Updated measurements of 

upper stratospheric ClO (near 4 hPa) show an overall trend 

that is consistent with the declines observed for HCl and 

ClONO2 (Figure 1-19) (update of Solomon et al., 2006). 

Short-term variability is observed, however, suggesting 

that other factors influence the proportion of stratospheric 

Cl present in reactive (e.g., ClO) versus unreactive (e.g.,

HCl and ClONO2) forms, as discussed previously (e.g., 

Siskind et al., 1998; Froidevaux et al., 2000). Such influences also imply that measured ClO trends are not directly comparable to changes in total stratospheric chlorine. 

Altitude (km)

Year

ppt Cl Figure 1-16. The evolution of total 

chlorine abundance (ppt) in the 

stratosphere, calculated based on 

the tropospheric chlorine time series 

shown in Figure 1-15, and including 

a constant contribution of 80 ppt of 

chlorine from VSLS species (consistent with 40–130 ppt estimated 

in Section 1.3). Typical profiles of 

mean age in the stratosphere (derived from observations of SF6 and 

CO2) and a parameterization of the 

age spectrum are used for the calculation of the effects of mixing and 

transport in the stratosphere (updated from Engel et al., 2002).
1.66

Chapter 1

1.4.2 Bromine in the Troposphere 

and Stratosphere

1.4.2.1 Tropospheric Bromine Changes

Bromine in the troposphere from halons and methyl bromide continued to decline during 2005–2008. By 

2008, 15.7 ± 0.2 ppt of tropospheric Br was measured at 

Earth’s surface from these long-lived chemicals by two 

independent networks (annual mean). These results indicate that total Br from these substances has declined by 

about 1 ppt in the troposphere from peak levels measured 

in 1998 (Figure 1-20) (Montzka et al., 2003), similar to 

that expected in the A1 scenario of the last Assessment 

(Daniel and Velders et al., 2007). This decline has been 

Given observed changes in methane, a compound that affects this proportion, the measured decline of −0.9 ± 0.2%/

yr in ClO for 1995–2007 (Table 1-13) is consistent with a 

somewhat slower rate of change in total chlorine (about 

−0.7%/yr) (see Solomon et al., 2006).

New observations of chlorofluorocarbonyl (COFCl) 

have been reported for the TTL and in the stratosphere (Fu 

et al., 2009; Rinsland et al., 2007). The contribution of 

this compound to total Cl above 35 km is expected to be 

small (<10 ppt) and accounts for only a small portion of 

total inorganic Cl in the upper stratosphere. Furthermore, 

COFCl is believed to result almost entirely from degradation of long-lived compounds containing Cl and F in the 

stratosphere, and so its abundance is not included as an 

additional term in sums of total organic Cl from long- and 

short-lived source gases.

0

1

2

3

4

5

6

1984.0 1989.0

0

1

2

3

4

5

6

Inorganic chlorine above

Jungfraujoch (46.5°N)

Inorganic chlorine above

Lauder (45°S)

June-to-November 

monthly means

Year-round monthly means

Cly

HCl

ClONO2

ClONO2

HCl

Cly

1994.0 1999.0 2004.0 2009.0

Total Column Abundance (x10 molec./cm ) 15 2

Calendar Year

Figure 1-17. Time series of monthly-mean total column HCl (red circles) and ClONO2 (green triangles) abundance (molecules per square centimeter), as measured above the Jungfraujoch (46.5°N) and Lauder (45°S) 

(updated from Mahieu et al., 2004 and Rinsland et al., 2003). Cly column estimates (blue triangles) are the 

sum of HCl and ClONO2 columns. Jungfraujoch data are shown only for June to November of each year; results from all months are displayed for Lauder. Fits to the data sets are given by the black curves. The orange 

curves and corresponding fits represent Cly (HCl + ClONO2) from the University of Leeds 2-D model, as in the 

previous Assessment report (Clerbaux and Cunnold et al., 2007).
1.67

ODSs and Related Chemicals

primarily a result of decreasing mixing ratios of CH3Br 

in response to reduced industrial production and emission 

after 1998. Though Br from halons increased throughout 

most of the 1998–2008 period, during 2005–2008 Br from 

halons peaked in air at Earth’s surface at 8.2–8.5 ppt and 

was no longer increasing (see Table 1-1). Surface-based 

data suggest that global surface Br from halons may have 

decreased slightly from 2007 to 2008 (at −0.3 to −0.5%/

yr).

1.4.2.2 Stratospheric Bromine Changes

New results from ongoing ground-based DOAS 

(Differential Optical Absorption Spectroscopy) measurements of stratospheric BrO at both Lauder (45°S) and 

Harestua (60°N) show changes over time in inorganic 

bromine (Bry) that are highly consistent with the observed 

tropospheric Br changes measured for halons and CH3Br 

(Hendrick et al., 2008) (Figure 1-21). Updated stratospheric balloon measurements (Dorf et al., 2006) (Figure 

1-21) do not provide as tight a constraint on Bry changes 

since 2000, but the available results are consistent with 

the tropospheric trends measured for Br from long-lived 

gases and the UV-visible measurements of Hendrick et 

al. (2008). Despite the significant contribution of noncontrolled, very short-lived Br source gases to inorganic 

Bry in the stratosphere (Section 1.3.3), temporal changes 

in stratospheric inorganic Br abundance appear to follow 

the measured changes in tropospheric Br from controlled 

substances (i.e., CH3Br and the halons) (Figure 1-21).

Table 1-14 is an update from the previous Assessment (Table 2-8 in Law and Sturges et al., 2007) and lists 

estimates of Bry from VSLS (Bry

VSLS) and their uncertainty 

ranges for 12 studies. The contribution of VSLS to Bry is 

derived in these studies by calculating Bry from measured 

BrO using a photochemical model and then subtracting the 

contribution of long-lived gases (halons and CH3Br). Seven of these studies were discussed previously, though one 

preliminary study in the previous Assessment is no longer 

considered. The estimate based on an analysis of MLS 

measurements of BrO by Livesey et al. (2006) shown in 

the previous Assessment has been updated with a newer 

study of the same data set by Kovalenko et al. (2007), who 

infer a VSLS contribution of 6.5 ± 5.5 ppt to stratospheric 

Table 1-13. Measured chlorine-related changes in the upper atmosphere.

Instrumentation

Technique / Location 

Species and

Altitude Region Time Period 

Rate of Change

(percent/yr)

(2-sigma error) 

ground-based microwave 1

Hawaii 

ClO

upper stratosphere

1995–2007 −0.9 (0.2)

spaceborne ACE-FTS 2

IR solar occultation 

HCl

upper stratosphere

Jan. 2004–Sept. 2009

 

−0.9 (0.2)

spaceborne Aura MLS 3

Microwave emission 

HCl

upper stratosphere

Aug. 2004–Jan. 2010

 

−0.6 (0.1)

ground-based FTIR

Jungfraujoch 4

Kirunawwq 5

HCl column

1996–2009

1996–2009 

−0.9 (0.1)

−0.8 (0.2)

ground-based FTIR

Jungfraujoch 4

Kiruna 5

ClONO2 column

1996–2009

1996–2009 

−0.9 (0.3)

−2.9 (0.4)

ground-based FTIR

Jungfraujoch 4

Kiruna 5

HCl + ClONO2 column

1996–2009

1996–2009

−0.9 (0.2)

−1.5 (0.2) 

1. Ground-based microwave ClO data (Hawaii, 20°N): Updated data and results following Solomon et al. (2006). Rate of change (%/yr) is referenced 

to 1995 (January). See Figure 1-19.

2. ACE-FTS HCl data (60°S–60°N average): Updated data and results following Froidevaux et al. (2006). Rate of change (%/yr) is referenced to the 

average abundance during the data time period. See Figure 1-18.

3. Aura MLS HCl data (60°S–60°N average): Updated data and results following Froidevaux et al. (2006). Rate of change (%/yr) is referenced to the 

average abundance during the data time period. See Figure 1-18.

4. Ground-based FTIR column data (Jungfraujoch, 46.5°N): Updated data and results following Mahieu et al. (2004) and Clerbaux and Cunnold et 

al. (2007). Rates of change (%/yr) are referenced to 1996. See Figure 1-17.

5. Ground-based FTIR column data (Kiruna, 67.8°N): Updated data and results following Mikuteit (2008). Rates of change (%/yr) are referenced to 

1996.
1.68

Chapter 1

Figure 1-18. Time series of near-global HCl abundances (ppb) at about 53 km (near 0.5 hPa) (updated 

from Froidevaux et al. (2006) and the previous 

Assessment report (Clerbaux and Cunnold et al., 

2007)). The satellite data sets have been updated 

for ACE-FTS (version 2.2 data, quarterly averages 

for 60°S–60°N, green dots) and for Aura MLS (version 2.2 monthly averages for same latitude range, 

red dots); both these data sets are for averages over 

0.46 and 0.68 hPa (about 50 to 54 km). The solid 

curves are expectations for HCl (total Cl × 0.96) near 

55 km based on the ground-based total tropospheric 

chlorine, and include a range of air-age spectrum 

widths (blue, black, and orange curves are for widths 

of 0, 2, and 4 years). The MLS data are interrupted 

after early 2006, because of a lifetime issue that 

affects the main HCl band’s retrievals so it has been 

turned on only sparingly since; the MLS points in 

early 2009 and 2010 represent averages from 3–4 

days in early January of each year. The Halogen 

Occultation Experiment (HALOE) data set (monthly 

averages, purple) ended in November 2005 and no 

updated results are available from the Atmospheric 

Trace Molecule Spectroscopy (ATMOS) instrument. 

Error bars (2-sigma estimates of systematic uncertainties) are given for the various time series and are 

shown typically only for the first and last points in the 

series (for clarity).

1985 1990 1995 2000 2005 2010

2.5

3.0

3.5

4.0

HCl (ppb)

MLS

HALOE

ATMOS

ACE

Year

1995 2000 2005

Year

0.35

0.40

0.45

0.50

0.55

0.60

ClO at 4 hPa (ppb)

Linear Fit after 1995

Typical ±1σ precision of

"weekly" averages

Figure 1-19. Time series 

of ClO abundances (ppb) 

above Hawaii (at 35–39 

km or about 4 hPa) derived from ground-based 

millimeter-wave measurements since 1992 

(update of Solomon et 

al., 2006). Black and 

red symbols are “weekly” 

and “yearly” averages 

of deseasonalized data, 

and error bars typically 

show two standard deviations about the means 

(see above reference). 

Dashed line gives a linear 

fit to the observations after 1995 (see also Table 

1-13).
1.69

ODSs and Related Chemicals

bromine. Two of the other four new estimates introduced 

in Table 1-14 are based on ground-based DOAS BrO measurements (Theys et al., 2007 and Hendrick et al., 2008) 

and suggest that a contribution of 7 ± 4 ppt and 6 ± 4 ppt 

of bromine, respectively, is needed from VSLS to account 

for the total measured bromine loading. McLinden et al. 

(2010) calculate a contribution from VSLS of 5 ± 5 ppt 

based on stratospheric BrO profiles measured by OSIRIS 

(Optical Spectrograph and InfraRed Imager System) during 2001–2008. Salawitch et al. (2010) inferred a value 

for Bry

VSLS of 7 (5 to 10) ppt from OMI (Ozone Monitoring Instrument) measurements of total column BrO and 

tropospheric partial column BrO calculated from aircraft 

profiles (Neuman et al., 2010) obtained during Arctic 

spring.

These new results support the conclusion, made in 

the last Assessment (Law and Sturges et al., 2007), that 

substantially more Bry is measured in the stratosphere 

than can be accounted for by CH3Br and the halons alone; 

the new “ensemble” value reported in Table 1-14 is 1 ppt 

higher than the 5 ppt previously derived, while the range 

has stayed the same with 3 to 8 ppt. In other words, the 

majority of studies available to date continue to suggest a 

non-zero, positive value for Bry

VSLS of 6 (3–8) ppt, which 

is large enough to affect ozone photochemistry in the lower middle stratosphere (see Section 1.3.4). Only 2 out of 

the 12 studies reported in Table 1-14 include a zero contribution from Bry

VSLS in their uncertainty range. Taking this 

ensemble value for Bry

VSLS of 6 (3–8) ppt derived from 

consideration of all available stratospheric Bry estimates, 

and adding tropospheric Br from CH3Br and the halons 

from 2003–2004 in order to account for lags due to transport, a 2008 total stratospheric bromine mixing ratio of 

22.5 (19.5–24.5) ppt can be derived. This result is comparable, given the range in results reported by different techniques, to the Bry abundance derived from stratospheric 

measurements of BrO made in 2008. Annual mean Bry

derived from the NH Harestua site (60°N) during 2008 averaged 22.5 ± 4 ppt (an update of Hendrick et al., 2008), 

and 19.5 ± 2.5 ppt Bry was derived from a single balloon 

profile made on 27 June 2008 in the tropics (5°S, 43°W; 

update of Dorf et al., 2006) (Figure 1-21).

These results imply that slightly less than half (40–

45%) of the 2008 stratospheric Bry burden is derived from 

controlled uses of halons (~8.3 ppt Br) and methyl bromide (0.4–0.9 ppt Br). As a result, more than half of the 

2008 stratospheric Bry burden is accounted for by natural 

sources of methyl bromide and other bromocarbons, and 

from quarantine and pre-shipment uses of methyl bromide 

not controlled by the Montreal Protocol. The total contribution of VSLS to stratospheric Br is estimated to be 

1–8 ppt Br from measurements of VSLS source gases and 

modeling of their chemistry and transport (Table 1-9; Section 1.3.3.3). This magnitude is derived from the observed 

abundances of brominated source gases (SG) at the tropopause plus the modeled contributions of product gases 

(PG) in this region of the atmosphere (Hossaini et al., 

2010; Aschmann et al., 2009; Liang et al., 2010).

The potential difference between Bry from SGs 

and Bry from BrO is also influenced by uncertainties in 

5

7

9

11

13

15

17

19

1960 1970 1980 1990 2000 2010

Ab, (WMO, 2003)

A1, (WMO, 2007)

UEA

NOAA/ESRL

AGAGE

CH3Br only (NOAA)

CH3Br only (AGAGE)

Bromine from C

H3Br and halons

 at Earth's surface (ppt)

Year

Figure 1-20. Tropospheric Br 

(ppt) from CH3Br and halons over 

time. Halons measured by AGAGE 

(green line; Prinn et al., 2000; Sturrock et al., 2002), UEA (red line; 

Reeves et al., 2005), and NOAA 

(blue line) are added to CH3Br data 

from NOAA (dashed black line; 

Montzka et al., 2003) or AGAGE 

(plus symbols; Simmonds et al., 

2004). Ab (dot-dot-dashed line indicated as WMO, 2003 (Montzka 

and Fraser et al., 2003)) and A1 

(solid black line indicated as WMO, 

2007 (Clerbaux and Cunnold et 

al., 2007)) baseline scenarios are 

shown for comparison. Updated 

considerations of preindustrial mixing ratios in the current Assessment cause the differences in earlier years (see Section 1.2.1.6 and 

Chapter 5).
1.70

Chapter 1

Year of stratospheric entry

Total stratospheric bromine (ppt)

Bry from DOAS BrO:

Balloonborne profiles

Balloonborne Langley observations

Ground-based profiles at Harestua

Ground-based profiles at Lauder

1980 1985 1990 1995 2000 2005 2010

26

6

10

16

18

22

30

8

12

14

20

24

28

Figure 1-21. Changes in total stratospheric Bry (ppt) derived from balloonborne BrO observations (squares) 

(update of Dorf et al., 2006) and annual mean mixing ratios calculated from ground-based UV-visible measurements of stratospheric BrO made at Harestua (60°N) and Lauder (45°S) (filled and open orange triangles, 

respectively) (adapted from Hendrick et al., 2008). These stratospheric trends are compared to trends in measured bromine (ppt) at Earth’s surface with additional constant amounts of Bry added (thin lines). Lines show 

global tropospheric bromine from methyl bromide as measured in ambient air and firn air with no correction 

for tropospheric CH3Br loss (dark blue line; Butler et al. (1999) through 1998 including a changing interhemispheric gradient—see Chapter 5; Montzka et al. (2003) thereafter); global tropospheric bromine from the sum 

of methyl bromide plus halons (purple line; Butler et al. (1999) and Fraser et al. (1999) through 1995; Montzka 

et al. (2003) thereafter); and bromine from CH3Br, halons, plus additional constant amounts of 3, 5, and 7 ppt 

Br (thin blue lines). Squares show total inorganic bromine derived from stratospheric measurements of BrO 

and photochemical modeling that accounts for BrO/Bry partitioning from slopes of Langley BrO observations 

above balloon float altitude (filled squares) and lowermost stratospheric BrO measurements (open squares). 

For the balloonborne observations, bold/faint error bars correspond to the precision/accuracy of the estimates, 

respectively. For the ground-based measurements (triangles), the error bars correspond to the total uncertainties in the Bry estimates. For stratospheric data, the date corresponds to the time when the air was last in the 

troposphere, i.e., sampling date minus estimated mean age of the stratospheric air parcel. For tropospheric 

data, the date corresponds to the sampling time, i.e., no corrections are applied for the time required to transport 

air from Earth’s surface to the tropopause. Preindustrial levels were 5.8 ± 0.3 ppt for CH3Br (Saltzman et al., 

2004) and 0 ppt for the halons (Reeves et al., 2005). This figure is an update to Figure 2-3 from the previous 

Assessment (Law and Sturges et al., 2007).
1.71

ODSs and Related Chemicals

Table 1-14. Estimates of inorganic bromine from very short-lived substances (Bry

VSLS) contribution to 

stratospheric bromine derived from BrO measurements. Update of Table 2-8 from the previous Assessment (Law and Sturges et al., 2007), extended with new results.

Data Source Bry

VSLS

Central Value (ppt)

Bry

VSLS

Range (ppt) Reference

Ground-based BrO

11 sites, 78°S–79°N

5 1–9 a Sinnhuber et al. (2002)

Ground-based BrO

Lauder, New Zealand, 45°S

6 3–9 Schofield et al. (2004)

Ground-based BrO

Arrival Heights, Antarctica, 78°S

6 3–9 Schofield et al. (2006)

DOAS balloon BrO profiles

5°S–68°N, 0–35 km

4.1 1.6–6.6 Dorf et al. (2006)

Aircraft & balloon BrO profiles,

22°S–35°N, 17–32 km and

GOME satellite column BrO,

60°S–60°N

7 4–10 a Salawitch et al. (2005)

SCIAMACHY satellite BrO profiles

80°S–80°N, 15–28 km

8.4 6.4–10.4 Sioris et al. (2006)

SCIAMACHY satellite BrO profiles

80°S–80°N, 15–28 km 

3 0–6 Sinnhuber et al. (2005)

Ground-based BrO

Harestua (60°N) and Lauder (45°S) 

6 2–10 Hendrick et al. (2008)

Ground-based BrO

Reunion Island (21°S, 56°E) 

7 3–11 b Theys et al. (2007)

MLS satellite BrO profiles

55°S–55°N, 10–4.6 hPa

6.5 1–12 Kovalenko et al. (2007)

OSIRIS satellite BrO profiles

80°S–80°N, 20–34 km

5 0–10 McLinden et al. (2010)

OMI total column BrO and aircraft tropospheric BrO profiles,

Arctic spring

7 c 5–10 Salawitch et al. (2010)

Ensemble 6 (3–8) d

DOAS, Differential Optical Absorption Spectroscopy; GOME, Global Ozone Monitoring Experiment; SCIAMACHY, 

Scanning Imaging Absorption Spectrometer for Atmospheric Cartography; MLS, Microwave Limb Sounder; OSIRIS, 

Optical Spectrograph and InfraRed Imager System; OMI, Ozone Monitoring Instrument.

___________

a Range estimated by previous Assessment (Law and Sturges et al., 2007), based on the uncertainty in stratospheric Bry inferred from BrO that was stated 

in the reference.

b Range estimated based on the Hendrick et al. (2008) study that uses the same method. c Salawitch et al. (2010) has a different treatment of CH2Br2 than the other studies listed here. Their estimate for Bry

VSLS would be slightly higher 

(~ ≤1 ppt) if CH2Br2 were treated in the same manner as in the other studies.

d Average and range of the central values of the 12 estimates of Bry

VSLS.
1.72

Chapter 1

absolute calibration scales for measurements of brominated organic SGs (Butler et al., 2010) and product gases. 

Furthermore, concentrations of VSLS sampled by a wholeair sampler represent only local influences for a small 

region over a short time in the tropical tropopause layer. 

In contrast, Bry from BrO is measured in an air parcel in 

which the Br content represents a weighted average of Br 

in air parcels with a distribution of mean ages and that has 

been influenced by stratospheric mixing processes.

In addition to observations of BrO, measurements 

of stratospheric bromine nitrate (BrONO2) have been 

analyzed from the Michelson Interferometer for Passive 

Atmospheric Sounding (MIPAS-E) instrument onboard 

the Envisat satellite for September 2002 and September 

2003 (Höpfner et al., 2009). The authors report that the 

maximum values of BrONO2 observed during night are 

always in the range of 20–25 ppt (with an uncertainty of 

20–25% depending upon altitude). Large uncertainties 

remain, however, in the absorption cross section and its 

temperature dependence, so reliable estimates of Bry from 

these measurements are not currently possible.

The contribution of tropospheric BrO to groundand satellite-based measurements of stratospheric Bry has 

been discussed in recent years. As in the last Ozone Assessment (Law and Sturges et al., 2007), it remains unclear 

whether BrO is distributed throughout the global troposphere with a background abundance of up to 3 ppt. Such 

a BrO background is supported by several studies (e.g., 

Richter et al., 2002; Van Roozendael et al., 2002; Theys et 

al., 2007), where comparisons of Global Ozone Monitoring Experiment (GOME) BrO vertical columns with model results, balloonborne observations, and ground-based 

measurements have led to the conclusion that a significant 

background of several ppt BrO is present in the free troposphere. However, other observations support much lower 

values that include a contribution of zero ppt Br (e.g., 

Schofield et al., 2004; Dorf et al., 2008). Tropospheric 

BrO levels of less than 1 ppt imply larger stratospheric 

BrO columns than can be accounted for by CH3Br and 

halons alone, and, as a result, an important contribution of 

VSLS to stratospheric Bry is necessary to explain the total 

column observations (e.g., Salawitch et al., 2005).

Similarly, two independent studies (also discussed 

in detail in the previous Assessment) show BrO profiles 

retrieved from Scanning Imaging Absorption Spectrometer for Atmospheric Cartography (SCIAMACHY) radiances that lead to different results. While Sioris et al. 

(2006) derive large BrO abundances in the tropical tropopause layer and lowermost stratosphere (LMS), consistent 

with Bry

VSLS of 8.4 ± 2 ppt, Sinnhuber et al. (2005) imply 

a much smaller value for Bry

VSLS of 3 ± 3 ppt. One cause 

for this discrepancy is that Sinnhuber et al. (2005) suggest the presence of 1 ± 0.5 ppt of tropospheric BrO while 

Sioris et al. (2006) suggest a much smaller level of global 

tropospheric BrO. Balloonborne DOAS observations of a 

BrO profile made in the tropics in 2005 found that BrO in 

the lower and middle troposphere is <1 ppt and compatible 

with zero within the uncertainties (Dorf et al., 2008). The 

authors derive a total contribution of 5.2 ± 2.5 ppt from 

brominated VSL SGs and inorganic PGs to stratospheric 

Bry and so suggest a Bry

VSLS magnitude in between the 

earlier results derived from SCIAMACHY.

The recent ARCTAS (Arctic Research of the Composition of the Troposphere from Aircraft and Satellites) 

and ARCPAC (Aerosol, Radiation, and Cloud Processes 

affecting Arctic Climate) campaigns included airborne in 

situ measurements of BrO and O3 (Neuman et al., 2010), 

placed in the footprint of satellite observations, to quantify 

the relative contribution of tropospheric and stratospheric 

partial columns to total column BrO and the relation 

between BrO hotspots (regions where column BrO is enhanced by 2 to 3 × 1013 radicals cm−2 relative to the zonal 

mean) and surface ozone depletion events. Salawitch et al. 

(2010) reported that aircraft in situ measurements of BrO 

and O3 near the surface often bear little relation to OMI BrO 

“hotspots.” The geographic location of numerous BrO hotspots was shown to be consistent with the location of high 

total column ozone and a low-altitude (~5 km) tropopause, 

suggesting a stratospheric origin to the BrO enhancements 

that implied, on average, Bry

VSLS equal to 7 (5–10) ppt. It 

has been noted that the largest source of uncertainty in deriving Bry from measurements of total column BrO is the 

rate constant of BrO + NO2 + M → BrNO3 + M (Sioris et 

al., 2006; Hendrick et al., 2008; Salawitch et al., 2010). 

Salawitch et al. (2010) suggest that the preponderance of 

prior observations of elevated BrO over Hudson Bay during spring may be related to a synoptic weather pattern 

known as the Hudson Bay low that is responsible for the 

very low-altitude tropopause (e.g., Liu and Moore, 2004). 

They stress that proper understanding of BrO “hotspots” 

requires accurate treatment of perturbations originating 

from both the stratosphere and troposphere, which generally has not been the focus of prior studies (i.e., many 

prior studies erroneously treat the stratospheric signal as 

spatially constant, ascribing all contributions to BrO hotspots as originating from below the tropopause).

In addition to the previously reported observations 

of BrO in the midlatitude marine boundary layer (Leser 

et al., 2003; Saiz-Lopez et al., 2004), BrO has now also 

been observed directly in the tropical and subtropical marine boundary layer. Mean daytime maxima of 2.5 ± 1.1 

ppt were reported for observations made from November 

2006 until June 2007 at Cape Verde (Read et al., 2008) 

and peak mixing ratios of 10.2 ± 3.7 ppt were measured 

during a ship cruise along the African coast during February 2007 (Martin et al., 2009). It is not clear to what extent this BrO influences Br abundances above the marine 

boundary layer.
1.73

ODSs and Related Chemicals

1.4.3 Iodine in the Upper Troposphere 

and Stratosphere

Updated results since the last Assessment continue 

to suggest that gaseous iodine-bearing compounds contribute very little iodine to the current lower stratosphere. 

Balloonborne solar occultation spectra of IO and OIO in 

the tropical upper troposphere/lower stratosphere (UT/LS) 

obtained during balloon flights in 2005 and 2008 confirm 

the low values of IO and OIO for the tropical upper troposphere and lower stratosphere (Butz et al., 2009). A 

photochemical model of these results yields corresponding upper limits for the total gaseous inorganic iodine burden (Iy) of 0.17 to 0.35 (+0.20/−0.08) ppt in the tropical 

upper troposphere (13.5 km to 16.5 km) and 0.09 to 0.16 

(+0.10/−0.04) ppt in the tropical lower stratosphere (16.5 

km to 21.0 km).

These observations do not preclude iodine reaching 

the lower stratosphere in chemical forms other than commonly considered in photochemical models, or in particulate matter. There is evidence that iodine that may have 

originated in the lower troposphere can be found in lower 

stratospheric aerosol (Murphy et al., 2007). But if present 

in the TTL or stratosphere in forms other than IO and OIO, 

the extent to which this iodine might become available for 

ozone-depleting reactions in other parts of the stratosphere 

is not known.

The inferred upper limits for IO and OIO concentrations in the lower stratosphere suggest that catalytic 

cycles involving iodine have only a minor contribution to 

total ozone loss, but it remains unclear whether reactive 

iodine contributes—possibly through coupling with chlorine or bromine—to the observed trend of declining ozone 

in the lower stratosphere (e.g., Butz et al., 2009).

1.4.4 Equivalent Effective Chlorine 

(EECl) and Equivalent Effective 

Stratospheric Chlorine (EESC)

Changes in the stratospheric burden of total inorganic halogen (Cl and Br) are estimated from measured 

tropospheric changes in ODS abundances with a number of 

different metrics such as equivalent effective stratospheric 

chlorine (EESC), EESC-Antarctica, and EESC-Midlatitudes; changes in total tropospheric halogen (Cl and Br) are 

addressed with the equivalent effective chlorine (EECl), 

and equivalent chlorine (ECl) metrics (Clerbaux and Cunnold et al., 2007). By accounting for transport times, the 

efficiency for different halogens to deplete ozone, mixing 

processes, and age-of-air-dependent ODS decomposition rates (i.e., fractional release values), inorganic halogen abundances and changes can be estimated with EESC 

for different regions of the stratosphere from tropospheric 

ODS measurements. As discussed in the last report and 

in recent publications, new approaches have been explored 

that enhance the usefulness of EESC for more accurately 

estimating halogen abundances and changes in specific 

stratospheric regions by considering that the extent of 

degradation of an ODS in the stratosphere is fairly well 

described by the mean stratospheric age of an air parcel 

(Newman et al., 2007). This new approach incorporates 

revised fractional release values for some compounds that 

are based primarily on observations (Schauffler et al., 

2003) but that are also consistent with models (Douglass et 

al., 2008; Daniel and Velders et al., 2007). New fractional 

releases have not been considered for HCFC-142b and 

HCFC-141b owing to large discrepancies between the values derived by observations (e.g., Schauffler et al., 2003) 

and those based on model calculations that have been used 

in past Ozone Assessments.

New methods have been devised to more readily 

convey the magnitude of decline in EESC over time by 

referencing the changes in EESC to peak levels observed 

in the mid-1990s and amounts inferred for 1980 (Hofmann 

and Montzka, 2009). Though ozone depletion was likely 

non-negligible in 1980 (see Chapter 2 and Chapter 5), the 

EESC value in 1980 is still used in this report as an important benchmark for ozone recovery, as in past reports.

Stratospheric air parcels in midlatitudes having 

a mean age of 3 years have substantially less inorganic 

halogen as EESC than an air parcel with a mean age of 5.5 

years typical of the Antarctic spring vortex (Figure 1-22, 

top panel) (Newman et al., 2007). Older air has larger 

EESC values because as air ages in the stratosphere, ODSs 

photochemically decompose so that a higher fraction of 

available halogen is present in inorganic forms.

As discussed in the last Assessment report (Daniel 

and Velders et al., 2007), calculating EESC in this way 

also demonstrates that EESC peaks at different times in 

different regions of the stratosphere. With the shorter 

transit times and higher relative contribution of rapidly declining short-lived ODSs in midlatitudes (e.g., CH3CCl3), 

EESC peaks earlier in midlatitudes than in polar regions 

(Figure 1-22, lower panel).

The date when the second derivative in EESC 

reaches a minimum also varies with stratospheric region. 

This date has been proposed as the time when a significant reduction in the rate of ozone decline might be expected, i.e., the “first stage” of ozone recovery (Yang 

et al., 2008). The second derivative in EESC reaches 

a minimum in early 1996 in midlatitudes (mean age = 

3 years and age-of-air spectrum width = 1.5 years) and 

in late 1997 in the Antarctic stratosphere (mean age = 

5.5 years and age-of-air spectrum width = 2.75 years) 

(Figure 1-22).

The similarity between global surface means 

derived for ODSs from independent surface sampling 
1.74

Chapter 1

networks (Table 1-1) suggests only small uncertainties 

in our estimate of EESC abundance and trend in recent 

years related to measurement errors. Larger uncertainties are associated with using tropospheric data to derive 

actual inorganic halogen abundances in different regions 

of the stratosphere. These additional uncertainties stem 

from an imperfect understanding of ODS decomposition 

rates in the stratosphere (fractional release factors), stratospheric mixing processes, and stratospheric air-age spectra 

(Newman et al., 2007).

Results from all three global surface sampling 

networks indicate that EESC and tropospheric EECl 

continued to decline during 2005–2008. By 2008, tropospheric EECl (weighted with fractional release factors 

appropriate for the midlatitude stratosphere) had dropped 

14% from its peak value measured in 1993–1994. Note 

that the magnitude of this decline is somewhat sensitive 

to the fractional release factors used. While a 14% drop 

is calculated with the updated fractional release factors 

discussed above, a drop in EECl of only 11% would be 

calculated with the fractional release factors used in the 

previous Assessment report.

Changes in stratospheric inorganic halogen abundance (EESC) through 2008 can also be estimated from 

tropospheric measurements of ODSs. By 2008, midlatitude EESC had decreased by about 11% from its peak 

value in 1997. This stratospheric decline is smaller than 

derived for tropospheric EECl because of the time lag associated with transporting air from the troposphere to the 

stratosphere. When referenced to EESC in 1980, this drop 

is 28% of the decrease required for EESC in midlatitudes 

to return to that benchmark level (Figure 1-22, bottom 

panel). These EESC declines calculated for midlatitudes 

are similar whether or not mixing processes are accounted 

for with an air-age spectrum.

An even smaller decline is derived for the inorganic halogen abundance in the Antarctic polar vortex. 

By 2008, polar EESC had decreased by about 5% from 

its peak value in 2002. The much smaller decline over 

Antarctica stems from less time having elapsed since peak 

Figure 1-22. Top panel: Equivalent effective 

stratospheric chlorine (EESC) (ppt) calculated 

for the midlatitude stratosphere from surface 

measurements (see Figure 1-1) and absolute 

fractional release values from Clerbaux and 

Cunnold et al. (2007) (black line) or with age-ofair-dependent fractional release values for the 

midlatitude stratosphere (red lines; mean age 

= 3 years) and the Antarctic springtime stratosphere (blue lines; mean age = 5.5 years). 

EESC calculated with stratospheric mixing 

processes included are shown as dashed 

colored lines for the different stratospheric regions (i.e., an air-age spectrum width equal to 

one-half the mean age; Newman et al., 2007). 

Different shades of the same colors represent 

EESC calculated with tropospheric halocarbon data from the different surface networks 

(NOAA and AGAGE, not always distinguishable from one another) without consideration 

of air-age spectra. Vertical dashed lines represent the date when the second derivative of 

EESC reaches a minimum (see text). Bottom 

panel: EESC derived for the midlatitude and 

polar stratospheric regions as a function of 

time plotted relative to peak abundances of 

1950 ppt for the midlatitude stratosphere and 

4150 ppt for the polar stratosphere. Percentages shown on right of this panel indicate the observed change in EESC relative to the change needed for 

EESC to return to its 1980 abundance (note that a significant portion of the 1980 EESC level is from natural 

emissions of CH3Cl and CH3Br). Colors and line styles represent the same quantities as in the top panel.

 0%

50%

60%

70%

80%

90%

100%

110%

EESC abundance 

(relative to the peak)

Midlatitude stratosphere

Polar stratosphere

-10%

-28%

1980 levels

500

1000

1500

2000

2500

3000

3500

4000

4500

EESC (ppt)

1980 1985 1990 1995 2000 2005 2010

% return to

1980 level

by the end 

 of 2008

Date of peak in the

troposphere

Year 
1.75

ODSs and Related Chemicals

abundances were observed in this older air and the dampening of changes by mixing air with a wider age spectrum 

across the maximum in EESC. This drop is approximately 

10% of the decrease required for EESC in polar regions to 

return to the 1980 benchmark level (Figure 1-22, bottom 

panel). The EESC declines calculated for polar regions 

are more sensitive to mixing processes than in midlatitudes (see Figure). Slightly smaller declines are calculated 

for polar regions when an air-age-spectrum width of 2.75 

years (one-half of the mean age of 5.5 years) is considered 

(a total decline of ~4% compared to 5% when mixing processes are not included in the EESC calculation).

In previous Assessment reports, it was noted that 

declines in the shorter-lived gases CH3CCl3 and CH3Br 

were the main reason for the observed declines in EECl 

and EESC. During the past four years, however, no single 

chemical class has dominated the decline in the total combined abundance of ozone-depleting halogen in the troposphere. From 2005 through 2008, the long-lived CFCs 

(−17 ppt EESC) contributed similarly to the EESC decline 

as did the short-lived CH3CCl3 (−20 ppt EESC) and CH3Br 

(−24 ppt EESC). Other compounds and compound classes 

contributed less to this decline (CCl4: −10 ppt EESC; halons: −4 ppt EESC), and HCFCs (+6 ppt EESC) added to 

this halogen burden over this period.

1.4.5 Fluorine in the Troposphere 

and Stratosphere

In contrast to the declines observed for tropospheric Cl and Br in recent years, the tropospheric abundance 

of fluorine (F) increased at a mean annual rate of 1.6 ± 

0.1%/yr (40 ± 4 ppt/yr) since 1996 (fluorine from CFCs 

-11, -12, -113, -114, -115; HCFCs -22, -141b, -142b, 

-124; halons -1211, -1301, -2402; HFCs -23, -152a, 

-134a, -143a, -125; PFCs -14, -116, -218; and SF6). This 

rate is substantially less than the annual increases of 60–

100 ppt/yr (annually 5.9 ± 0.3%/yr) observed for tropospheric F during the 1980s as CFC abundances were rapidly increasing. Many replacement compounds (HCFCs, 

HFCs, and PFCs) and other gases contributing F to the 

atmosphere degrade very slowly in the stratosphere and, 

therefore, do not affect stratospheric hydrogen fluoride 

(HF) changes as much as shorter-lived gases. Tropospheric F from compounds having stratospheric lifetimes 

<100 yrs (CFCs, halons, HCFC-141b, and HFC-152a; 

see Table 1-3) peaked in 2001 and was decreasing slowly during 2005–2008 (trend of −0.4 ± 0.1%/yr). Stratospheric changes in HF are likely to be between these two 

ranges (−0.4%/yr to +1.6%/yr). 

Stratospheric column FTIR measurements of HF, 

carbonyl fluoride (COF2), and Fy

*

 (where Fy

*

 = HF + 

2×COF2) also suggest a decrease in accumulation rate 

for atmospheric F beginning in 1999 (Figure 1-23). During 2005–2008 the rate of increase was 0.4%/yr, which is 

smaller than observed during the late 1980s–early 1990s. 

This change in rate is consistent with tropospheric changes 

in F-containing gases propagating to the stratosphere and 

a lag associated with transport. While the abundance and 

rates of change observed for stratospheric Fy

*

 are fairly 

well calculated with a 2-D model (Chipperfield et al., 

1997) in which tropospheric changes for a subset of Fcontaining gases were considered, the calculated F abundance is likely a lower limit for F as some less abundant 

fluorine source gases were not included in the model. 

Data from other FTIR sites have also shown continued 

increases in HF column abundances. The measurements 

from Kiruna, Sweden, for example, indicate that an average increase in HF column of 1%/yr (± 0.3%/yr) occurred 

during 1996–2008 (relative to 2000; updates to Mikuteit 

(2008)). The rates of increase in stratospheric HF columns 

derived from these ground-based observations are within 

the expected range derived from surface observations 

(−0.4%/yr to +1.6%/yr).

1.5 CHANGES IN OTHER TRACE GASES THAT 

INFLUENCE OZONE AND CLIMATE

In addition to the Cl- and Br-containing ODSs, 

other trace substances can affect stratospheric ozone. The 

most important substances in this group are the greenhouse gases. As greenhouse gases have been recently 

assessed by the Working Group I of the IPCC Fourth 

Assessment Report (AR4) (Forster et al., 2007; Denman 

et al., 2007) only updates on recent trends of mole fractions (Table 1-15), and information on sources and sinks 

are provided. Furthermore, in Figure 1-24 the development of the adiative forcing of the greenhouse gases discussed in this section is shown. In Section 1.5.1, CH4, 

N2O, and sulfur compounds are discussed, which, apart 

from their direct influence on the abundance of the stratospheric ozone, also have an effect on the radiative forcing. A special focus in this section is the renewed increase 

of atmospheric CH4 mixing ratios (Rigby et al., 2008; 

Dlugokencky et al., 2009) and the re-evaluated influence 

of N2O on stratospheric ozone (Ravishankara et al., 2009). 

Sulfur trace gases and SO2 as their main degradation product can potentially reach the stratosphere and can influence 

stratospheric ozone by enhancing its degradation through 

heterogeneous reactions. In Section 1.5.2, fluorinated 

greenhouse gases, which only indirectly influence stratospheric ozone via the greenhouse effect, are discussed.

The influence of rockets on stratospheric ozone will 

be discussed in Section 1.5.3, as their emissions could be 

highly relevant because they are emitted either in or near 

the stratospheric ozone layer.
1.76

Chapter 1

1.5.1 Changes in Radiatively Active Trace 

Gases that Directly Influence Ozone

In this section substances that are not controlled by 

the Montreal Protocol but that nevertheless have an influence on stratospheric ozone are discussed.

1.5.1.1 Methane (CH4)

Apart from its well-established influence on radiative forcing, methane (CH4) is both a sink of reactive chlorine and a source of water vapor in the stratosphere and, 

therefore, influences the availability of inorganic halogen 

for depleting stratospheric ozone (see Section 1.4.1). Furthermore, CH4 is a significant sink of tropospheric OH 

radicals and, hence, changes in its abundance can lead to 

changes in the lifetimes of ozone-depleting substances 

removed by OH (e.g., CH3CCl3, HCFCs, VSLS).

A detailed description of the global methane budget 

and its uncertainties through 2005 is given in the IPCC 

AR4 (Forster et al., 2007). They reported that the global 

average CH4 mole fraction in 2005 was 1774 ppb, which 

far exceeds the natural range during the last 650,000 years, 

and that the observed increase since preindustrial times is 

very likely due to anthropogenic activities (agriculture, 

fossil fuel use, etc.). Here we focus on important results 

published afterwards and any similarities and differences 

between these recent and earlier studies.

For the IPCC A2 scenario, Portmann and Solomon 

(2007) compute methane-induced ozone increases over 

the 21st century of 2–12% below 20 km, and 0–4% between 20 and 40 km. On the other hand, they compute 

ozone decreases of 0–12% between 40 and 60 km. By 

2008, the radiative forcing of CH4 arising from mixing ratio changes since 1750 had reached about 0.5 W/m2

, second only to CO2. Indirect effects of atmospheric methane 

related to changes in stratospheric H2O and tropospheric 

ozone suggest that the present-day net radiative forcing of 

CH4 could be as large as 0.7 W/m2 (Forster et al., 2007).

After eight years of minimal net change, the mole 

fractions of CH4 began to increase in 2007 in both hemispheres (Rigby et al., 2008; Dluogokencky et al., 2009). 

The growth rate was about 0.9 ± 3.3 ppb/yr from 1998–

2006, while the global growth rate for 2007–2008 averaged 

5.9 ppb/yr (NOAA) or 8.4 ppb/yr (AGAGE) (Table 1-15). 

The growth rate anomalies during 2006–2008 have been 

similar in magnitude to those observed in some years since 

1990. The inverse analysis of Rigby et al. (2008) suggests 

that the renewed increase in growth rate was attributable 

either to increasing tropical and high latitude emissions or 

to a smaller high-latitude emissions increase along with a 

few percent low-latitude OH decrease (or to some combination of the two). A second analysis of spatial gradients 

Year

1983.0 1986.0 1989.0 1992.0 1995.0 1998.0 2001.0 2004.0 2007.0 2010.0

Total colum abundance (x10 molec./cm ) 15 2

0.0

0.5

1.0

1.5

2.0

2.5 Inorganic fluorine above

Jungfraujoch

June to November monthly means

Non parametric LS fit

HF

2 x COF2

Fy

*

2D-Model Fy

*

Figure 1-23. Time series of monthlymean total column HF (red circles) 

and COF2 (green triangles) (molecules per square centimeter), as derived from the Jungfraujoch (46.5°N) 

database, updated from the previous Assessment (Clerbaux and 

Cunnold et al., 2007). Duchatelet et 

al. (2009) have provided an updated 

analysis for the COF2 results shown 

here, using a multispectral, multiwindow approach. Data values are 

limited to June to November months 

to reduce variability. Fits to the data 

sets are given by the black curves. 

Fy

*

 estimates (blue triangles) are 

calculated from HF + 2×COF2. Also 

shown are column Fy

*

 estimates 

from a 2-D model (based on Chipperfield et al., 1997) shown by the 

orange curve, with corresponding fit, 

that are derived from surface observations of CFC-11, CFC-12, CFC113, and HCFC-22.
1.77

ODSs and Related Chemicals

Table 1-15. Mole fractions of CO2, CH4, N2O, SF6, SO2F2, COS, and selected HFCs and PFCs.

Mole Fraction Annual Change in 

Mole Fraction

2005 2006 2007 2008 2005/06 2006/07 2007/08

CO2 [ppm] N 378.8 381.0 382.7 384.8 2.2 1.7 2.1

CH4 [ppb] N 1774.7 1775.4 1781.7 1787.6 0.7 6.3 5.9

CH4 [ppb] A 1774.2 1774.6 1780.8 1789.2 0.4 6.2 8.4

N2O [ppb] N 319.0 319.8 320.5 321.5 0.8 0.7 1.0

N2O [ppb] A 319.2 319.9 320.6 321.6 0.7 0.7 1.0

SF6 [ppt] N 5.6 5.9 6.2 6.4 0.3 0.3 0.2

SF6 [ppt] A 5.6 5.9 6.2 6.4 0.3 0.3 0.2

HFC-134a [ppt] N 34.4 38.8 43.2 47.6 4.4 4.4 4.4

HFC-134a [ppt] A 34.6 38.9 43.3 48.2 4.3 4.4 4.9

HFC-23 [ppt] A * 19.0 20.0 21.0 21.8 1.1 1.0 0.8

HFC-152a [ppt] A 4.1 4.5 5.3 5.9 0.4 0.8 0.6

HFC-143a [ppt] A 5.8 6.6 7.5 8.5 0.8 0.9 1.0

HFC-32 [ppt] A 1.3 1.6 2.1 2.7 0.3 0.5 0.6

HFC-125 [ppt] A 3.9 4.5 5.2 6.1 0.6 0.7 0.9

HFC-365mfc [ppt] A 0.1 0.3 0.4 0.4 0.2 0.1 0.1

HFC-245fa [ppt] ** 0.4 0.6 0.7 1.0 0.2 0.1 0.4

PFC-14 [ppt] A 75.1 75.7 76.4 77.1 0.6 0.7 0.7

PFC-116 [ppt] A 3.7 3.8 3.8 3.9 0.1 0.0 0.1

PFC-218 [ppt] A 0.4 0.4 0.5 0.5 0.0 0.1 0.0

SO2F2 [ppt] A 1.35 1.42 1.47 1.51 0.07 0.05 0.04

COS [ppt] N 488 491 494 491 2.2 2.9 −2.1

Data are global surface means. 

These observations are updated from the following sources: Conway et al. (1994), Dlugokencky et al. (2009), Geller et 

al. (1997), Greally et al. (2005, 2007), Hall et al. (2007), Miller et al. (2010), Montzka et al. (1996, 2007), Mühle et 

al. (2009, 2010), O’Doherty et al. (2004, 2009), Prinn et al. (2000), Rigby et al. (2008), Stemmler et al. (2007), and 

Vollmer et al. (2006).

Annual changes in mole fraction are drived from the difference between year x and x-1.

N denotes data from NOAA.

A denotes data from AGAGE.

________

* Global averaged mixing ratios for HFC-23 before 2007 have been modeled using archived air data from the SH (Miller et al., 2010). ** Data are an average of measurements from the Jungfraujoch, Switzerland (47°N) and Cape Grim, Australia (40°S) (updated from Vollmer et al., 2006).
1.78

Chapter 1

and Arctic isotopic signals by Dlugokencky et al. (2009) 

suggests primarily an increased wetland source in both 

the high latitudes and tropics for 2007, while the 2008 increase was predominantly seen at tropical latitudes. They 

argued for little influence from OH variations. Furthermore, an increase in interhemispheric exchange during the 

cool phase of the El Niño-Southern Oscillation (ENSO) 

(consistent with observed changes in the SF6 mean interhemispheric difference), could have contributed, in part, to 

changes in Southern Hemisphere growth rates. Bousquet 

et al. (2006) suggested that the relatively stable period for 

CH4 mixing ratios between 1999 and 2006 was a fortuitous combination of a decreasing wetland source masking 

an increasing anthropogenic energy-related source, with 

the chemical loss of CH4 due to OH playing a potential 

role in the observed atmospheric variability. In compariYear Year

Figure 1-24. Left panels: The evolution of radiative forcings from the sum of the major greenhouse gases (CO2, 

CH4, N2O), the ODSs (CFCs, HCFCs, halons, CH3CCl3, CCl4), and the Kyoto Protocol synthetic gases (HFCs, 

PFCs, SF6) in W/m2. Right panels: The upper panel shows the evolution of radiative forcings (note change in 

scale) from individual HCFCs and their sum, and from the sum of Kyoto Protocol synthetic gases (HFCs, PFCs, 

and SF6), which appear individually or grouped together in the lower right panel (in mW/m2). Forcings are calculated from background mixing ratios in Table 1-1 and Table 1-15 and radiative efficiencies given in Chapter 5.
1.79

ODSs and Related Chemicals

son, Chen and Prinn (2006) evaluated growth rates and 

distributions of CH4 during 1996–2001 relative to earlier 

years and found decreased energy-related emissions and 

increased emissions from rice, with the 1998 anomaly due 

to increased global wetland and wildfire emissions. Chen 

and Prinn (2006) note that their inferred increased rice 

emissions (about 25 Tg CH4/yr) could also be attributed to 

surrounding non-rice wetland emissions.

While the possibility of very large methane emissions from plants (as proposed by Keppler et al., 2006) has 

been ruled out (Ferretti et al., 2007; Beerling et al., 2008; 

Bloom et al., 2010a), laboratory experiments have shown 

that UV-irradiated plant pectin produces CH4 (Vigano et 

al., 2008). While not ascribed to aerobic methanogenesis, Carmo et al. (2006) presented evidence for a canopy 

source of CH4 from upland forest and Rice et al. (2010) 

reported a “bottom-up” estimate of global CH4 emissions 

of 60 ± 20 Tg/year from trees in flooded soils.

In the last decade, satellite observations have become available that fill an important gap in the global 

coverage of the temporal and spatial variations of CH4

(e.g., Frankenberg et al., 2005; Frankenberg et al., 2008). 

Furthermore, satellite observations have also detected an 

increase in global mixing ratios in recent years (Bloom et 

al., 2010b) and identified increased wetland emissions as 

a potential cause, consistent with in situ measurements.

Future trends in atmospheric CH4 are highly uncertain and depend to a large extent on the proportion of 

the carbon stored in Arctic permafrost that emanates as 

CH4 following permafrost thaw and the extent of permafrost thawing. Tarnocai et al. (2009) estimate that frozen 

Arctic soils contain about 1670 petagrams (Pg) of carbon. 

While current observations do not imply a large role for 

thawing permafrost, this could change with large future 

Arctic warming. Additionally, large stores of CH4 exist 

in the form of frozen CH4 hydrates on the sea floor (Buffett, 2004). These potentially could be liberated either 

by warming induced by ocean circulation changes or via 

attempted extraction for fuel usage. Changes in wetland 

extent, agricultural practice, and fossil fuel extraction (excluding hydrates) could also affect future atmospheric CH4

levels, but have much smaller potential for impact than the 

large frozen soil-carbon reservoirs mentioned above.

1.5.1.2 Nitrous Oxide (N2O)

The photochemical degradation of nitrous oxide 

(N2O) in the stratosphere leads to ozone-depleting nitric 

oxide (NO) and nitrogen dioxide (NO2) and to important 

free radical reservoir species (e.g., HNO3, ClONO2). In 

addition, N2O is an important greenhouse gas.

Forster et al. (2007) reported that the global average 

mole fraction of N2O in 2005 was 319 ppb compared to a 

preindustrial value of 270 ppb. The growth rate has been 

approximately constant since 1980 with more than onethird of its emissions being anthropogenic (agriculture, 

etc.). In 2005–2008 the average growth rate of N2O was 

0.8 ppb/yr (NOAA and AGAGE), with an average global 

mixing ratio in 2008 of 321.5 ppb (NOAA) and 321.6 

ppb (AGAGE) (Table 1-15). N2O is an important ozonedepleting and greenhouse gas. For the A2 IPCC scenario, Portmann and Solomon (2007) compute 21st century 

ozone decreases up to 8% in the 20–40 km altitude region 

from changes in N2O alone. By comparing the Ozone 

Depletion Potential-weighted anthropogenic emissions of 

N2O with those of ozone-depleting substances, Ravishankara et al. (2009) show that N2O emissions currently are 

the single most important emissions of a chemical that depletes ozone. Yet, N2O is not controlled by the Montreal 

Protocol. The findings of Ravishankara et al. (2009) are 

discussed in more depth in Sections 5.2 and 5.4 of Chapter 

5. Future anthropogenic emissions of N2O may increase if 

nitrogen-containing fertilizer use is enhanced for the production of biofuels, as discussed by Crutzen et al. (2008). 

Limiting or reducing future N2O emissions, one target of 

the Kyoto Protocol, would enhance the recovery of the 

ozone layer from its depleted state and would also reduce 

anthropogenic forcing of the climate system.

Recently, Huang et al. (2008) used AGAGE, 

NOAA/ESRL, and Commonwealth Scientific and Industrial Research Organisation (CSIRO) observations, the 

3-D MATCH model (1.8° × 1.8°), and a Kalman filter to 

deduce regional N2O emissions. The effects of model errors were assessed using a large ensemble of 2-D model 

inversions. They concluded that global N2O emissions 

with 66% probability errors are 16.3+

−1

1.

.

5

2 and 15.4+−1

1.

.

7

3 Tg 

N (N2O)/yr, for 1997–2001 and 2001–2005, respectively. 

Emissions from the equator to 30°N increased significantly from the earlier Bouwman et al. (1995) estimates, while 

emissions from the southern oceans (30°S–90°S) decreased significantly. Relative to Bouwman et al. (1995), 

Huang et al. (2008) found that land emissions from South 

America, Africa, and China/Japan/South East Asia are 

larger, while land emissions from Australia/New Zealand 

are smaller. Their study also showed a shift of the oceanic sources from the extratropical to the tropical oceans 

relative to Bouwman et al. (1995). Between the periods 

1996–2001 and 2002–2006, emissions increased in China/

Japan/South East Asia, 0°–30°N oceans, and North West 

Asia; emissions decreased in Australia/New Zealand, 

30°S–90°S oceans, 30°N–90°N oceans, and Africa. The 

lower tropical ocean emissions in 1997–2001 relative to 

2002–2005 could result from the effects of the 1997–1998 

El Niño in the earlier period.

The N2O fluxes from the equator to 30°N region 

reported by Hirsch et al. (2006) and Huang et al. (2008) 

are significantly larger than estimated by Bouwman et 

al. (1995) and Prinn et al. (1990), while the emissions 
1.80

Chapter 1

from the southern oceans (30°S–90°S) are significantly 

smaller. The differences between the two recent and two 

earlier studies may be due to either real long-term variations or different modeling and methodological approaches. Huang et al. (2008) found that 23 ± 4% of the global 

total N2O emissions come from the oceans, which is at 

the low range of the Hirsch et al. (2006) and Bouwman 

et al. (1995) estimates. Overall, the Hirsch et al. (2006) 

and Huang et al. (2008) studies show reasonably good 

agreement. Considering the fact that Hirsch et al. (2006) 

only used N2O measurements from NOAA/ESRL for the 

1998 to 2001 time period and the inversion techniques and 

transport models used in the two studies differ as well, 

these good agreements may indicate that the relative contributions of regional N2O surface fluxes to the global total 

may indeed have changed significantly from the Bouwman et al. (1995) estimates for 1990 and the Prinn et al. 

(1990) estimates for 1978–1988. In comparison with 

Prinn et al. (1990), the Huang et al. (2008) study shows 

smaller surface emissions in the 30°S–90°S and 30°N–

90°N regions, significantly higher fluxes in the 0°–30°N 

region, and similar fluxes in the 0°–30°S region. Huang 

et al. (2008) concluded that uncertainty in modeling troposphere and stratosphere exchange is the most significant 

source of uncertainty in deriving regional N2O emissions.

High-precision measurements of atmospheric N2O 

over the last decade reveal subtle signals of interannual 

variability (IAV) superimposed upon the more prominent 

growth trend. Nevison et al. (2007) explored the causes of 

both seasonal and interannual variability using comparisons of a 1993–2004 3-D model simulation to observations of N2O at five AGAGE stations. The model does not 

include a stratospheric sink and thus isolates the effects of 

surface sources and tropospheric transport. Both model 

and observations yield correlations in seasonal and interannual variability among species, but only in a few cases 

are model and observed variability correlated to each other. The results suggest that tropospheric transport contributes significantly to observed variability, especially at the 

Samoa station. However, some features of observed variability (e.g., at Mace Head, Ireland) are not explained by 

the model simulation and appear more consistent with the 

effects of downward mixing of N2O-poor stratospheric air.

Finally, Jiang et al. (2007) carried out a systematic 

study of the N2O seasonal cycle and its latitudinal variation using AGAGE and NOAA/ESRL data. The seasonal 

cycles were statistically significant at Alert (82°N, 62°W), 

Barrow (71°N, 157°W), Mace Head (53°N, 10°W), Cape 

Kumukahi (19°N, 155°W), Cape Matatula (14°S, 171°W), 

Cape Grim (41°S, 145°E), and South Pole (90°S, 102°W). 

The amplitude (peak to peak) of the seasonal cycle varies 

from 0.29 ppb at the South Pole to 1.15 ppb at Alert. The 

month at which the seasonal cycle is at a minimum varies monotonically from April (South Pole) to September 

(Alert). The seasonal cycle in the Northern Hemisphere 

shows the influence of the stratosphere, owing to seasonal 

variations in exchange between the stratosphere and troposphere in the Arctic; the seasonal cycle in the Southern Hemisphere suggests greater influence from surface 

sources.

1.5.1.3 COS, SO2, and Sulfate Aerosols

Carbonyl sulfide (COS) and other sulfur-containing 

gases such as sulfur dioxide (SO2) are important precursors of sulfate aerosols in the stratosphere (Notholt et al., 

2005; SPARC, 2006), which catalyze ozone depletion 

by ODSs (e.g., Newman and Pyle et al., 2003; Danilin 

and McConnell, 1995) and affect the radiative balance of 

the atmosphere (e.g., Forster et al., 2007). The relative 

contributions of the sulfate aerosol precursors are difficult to quantify because SO2 has a short lifetime, and the 

nonvolcanic stratospheric sulfur (S) burden is not well 

determined.

COS is the most abundant sulfur-containing trace 

gas in the atmosphere. Long-term trends in COS mixing 

ratios suggest that global changes during the past decade 

have been fairly small (Aydin et al., 2008; Montzka et al., 

2004; Montzka et al., 2007; Zander et al., 2005). Updated 

data show a mean global surface mixing ratio of 491 ppt in 

2008 and a mean rate of increase of 1.8 ppt/yr during 2000–

2008 (Table 1-15). The current background concentration 

is more than 60% higher than preindustrial values of ~300 

ppt (Montzka et al., 2004; Aydin et al., 2008). Long-term 

COS trends from long-path infrared solar absorption measurements above Jungfraujoch (Switzerland) (updated from 

Zander et al., 2005) show an annual increase in the total 

vertical column of 0.79 ± 0.09% over 2000–2008 (relative 

to January 2000 values) and an increased growth rate of 

1.26 ± 0.29%/yr over 2005–2008 (relative to January 2005 

values). Over decadal-to-centennial periods, COS mixing 

ratios appear correlated to anthropogenic sulfur emissions (Montzka et al., 2004). The updated Jungfraujoch 

data also show this behavior, exhibiting decreases in total 

column COS during the late 1980s to early 1990s. Both 

ground-based flask sampling and Jungfraujoch remote 

sensing results then show a reversal, with increases since 

the mid-2000s, concurrent with 5%/yr increases in global 

coal combustion since then (BP, 2009).

Global flask measurements show large seasonal 

changes across broad geographic scales (Montzka et al., 

2007). The observed variations suggest an important role 

for the oceans in determining the seasonality in the SH 

and uptake by the terrestrial biosphere during the growing 

season in the NH. The amplitude of observed seasonal 

changes in the NH suggests significantly larger vegetative 
1.81

ODSs and Related Chemicals

uptake of COS and, as a result, a shorter global lifetime 

of 2–4 years (Montzka et al., 2007; Suntharalingam et al., 

2008) than derived in earlier studies.

COS mixing ratio distributions derived from satellites have provided an estimate of COS stratospheric loss. 

From correlations to measured CFC abundances, a stratospheric lifetime of 64 ± 21 years can be derived. This 

lifetime suggests that COS contributes 34–66 Gg S/yr to 

the stratosphere (Barkley et al., 2008). Given the wide 

range in estimates of sulfur emissions necessary to maintain the stratospheric sulfate aerosol layer, it is difficult to 

accurately estimate the contribution of sulfur from COS to 

this layer. Some estimates from midlatitude data suggest 

that COS could contribute ~50% of aerosol sulfur mass 

being transported to the midlatitude lower stratosphere 

(Martinsson et al., 2005). Other results implying 300–400 

Gg S/yr to explain observed seasonal changes above some 

NH sites would suggest a much smaller relative contribution of COS to stratospheric sulfur (Hofmann et al., 2009). 

This latter study also provided evidence for large changes 

in stratospheric aerosol during recent years. Because COS 

mixing ratios in the background atmosphere have not increased by more than a few percent since 2000, the cause 

for this aerosol increase is not likely attributable entirely 

to the observed COS changes (Hofmann et al., 2009).

Sources of atmospheric SO2 have been reviewed 

by Stern (2005) and Fisher et al. (2007). Anthropogenic 

emissions of SO2 are mainly due to fossil fuel burning and 

metal smelting. Natural sources consist of the oxidation of 

COS, dimethyl sulfide (DMS), carbon disulfide (CS2), and 

hydrogen sulfide (H2S), as well as emissions from volcanic activities. Based on “bottom-up” estimates (see Box 

1-1), anthropogenic sulfur emissions reached a maximum 

of ~75 Tg S/yr in the 1980s (Stern, 2005). Emissions 

subsequently decreased as a consequence of legislation 

addressing enhanced acidification (in industrialized countries) and because of reduced industrial activity in Eastern 

Europe (Stern, 2005). In 2000, global emissions were estimated to be 55–62 Tg S/yr (Fisher et al., 2007), of which 

~17 Tg S/yr was emitted from Asia (Stern, 2005; Klimont 

et al., 2009). In some projections future emissions from 

Asia are expected to increase substantially (e.g., 40 Tg S/

yr by 2030; Klimont et al., 2009), with the main increase 

expected to occur either in China (Fisher et al., 2007) or 

in India (Klimont et al., 2009). SO2 mixing ratios decline 

rapidly with height and distance from the source regions to 

a few tens of ppt (SPARC, 2006), and therefore the same 

limitations for its transport into the stratosphere apply as 

for VSLS discussed in Section 1.3.

Natural sources of atmospheric SO2, which consist 

of volcanic eruptions and the oxidation of primarily oceanic 

substances (COS, DMS, and H2S), are estimated to account 

for 17–41 Tg S/yr (Haywood and Boucher, 2000). Infrequent explosive volcanic eruptions do temporarily enhance 

the sulfate aerosol burden of the stratosphere dramatically. 

For example, in 1991 the eruption of Mt. Pinatubo added 

~10 Tg S into the atmosphere (Read et al., 1993; Guo et al., 

2004), which partly reached the stratosphere and significantly affected the atmospheric radiative balance. Deliberately enhancing S in the stratosphere is being discussed as a 

geoengineering option for mitigating the heating influence 

from greenhouse gases (see Chapter 5).

1.5.2 Changes in Radiative Trace Gases 

that Indirectly Influence Ozone

Carbon dioxide and fluorinated compounds are 

greenhouse gases and have no direct effect on stratospheric ozone. Nevertheless, they are discussed in this section, 

as their impact on global warming can indirectly lead to 

changes in stratospheric ozone, as discussed in Chapter 4 

of this Assessment.

1.5.2.1 Carbon Dioxide (CO2)

Since 1750 CO2 has been the most important anthropogenic greenhouse gas and it still continues to dominate 

atmospheric radiative forcing at 1.74 W/m2 in 2008 (Hofmann et al., 2006) (Figure 1-24). The 2003.5–2008.5 increase in CO2 radiative forcing (0.14 W/m2

) was six times 

greater than the total direct forcing increase from all other 

radiatively active, long-lived gases in this same period.

In 2008, the global average CO2 mole fraction was 

384.8 ppm (Table 1-15). The global growth rate of CO2

averaged 2.1 ppm/yr between 2005 and 2008, when derived with results from all 37 marine boundary layer sites 

in the NOAA/ESRL surface air sampling network. This 

is similar to the average growth rate for the previous fouryear period (2.0 ppm/yr), but significantly higher than the 

average growth rate in the 1990s (1.5 ppm/yr). The increase in atmospheric growth rate in the last 8 years corresponds closely with acceleration in the combustion of 

fossil fuels (Boden et al., 2009; BP, 2009). The increase 

in global fossil fuel-based CO2 emissions averaged 0.8%/

yr in the 1990s and 3.2%/yr between 2000 and 2008 (with 

estimated emissions of 8.7 Pg C/yr in 2008).

The fraction of fossil carbon remaining in the atmosphere (airborne fraction) shows no discernible trend 

over the past 20 years, with 5-year means stable at about 

50% (Knorr et al., 2009). There is significant interannual 

variability in the atmospheric growth (and airborne fraction), of which about 50% can be explained by ENSO and 

volcanism (Bacastow, 1976; Reichenau and Esser, 2003) 

and their effect on carbon balance of the land biosphere.

Although the number of coupled climate carbon 

cycle models and their mechanistic sophistication has increased, there is still no consensus as to the centennial-
1.82

Chapter 1

scale fate of the airborne fraction and the 50% radiative 

forcing discount that it has provided in recent decades. 

Most models predict a declining efficiency in the uptake 

of atmospheric carbon by the oceans by the middle of the 

century (Friedlingstein et al., 2006). Whether or not this 

decline can be observed already in the Southern Ocean is 

a matter of scientific debate and is discussed in detail in 

Chapter 4. The fate of the land biosphere sink of CO2

is also uncertain, with some models predicting increased 

uptake due to boreal forest expansion and general CO2 fertilization, while others predict drought-stress related die 

back of forests and losses in productivity (Friedlingstein 

et al., 2006). The coupling of the carbon cycle to climate 

variability remains a first-order uncertainty in the prediction of future climate. This is especially true considering 

the uncertain fate of an estimated 1000 Pg C of carbon 

frozen in the top 3 meters of Arctic soils (Tarnocai et al., 

2009), which have not been considered in coupled climate 

carbon cycle models.

1.5.2.2 Fluorinated Greenhouse Gases

With the exception of carbon tetrafluoride (CF4), 

the fluorinated compounds discussed in this section are 

virtually entirely of anthropogenic origin. Hydrofluorocarbons (HFCs) are replacement chemicals for the longlived ODSs in various applications such as refrigeration, 

fire extinguishers, propellants, and foam blowing (IPCC/

TEAP, 2005). The very long-lived perfluorocarbons 

(PFCs) and sulfur hexafluoride (SF6) have been emitted 

over past decades from established industrial processes 

and, in the case of PFCs, recently from newer applications 

largely in the electronics sector. The newest compounds 

detected in the atmosphere are sulfuryl fluoride (SO2F2) 

(used as a replacement for CH3Br) and nitrogen trifluoride 

(NF3) used in the electronic sector.

By mid-2008, the contribution of the fluorinated 

substances (HFCs, PFCs, SF6) to radiative forcing had increased to ~24 mW/m2

. Contributions were 16 mW/m2

for HFCs (including HFC-23), 3.4 mW/m2 for SF6, and 

5.4 mW/m2 from the sum of CF4, C2F6, and C3F8 (Mühle et 

al., 2010), assuming a natural background for CF4 of 34.7 

ppt (Figure 1-24). These contributions to total radiative 

forcing by 2008 are small compared to CO2 (1740 mW/

m2), CH4 (500 mW/m2), CFCs (262 mW/m2), and N2O 

(170 mW/m2), but amount to about half of the radiative 

forcing from HCFCs (45 mW/m2) (Figure 1-24). However, these values are radiative forcings since preindustrial 

times. Considering changes only over the past 5 years 

(2003.5–2008.5), CO2 direct radiative forcing increased 

by 139 mW/m2, that from N2O increased by 12 mW/m2, 

and that from CH4 increased by 4 mW/m2. Radiative forcing from the sum of HFCs, PFCs, and SF6 increased by 

8 mW/m2 over this same period, whereas the ODSs controlled under the Montreal Protocol have positive and 

negative contributions (CFCs, Cl-solvents, and halons: 

−8 mW/m2; HCFCs: +8 mW/m2). This means that during 

2003.5–2008.5, the change in direct radiative forcing from 

the sum of HFCs, PFCs, and SF6 was comparable to the 

change in direct radiative forcing from either CH4 or N2O. 

Further impacts of the potential large future increase of 

fluorinated greenhouse gases are discussed in Chapter 5 of 

this Assessment.

Summed emissions from HFCs used primarily as replacements for ODSs (i.e., not HFC-23) have increased since 

2004. The sum of emissions (weighted by GWP) of HFCs 

used as ODS replacements increased by 8–9%/yr from 2004 

to 2008, and by 2008 amounted to 0.39 ± 0.03 gigatonnes 

of CO2 equivalents per year (Gt CO2-eq/yr). Emissions of 

HFC-23, predominantly the result of HCFC-22 production, 

contributed an additional ~0.2 Gt CO2-eq/yr emission during 

2006–2008 (Montzka et al., 2010). Recently published data 

with higher time resolution suggest that the 2008 HFC-23 

emission may be slightly less than this 3-year average (~0.17 

Gt CO2-eq/yr in 2008; Miller et al., 2010).

The following paragraphs are an update on the 

abundance and the sources of different types of fluorinated 

greenhouse gases (GHGs).

HFC-134a (CH2 FCF3 )

HFC-134a has replaced CFC-12 as the preferred 

refrigerant in refrigeration and mobile air conditioning 

(MAC), and it also has a minor usage in foam-blowing 

applications. Observed global abundances and rates of 

change estimated by two independent global sampling 

networks (NOAA and AGAGE) are in good agreement. 

HFC-134a has been growing steadily and reached a global 

mean mole fraction of ~48 ppt in 2008 (Table 1-15), with 

an average trend of 4.6 ppt/yr (or ~10%/yr) in 2007–2008. 

HFC-134a contributed ~8 mW/m2 to atmospheric radiative forcing in 2008.

A 12-box model analysis of measured changes 

in the global atmospheric abundance of HFC-134a was 

used to derive global emissions of 149 ± 27 Gg/yr during 2008 (approximately 125 ± 16 Gg/yr was derived for 

2005–2006). Stohl et al. (2009) used regional-scale inversions with a global coverage to derive global HFC-134a 

emissions in 2005–2006 of 130–140 Gg/yr, similar to the 

12-box model result for these years once uncertainties are 

considered. Analyses of pollution events observed by 

high-frequency measurements and other regional studies provide insights into regional contributions to these 

global emissions. In the Stohl et al. (2009) study for the 

period 2005–2006, about 40 Gg/yr was attributed to North 

America, 25 Gg/yr to Europe, 43 Gg/yr to Asia, and lesser 

amounts were derived for other regions of the globe. For 

the United States alone, Stohl et al. (2009) derived 28–35 
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ODSs and Related Chemicals

Gg/yr of HFC-134a emission in 2005 and 2006, which is 

similar to the 27 (12–39) Gg/yr HFC-134a emissions from 

the United States estimated from aircraft measurement 

campaigns in 2004 and 2006 (Millet et al., 2009). In a 

separate study in which high-frequency data from three 

stations in Eastern Asia were considered, East Asian emissions of HFC-134a were estimated at 19.2 Gg/yr in 2008, 

of which 12.9 ± 1.7 Gg/yr was attributed to China (Stohl et 

al., 2010). This is only slightly higher than the HFC-134a 

emission estimate for China of 8.7 (6.5–12) Gg/yr in 2008 

derived from high-frequency atmospheric measurements 

in South Korea (Kim et al., 2010).

Because of the long lifetime of HFC-134a (13.4 

years, Table 1-3) and a relatively high GWP100 (1370, 

Chapter 5), the use of HFC-134a will eventually be phased 

out in Europe. This will lead to a very gradual phasedown of the use of HFC-134a in cars, which is expected 

to also take place outside Europe because of the global nature of the car industry. However, in developing countries 

the potential for growth of HFC-134a is potentially large 

(Velders et al., 2009).

HFC-23 (CHF3 )

HFC-23 is primarily emitted to the atmosphere 

from over-fluorination of chloroform during the production of HCFC-22. Other minor emissions of HFC-23 arise 

from the electronics industry, refrigeration, and fire extinguishers (Oram et al., 1998). Due to its long lifetime 

of 222 years (Table 1-3) and continued emissions, HFC23 global mixing ratios reached 22 ppt in 2008, with a 

growth rate of 0.83 ppt/yr (Miller et al., 2010). At this 

global abundance, HFC-23 contributed ~4 mW/m2 to the 

atmospheric radiative forcing in 2008 (Figure 1-24).

A study of Antarctic firn air suggested a 50% increase in global HFC-23 emissions from 8.7 ± 2 Gg/yr during the 1990s to a mean of 13.5 ± 2 Gg/yr during 2006–

2008 (Montzka et al., 2010). HFC-23 emissions increased 

even though emissions reported by developed countries 

declined from 6–8 Gg/yr in the late 1990s to 2.8 Gg/yr in 

2007 and despite the destruction by incineration of 5–7 Gg 

of HFC-23 in developing countries in 2007–2008 through 

United Nations Framework Convention on Climate Change 

(UNFCCC) Clean Development Mechanism (CDM) projects (Montzka et al., 2010). The increase inferred for global 

HFC-23 emissions is coincident with a substantial increase 

in HCFC-22 production in developing countries, which 

accounted for 60% of global HCFC-22 production in 2007. 

The mean yield of HFC-23 emission from global HCFC-22 

production during 2006–2008 was estimated at 1.7 ± 0.3%, 

which is slightly lower than the mean of 2.3% derived for 

the early 1990s (Montzka et al., 2010).

These firn-air-derived global emission estimates 

for HFC-23 are largely consistent with a recent analysis 

of archived air (back to 1978) and ongoing remote atmospheric measurements at multiple sites since late 2007 

(Miller et al., 2010). These ongoing data provide higher 

time resolution during 2006–2008 than the analysis of firn 

measurements, however, and suggest that global HFC-23 

emissions declined after 2006 and were 12.0 (+0.6/−0.7) 

Gg/yr in 2008.

East Asian HFC-23 emission magnitudes and distributions have been derived using inversion modeling of 

in situ measurements at three locations (Stohl et al., 2010). 

Emissions from this region during 2008 were dominated 

by 6.2 ± 0.7 Gg/yr from China. This analysis yielded enhanced emissions from locations where HCFC-22 production facilities are known to be located, both in China and 

Japan (Stohl et al., 2010). The uncertainties (1 standard 

deviation) quoted in Stohl et al. (2010) do not include any 

systematic errors in their dispersion model or in the meteorological input data used in their inversion analysis. A 

higher HFC-23 emission of 12 (8.6–15 Gg/yr) was derived 

in 2008 for China recently from a combined inversion/

ratio method based on data from the Korean Gosan station (Kim et al., 2010). This is similar to the 10 ± 5 Gg/

yr inferred from China for 2004 and 2005 in an earlier 

study (Yokouchi et al., 2005). Substantial HFC-23 emissions from China are likely ongoing in 2008 because less 

than half of HCFC-22 production in developing countries 

during this time was associated with the UNFCCC CDM 

projects (Montzka et al., 2010).

HFC-152a (CH3CHF2 )

HFC-152a is used as a foam-blowing agent and as 

an aerosol propellant (Greally et al., 2007). In 2008 its 

globally averaged mole fraction was 5.9 ppt and its radiative forcing 0.5 mW/m2

. It has a relatively short lifetime of 

1.5 years (Table 1-3), due to efficient removal by OH oxidation. HFC-152a has the smallest GWP100 (133; Chapter 

5) of all major HFCs. Given its short lifetime, the increase 

in background concentrations (0.6 ppt/yr (+11%/yr) in 

2007–2008) implies a substantial increase of emissions 

in recent years. Emissions derived from the observed atmospheric change were 50 Gg/yr in 2008 (AGAGE data), 

compared to 28 Gg/yr in 2004 (Greally et al., 2007).

As was found for HFC-134a, an analysis of mixing 

ratio enhancements above background levels from highfrequency measurements suggest substantial emissions 

of HFC-152a from North America (12.3–15.1 Gg/yr), 

Asia (9.6–9.8 Gg/yr), and Europe (3.5–3.9 Gg/yr) during 

2005–2006 (Stohl et al., 2009). The total global emissions 

derived from this regional study (29 Gg/yr in 2005 and 33 

Gg/yr in 2006) are similar to the global emission totals 

derived from the 12-box analysis of mean global mixing 

ratio changes (34 ± 4 in 2005 and 41 ± 4 Gg in 2005 and 

2006, respectively) once uncertainties are considered.
1.84

Chapter 1

HFC-143a (CH3CF3 )

Mixing ratios of HFC-143a, which is used mainly 

in refrigerant blends, of 1.8 ppt have been reported in 1997 

by Culbertson et al. (2004), with a strong increase (25%/

yr) in the second half of the 1990s. Recent independent 

measurements from the AGAGE network show that HFC143a increased to 8.5 ppt in 2008, increasing by 1.0 ppt/

yr (13%/yr) in 2007–2008 (updated from Greally et al., 

2005), resulting in a radiative forcing of 1 mW/m2 and 

global emissions of 17 Gg/yr in 2008.

HFC-32 (CH2 F2 )

Mixing ratios of HFC-32, which is mainly used in 

refrigerant blends, were reported by Greally et al. (2005) 

to be 0.7 ppt in 2004 at Mace Head (Ireland). Updated 

measurements from the AGAGE network show that mixing ratios have increased to 2.7 ppt in 2008 with a yearly 

increase of 0.6 ppt (26%) in 2007–2008. In 2008, emissions of 8.9 Gg/yr have been derived by applying the 12-

box model to AGAGE measurements.

HFC-125 (CHF2 CF3 )

HFC-125 is used in refrigeration blends and for fire 

suppression (IPCC/TEAP, 2005). Background mixing 

ratios have grown to 6.1 ppt in 2008, which results in a 

direct radiative forcing contribution of 1.7 mW/m2. The 

interhemispheric gradient (2008) and growth rate (2007–

2008) were 1.4 ppt and 0.9 ppt/yr (16%/yr), respectively 

(O’Doherty et al., 2009; Table 1-15), from which global 

emissions increasing from 7.5 Gg in 2000 to 22 Gg in 

2008 are derived. European emissions in 2007 have been 

estimated at 3.7–5.5 Gg/yr (O’Doherty et al., 2009).

HFC-365mfc (CF3CH2 CF2 CH3 ) and HFC-245fa 

(CF3CH2 CHF2 )

HFC-365mfc and HFC-245fa are replacements for 

HCFC-141b in foam-blowing applications (Vollmer et al., 

2006). HFC-365mfc showed low mixing ratios of 0.05 

ppt in early 2003 at Jungfraujoch (Switzerland) (Stemmler et al., 2007). An update, using AGAGE and SOGE 

measurements, shows a global mixing ratio of 0.44 ppt in 

2008, with a mixing ratio increase of 0.05 ppt (11%/yr) 

in 2007–2008, and global emissions of 3 Gg/yr in 2008. 

Vollmer et al. (2006) reported HFC-245fa mixing ratios 

of 0.3 ppt at Jungfraujoch in 2004 and estimated global 

emissions of 5.1–5.9 Gg/yr during 2005. The remoteatmosphere mixing ratio of HFC-245fa (results from one 

site in each hemisphere) reached 1.0 ppt in 2008 and had 

increased by 0.4 ppt (32%) from 2007 to 2008 (Vollmer 

et al., 2006).

HFC-227ea (CF3CHFCF3 )

HFC-227ea is mainly used for fire suppression and 

to a lesser extent in metered dose inhalers, refrigeration, 

and foam blowing. Laube et al. (2010) have recently reported the use of firn air measurements from Greenland to 

reconstruct the atmospheric history of HFC-227ea in the 

Northern Hemisphere. These results indicated the mixing 

ratio has grown from less that 0.1 ppt in the 1990s to 0.59 

ppt in 2007. The rate of growth increased from 0.026 ppt/

yr in 2000 to 0.057 ppt/yr in 2007. Global emissions were 

estimated to be ~1.8 Gg/yr in 2007.

Perfluorocarbons (PFCs)

Perfluorocarbons (PFCs) have very large radiative 

efficiencies and lifetimes in the range of 2,000 to 50,000 

years (Table 1-3). PFC-14 (CF4) is largely emitted as a 

by-product of aluminium production and to a smaller degree from the electronics industry (plasma etching), while 

for PFC-116 (C2F6) both sources are significant. PFC218 (C3F8) is largely emitted by the electronics industry, 

with very small contributions from aluminium smelting 

and increasing contributions from refrigeration use. The 

origin of PFC-c-318 (c-C4F8) is uncertain but possibly 

due to a combination of electronics (plasma etching) and 

plastics (PTFE) thermal decomposition (Harnisch, 1999; 

Harnisch, 2000).

Recently developed gas chromatography-mass 

spectrometry (GC-MS) instrumentation has been deployed 

at AGAGE sites, achieving significantly improved precisions on annual means for CF4 (~0.1%), C2F6 (~0.8%), 

and C3F8 (~3%) (Greally et al., 2005; Miller et al., 2008; 

Mühle et al., 2010). Global mixing ratios (2008) and annual growth rates (2007–2008) for CF4, C2F6, and C3F8 of 

77.1 ppt and 0.9%/yr, 3.9 ppt and 2.6%/yr, and 0.5 and 

5.2%/yr, respectively, were measured in the AGAGE network. The radiative forcing contributions by 2008 for CF4, 

C2F6, and C3F8 were 4, 1, and 0.1 mW/m2, respectively.

The preindustrial CF4 level has been re-estimated 

at 34.7 ± 0.2 ppt (Mühle et al., 2010), based on analysis of 

Greenland and Antarctic firn air samples, slightly less than 

the 39 ± 6 ppt estimated previously (Harnisch et al., 1996). 

Direct evidence of a natural source (crustal degassing) of 

CF4 sufficient to maintain preindustrial CF4 abundances 

(34 ppt) has been found in desert groundwaters (Deeds et 

al., 2008).

From the early 1970s to the late 1990s, Greenland 

and Antarctic firn data suggest a small decline in CF4

emissions from 13 Gg/yr to 11–12 Gg/yr (Worton et al., 

2007). Mühle et al. (2010) report that emissions were ∼15 

Gg/yr in 1975, rising to ∼18 Gg/yr around 1980, generally 

declining to ∼11 Gg/yr in 2000, and stabilizing at ∼11 Gg/

yr thereafter. They derive a growing difference between 
1.85

ODSs and Related Chemicals

emissions reported by the global aluminium industry and 

emissions derived from atmospheric measurements in 

the last years. They suggest that either nonmetallic CF4

emissions (possibly from the electronics sector) are growing or that the “bottom-up” approach to estimating CF4

emissions from the aluminium industry is underestimating emissions, or a combination of both. The influence of 

the semiconductor industry on global C2F6 emissions can 

be seen in both “bottom-up” and “top-down” estimates of 

emissions, increasing from about 0.5 Gg/yr to 1.8 Gg/yr 

between 1990 and 2001 (Worton et al., 2007).

Sulfur Hexafluoride (SF6 )

Sulfur hexafluoride (SF6) is an important greenhouse gas because it combines a high radiative efficiency 

with a very long lifetime and a considerable annual increase. Global average (NOAA, AGAGE) mixing ratios 

of SF6 reached 6.4 ppt in 2008, with a yearly (2007–2008) 

increase of 0.2 ppt/yr (3%), resulting in a contribution 

to radiative forcing of 3.4 mW/m2 by 2008. Levin et al. 

(2010) found a similar mixing ratio of 6.7 ppt in 2008 and 

inferred a global emission of 7.16 Gg/yr in 2008. Comparable global SF6 growth rates (0.2–0.3 ppt/yr) have 

been derived from the MIPAS satellite data in 2002–2004 

(Stiller et al., 2008), and solar spectroscopy at Jungfraujoch has been used to derive long-term trends in the total 

column of SF6 (Zander et al., 2008).

Nitrogen Trifluoride (NF3 )

NF3, which is used as a replacement for PFCs in 

plasma etching, in the semiconductor industry, and in the 

production of flat panel displays, has recently been discovered in the atmosphere by Weiss et al. (2008). Although 

NF3 has a high GWP it is not currently included in the 

Kyoto Protocol (Prather and Hsu, 2008). In 2008 the mean 

global tropospheric concentration was 0.45 ppt, increasing 

at 0.05 ppt/yr, or 11%/yr (Weiss et al., 2008). “Bottomup” emissions in 2006 were estimated at 0.14 Gg/yr (Robson et al., 2006). This emission figure was corrected to be 

0.62 Gg/yr in 2008, based on the measured global background abundance and trend by Weiss et al. (2008).

The lifetime and GWP of NF3 have been revised 

since the previous Assessment by Prather and Hsu (2008). 

As a result the NF3 lifetime (740 years) and GWP100

(17,200) given in IPCC’s 4th Assessment (Forster et al., 

2007) have now been revised to 500 years and 17,500, respectively (see Section 1.2.2 and Chapter 5).

Sulfuryl Fluoride (SO2 F2 )

Sulfuryl fluoride (SO2F2) is used as a fumigant 

to replace methyl bromide (except for quarantine/preshipment uses). The global total atmospheric lifetime of 

SO2F2 has been recently assessed to be 36 ± 11 years 

(Mühle et al., 2009), which is significantly longer than previous estimates (<5 years) (Table 1-3). Oceanic hydrolysis 

is the major sink with the global oceanic uptake atmospheric 

lifetime being 40 ± 13 years. Compared to hydrolysis, gasphase tropospheric and stratospheric loss processes are 

only marginally important (Papadimitriou et al., 2008b; 

Dillon et al., 2008). The global tropospheric background 

concentration of SO2F2 has increased from ~0.3 ppt in 

1978 to 1.51 ppt in 2008, with a yearly increase (2007–

2008) of 0.04 ppt (3%). Papadimitriou et al. (2008b) calculated that the GWP100 of SO2F2 is 4780, similar to that of 

CFC-11, using newly measured infrared absorption cross 

sections and the lifetime reported by Mühle et al. (2009). A 

slightly lower GWP100 of 4740 is reported for SO2F2 in 

Table 1-11 owing to a slightly different radiative efficiency 

being used in the calculations appearing in Chapter 5 (see 

note 19 to Table 1-11). Sulbaek Anderson et al. (2009b) 

list a GWP100 range of 120–7600 for SO2F2 lifetimes from 

1–100 years. The emissions calculated from atmospheric 

observations increased from ~0.6 Gg in 1978 to ~1.9 Gg in 

2007. Global production magnitudes of SO2F2 are on average 1.5 times emissions deduced from global atmospheric 

measurements. This suggests that about one-third of SO2F2

may be destroyed during application (like CH3Br) or the 

presence of additional losses.

1.5.3 Emissions of Rockets and Their 

Impact on Stratospheric Ozone

In this section the historic and actual emissions of 

rockets and their potential to impact stratospheric ozone 

are discussed. The future emissions and their impacts will 

be discussed in Section 5.4.2.5 of Chapter 5 of this Assessment.

A variety of propellant combinations contribute to 

the global emissions of rockets. All produce gases and particles that affect ozone chemistry in the stratosphere and 

mesosphere to varying degrees. The emissions from solid 

rocket motors (SRM) are much better understood than the 

emissions from the three liquid propellant types—liquid 

oxygen/kerosene (O2/HC), cryogenic (O2/H2), hypergolic 

(N2O4/hydrazine), and hybrid (N2O/HC).

It is thought that chlorine and aluminum oxide 

(“alumina”) from SRMs account for most of the global 

ozone loss associated with worldwide rocket launches. 

Global, annually averaged ozone losses due to these emissions are estimated to be of the order of 0.1% or less (Ross 

et al., 2009), though this still has not been confirmed with 

observations or models.

Since the last Assessment of rocket impacts on 

stratospheric ozone (Newman and Pyle et al., 2003), there 
1.86

Chapter 1

have been some important developments that impact projections of ozone losses due to rockets. First, the decreasing trend in rocket launches at the start of the 21st century 

has reversed, with the global annual launch rate increasing by 20% since 2005 (http://planet4589.org/space/log/

launchlog.txt). Second, recent studies of the mass fraction 

of SRM sub-micron alumina emissions call into question 

the assumptions that have gone into previous model estimates of the contribution of heterogeneous chlorine reactions on alumina to global ozone loss. Values for that submicron mass ranging from 2% (Danilin et al., 2001), 8% 

(Schmid et al., 2003), 12% (Brady and Martin, 1997), and 

50% (Gossé et al., 2006) have now been reported. These 

sub-micron particles have a disproportionately large impact on ozone abundances, not only because of their high 

surface-to-mass ratios, but also because of their relatively 

long lifetimes above the tropopause compared to larger 

particles. In addition, the previously assumed heterogeneous reaction rates may be low compared to the actual 

reactivity in stratospheric plumes (Danilin et al., 2003). If 

the larger values of these parameters are confirmed, previous studies will have significantly underestimated ozone 

losses from SRMs by as much as a factor of ten. It will 

also be the case that rocket emissions of alumina will have 

a greater impact on ozone than rocket emissions of chlorine, in which case much larger impacts could be expected 

in cold, low-UV regions of the stratosphere (e.g., lowermost stratosphere and polar vortices in springtime). The 

geographic variations in ozone loss due to rockets have yet 

to be studied in detail.

There have been few studies of the role of rocket 

emissions of nitrogen oxides (NOx). An estimate of global 

ozone loss from a hypergolic propellant (N2O4/hydrazine) 

rocket did show that it caused approximately 2% of the 

ozone loss from an SRM rocket of approximately the same 

payload (Ross et al., 2004). However, there are no measurements of NOx emissions from these rockets to validate 

the NOx emissions assumed in the model.

The impacts on ozone by rocket emissions other 

than alumina and chlorine remain unclear (e.g., water 

vapor, NOx, HCs, and soot). Given the tendency for heterogeneous reactions to enhance ozone destruction, it is 

likely that all H2O-containing particles produced by rockets, directly or as ice nucleation sources (e.g., alumina and 

soot), will be net-destroyers of ozone. Since rockets emit 

exhaust throughout the stratosphere, the net impact of NOx

emissions is most-likely ozone destruction via catalytic 

NOx reactions and increased aerosol surface area with subsequent halogen activation. Rocket emissions of H2O into 

the winter polar stratosphere and summer mesosphere can 

also increase occurrence frequencies of polar stratospheric 

and mesospheric clouds (Stevens et al., 2005; Meier, et 

al., 2010). At current launch rates, the impacts of nonSRM rocket emissions are thought to be less important 

than SRM emissions of alumina and chlorine (Ross et al., 

2009).

REFERENCES

Abbatt, J.P.D., and J.G. Anderson, High-pressure discharge flow kinetics and frontier orbital mechanistic analysis for OH + CH2CCl2, cis-CHClCHCl, 

trans-CHClCHCl, CFClCF2, and CF2CCl2 →

products, J. Phys. Chem., 95 (6), 2382-2390, doi: 

10.1021/j100159a049, 1991.

Acerboni, G., J.A. Beukes, N.R. Jensen, J. Hjorth, G. Myhre, 

C.J. Nielsen, and J.K. Sundet, Atmospheric degradation and global warming potentials of three perfluoroalkenes, Atmos. Environ., 35, 4113-4123, 2001.

Aizawa, M., K. Asaoka, M. Atsumi, and T. Sakou, Seaweed bioethanol production in Japan – The Ocean 

Sunrise Project, Oceans 2007, 1-5, doi: 10.1109/

OCEANS.2007.4449162, 2007.

Allen, N.D.C., P.F. Bernath, C.D. Boone, M.P. Chipperfield, D. Fu, G.L. Manney, D.E. Oram, G.C. Toon, 

and D.K. Weisenstein, Global carbon tetrachloride 

distributions obtained from the Atmospheric Chemistry Experiment (ACE), Atmos. Chem. Phys., 9 (19), 

7449-7459, doi: 10.5194/acp-9-7449-2009, 2009.

Andreae, M.O., and P. Merlet, Emission of trace gases 

and aerosols from biomass burning, Global Biogeochem. Cycles, 15 (4), 955-966, doi: 10.1029/ 

2000GB001382, 2001.

Aschmann, J., B.-M. Sinnhuber, E.L. Atlas, and S.M. 

Schauffler, Modeling the transport of very shortlived substances into the tropical upper troposphere 

and lower stratosphere, Atmos. Chem. Phys., 9 (23), 

9237-9247, doi: 10.5194/acp-9-9237-2009, 2009.

Atkinson, R., D.L. Baulch, R.A. Cox, J.N. Crowley, R.F. 

Hampson, R.G. Hynes, M.E. Jenkin, M.J. Rossi, J. 

Troe, and T.J. Wallington, Atmos. Chem. Phys., 8, 

4141-4496, 2008. Updates at: IUPAC Subcommittee for Gas Kinetic Data Evaluation: http://www.

iupac-kinetic.ch.cam.ac.uk.

Aydin, M., M.B. Williams, C. Tatum, and E.S. Saltzman, 

Carbonyl sulfide in air extracted from a South Pole ice 

core: A 2000 year record, Atmos. Chem. Phys., 8 (24), 

7533-7542, doi: 10.5194/acp-8-7533-2008, 2008.

Bacastow, R.B., Modulation of atmospheric carbon dioxide by the Southern Oscillation, Nature, 261, 116-

118, doi: 10.1038/261116a0, 1976.

Barkley, M.P., P.I. Palmer, C.D. Boone, P.F. Bernath, 

and P. Suntharalingam, Global distributions of 

carbonyl sulfide in the upper troposphere and 

stratosphere, Geophys. Res. Lett., 35, L14810, doi: 
1.87

ODSs and Related Chemicals

10.1029/2008GL034270, 2008.

Beerling, D.J., T. Gardiner, G. Leggett, A. McLeod, 

and W.P. Quick, Missing methane emissions 

from leaves of terrestrial plants, Global Change 

Biol., 14 (8), 1821-1826, doi: 10.1111/j.1365-

2486.2008.01607.x, 2008.

Bell, N., L. Hsu, D.J. Jacob, M.G. Schultz, D.R. Blake, 

J.H. Butler, D.B. King, J.M. Lobert, and E. Maier-Reimer, Methyl iodide: Atmospheric budget 

and use as a tracer of marine convection in global 

models, J. Geophys. Res., 107 (D17), 4340, doi: 

10.1029/2001JD001151, 2002.

Bergamaschi, P., C. Frankenberg, J.F. Meirink, M. Krol, 

M.G. Villani, S. Houweling, F. Dentener, E.J. Dlugokencky, J.B. Miller, L.V. Gatti, A. Engel, and 

I. Levin, Inverse modeling of global and regional 

CH4 emissions using SCIAMACHY satellite retrievals, J. Geophys. Res., 114, D22301, 28, doi: 

10.1029/2009JD012287, 2009.

Berthet, G., J.G. Esler, and P.H. Haynes, A Lagrangian 

perspective of the tropopause and the ventilation of 

the lowermost stratosphere, J. Geophys. Res., 112, 

D18102, doi: 10.1029/2006JD008295, 2007.

Bhatia, A., H. Pathak, P.K. Aggarwal, and N. Jain, Tradeoff between productivity enhancement and global 

warming potential of rice and wheat in India, Nutr. 

Cycl. Agroecosyst., 86 (3), 413-424, doi: 10.1007/

s10705-009-9304-5, 2010.

Bilde, M., T.J. Wallington, C. Ferronato, J.J. Orlando, G.S. 

Tyndall, E. Estupiñan, and S. Haberkorn, Atmospheric chemistry of CH2BrCl, CHBrCl2, CHBr2Cl, 

CF3CHBrCl, and CBr2Cl2, J. Phys. Chem. A, 102 

(11), 1976-1986, doi: 10.1021/jp9733375, 1998.

Blei, E., C.J. Hardacre, G.P. Mills, K.V. Heal, and M.R. 

Heal, Identification and quantification of methyl 

halide sources in a lowland tropical rainforest, Atmos. Environ., 44 (8), 1005-1010, doi: 10.1016/j.

atmosenv.2009.12.023, 2010.

Bloom, A.A., J. Lee-Taylor, S. Madronich, D.J. Messenger, P.I. Palmer, D.S. Reay, and A.R. McLeod, 

Global methane emission estimates from ultraviolet irradiation of terrestrial plant foliage, New 

Phytologist, 187, 417-425, doi: 10.1111/j.1469-

8137.2010.03259.x, 2010a.

Bloom, A.A, P.I. Palmer, A. Fraser, D.S. Reay, and C. 

Frankenberg, Large-scale controls of methanogenesis inferred from methane and gravity spaceborne 

data, Science, 327 (5963), 322-325, doi: 10.1126/

science.1175176, 2010b.

Boden, T.A., G. Marland, and R.J. Andres, Global, regional, 

and national fossil-fuel CO2 emissions, Carbon Dioxide Information Analysis Center, Oak Ridge National 

Laboratory, U.S. Department of Energy, Oak Ridge, 

TN, U.S.A., doi: 10.3334/CDIAC/00001, 2009.

Bönisch, H., A. Engel, J. Curtius, Th. Birner, and P. Hoor, 

Quantifying transport into the lowermost stratosphere using simultaneous in-situ measurements of 

SF6 and CO2, Atmos. Chem. Phys., 9 (16), 5905-

5919, doi: 10.5194/acp-9-5905-2009, 2009.

Bousquet, P., P. Ciais, J.B. Miller, E.J. Dlugokencky, 

D.A. Hauglustaine, C. Prigent, G.R. Van der Werf, 

P. Peylin, E.-G. Brunke, C. Carouge, R.L. Langenfelds, J. Lathière, F. Papa, M. Ramonet, M. 

Schmidt, L.P. Steele, S.C. Tyler, and J. White, 

Contribution of anthropogenic and natural sources 

to atmospheric methane variability, Nature, 443, 

439-443, doi: 10.1038/nature05132, 2006.

Bouwman, A.F., K.W. Van der Hoek, and J.G.J. Olivier, 

Uncertainties in the global source distribution of 

nitrous oxide, J. Geophys. Res., 100 (D2), 2785-

2800, 1995.

BP (British Petroleum), BP Statistical Review of World 

Energy 2009, available: http://www.bp.com/liveassets/bp_internet/globalbp/globalbp_uk_english/

reports_and_publications/statistical_energy_review_2008/STAGING/local_assets/2009_downloads/statistical_review_of_world_energy_full_report_2009.pdf, 2009.

Brady, B.B., and L.R. Martin, Modeling the multiphase 

atmospheric chemistry of launch clouds, J. Space. 

Rockets, 34 (6), 780-784, 1997.

Buffett, B., and D. Archer, Global inventory of methane 

clathrate: Sensitivity to changes in the deep ocean, 

Earth Planet. Sci. Lett., 227 (3-4), 185-199, doi: 

10.1016/j.espl.2004.09.005, 2004.

Butler, J.H., S.A. Montzka, A.D. Clarke, J.M. Lobert, and 

J.W. Elkins, Growth and distribution of halons in 

the atmosphere, J. Geophys. Res., 103 (D1), 1503-

1511, doi: 10.1029/97JD02853, 1998.

Butler, J.H., M. Battle, M.L. Bender, S.A. Montzka, A.D. 

Clarke, E.S. Saltzman, C.M. Sucher, J.P. Severinghaus, and J.W. Elkins, A record of atmospheric 

halocarbons during the twentieth century from 

polar firn air, Nature, 399 (6738), 749-755, doi: 

10.1038/21586, 1999.

Butler, J.H., D.B. King, J.M. Lobert, S.A. Montzka, S.A. 

Yvon-Lewis, B.D. Hall, N.J. Warwick, D.J. Mondeel, M. Aydin, and J.W. Elkins, Oceanic distributions and emissions of short-lived halocarbons, 

Global Biogeochem. Cycles, 21, GB1023, doi: 

10.1029/2006GB002732, 2007.

Butler, J.H., T.G. Bell, B.D. Hall, B. Quack, L.J. Carpenter, and J. Williams, Technical Note: Ensuring 

consistent, global measurements of very short-lived 

halocarbon gases in the ocean and atmosphere, Atmos. Chem. Phys., 10 (2), 327-330, doi: 10.5194/

acp-10-327-2010, 2010.

Butz, A., H. Bösch, C. Camy-Peyret, M.P. Chipperfield, 
1.88

Chapter 1

M. Dorf, S. Kreycy, L. Kritten, C. Prados-Roman, 

J. Schwärzle, and K. Pfeilsticker, Constraints on 

inorganic gaseous iodine in the tropical upper troposphere and stratosphere inferred from balloonborne solar occultation observations, Atmos. Chem. 

Phys., 9 (18), 7229-7242, doi: 10.5194/acp-9-7229-

2009, 2009.

Calvert, J.G., R.G. Derwent, J.J. Orlando, G.S. Tyndall, 

and T.J. Wallington, Mechanisms of Atmospheric 

Oxidation of the Alkanes, 1008 pp., Oxford University Press, New York, NY, 2008.

Carmo, J.B., M. Keller, J.D. Dias, P.B. Camargo, and P. 

Crill, A source of methane from upland forests in 

the Brazilian Amazon, Geophys. Res. Lett., 33, 

L04809, doi: 10.1029/2005GL025436, 2006.

Carpenter, L.J., C.E. Jones, R.M. Dunk, K.E. Hornsby, 

and J. Woeltjen, Air-sea fluxes of biogenic bromine 

from the tropical and North Atlantic Ocean, Atmos. 

Chem. Phys., 9 (5), 1805-1816, doi: 10.5194/acp-9-

1805-2009, 2009.

Chen, J., V. Young, H. Niki, and H. Magid, Kinetic and 

mechanistic studies for reactions of CF3CH2CHF2

(HFC-245fa) initiated by H-atom abstraction using 

atomic chlorine, J. Phys. Chem. A, 101 (14), 2648-

2653, doi: 10.1021/jp963735s, 1997.

Chen, L., S. Kutsuna, K. Tokuhashi, and A. Sekiya, Kinetics of the gas-phase reaction of CF3OC(O)H with 

OH radicals at 242-328 K, Int. J. Chem. Kinet., 36

(6), 337-344, doi: 10.1002/kin.20004, 2004a.

Chen, L., S. Kutsuna, K. Tokuhashi, and A. Sekiya, Kinetics study of the gas-phase reactions of C2F5OC(O)H 

and n-C3F7OC(O)H with OH radicals at 253-328 

K, Chem. Phys. Lett., 400 (4-6), 563-568, 2004b.

Chen, L., S. Kutsuna, K. Tokuhashi, and A. Sekiya, 

Kinetics study of the gas-phase reactions of 

CHF2CF2OCHF2 and CF3CHFCF2OCH2CF2CF3

with OH radicals at 253-328 K, Chem. Phys. Lett., 

403 (1-3), 180-184, 2005a.

Chen, L., S. Kutsuna, K. Tokuhashi, A. Sekiya, R. Tamai, and Y. Hibino, Kinetics and mechanism of 

(CF3)2CHOCH3 reaction with OH radicals in an environmental reaction chamber, J. Phys. Chem. A, 109 

(21), 4766-4771, doi: 10.1021/jp050491f, 2005b.

Chen, Y.-H., and R.G. Prinn, Estimation of atmospheric 

methane emissions between 1996 and 2001 using a three-dimensional global chemical transport 

model, J. Geophys. Res., 111, D10307, 25, doi: 

10.1029/2005JD006058, 2006.

Chipperfield, M.P., M. Burton, W. Bell, C.P. Walsh, T. 

Blumenstock, M.T. Coffey, J.W. Hannigan, W.G. 

Mankin, B. Galle, J. Mellqvist, E. Mahieu, R. Zander, 

J. Notholt, B. Sen, and G.C. Toon, On the use of HF 

as a reference for the comparison of stratospheric observations and models, J. Geophys. Res., 102 (D11), 

12901-12919, doi: 10.1029/96JD03964, 1997.

Christensen, L.K., J. Sehested, O.J. Nielsen, M. Bilde, 

T.J. Wallington, A. Guschin, L.T. Molina, and 

M.J. Molina, Atmospheric chemistry of HFE-7200 

(C4F9OC2H5): Reaction with OH radicals and fate 

of C4F9OCH2CH2O(·) and C4F9OCHO(·)CH3 radicals, J. Phys. Chem. A, 102 (25), 4839-4845, doi: 

10.1021/jp981128u, 1998.

Class, Th., and K. Ballschmiter, Global baseline pollution 

studies, X. Atmospheric halocarbons: Global budget estimations for tetrachloroethene, 1,2-dichloroethane, l,l,l,2-tetrachloroethane, hexachloroethane 

and hexachlorobutadiene. Estimation of the hydroxyl radical concentrations in the troposphere of 

the northern and southern hemisphere, Fresenius’ 

Z. Anal. Chem., 327 (2), 198-204, doi: 10.1007/

BF00469817, 1987.

Clerbaux, C., and D.M. Cunnold (Lead Authors), 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, and Y. Yokouchi, 

Long-lived compounds, Chapter 1 in Scientific Assessment of Ozone Depletion: 2006, Global Ozone 

Research and Monitoring Project—Report No. 50, 

572 pp., World Meteorological Organization, Geneva, Switzerland, 2007.

Collins, W.J., R.G. Derwent, C.E. Johnson, and D.S. 

Stevenson, The oxidation of organic compounds 

in the troposphere and their global warming potentials, Clim. Change, 52 (4), 453-479, doi: 

10.1023/A:1014221225434, 2002.

Conway, T.J., P.P. Tans, L.S. Waterman, K.W. Thoning, 

D.R. Kitzis, K.A. Masarie, and N. Zhang, Evidence 

for interannual variability of the carbon cycle from 

the National Oceanic and Atmospheric Administration/Climate Monitoring and Diagnostics Laboratory 

Global Air Sampling Network, J. Geophys. Res., 99

(D11), 22831-22855, doi: 10.1029/94JD01951, 1994.

Cox, M.L., P.J. Fraser, G.A. Sturrock, S.T. Siems, and 

L.W. Porter, Terrestrial sources and sinks of halomethanes near Cape Grim, Tasmania, Atmos. 

Environ., 38 (23), 3839-3852, doi: 10.1016/j.atmosenv.2004.03.050, 2004.

Crutzen, P.J., A.R. Mosier, K.A. Smith, and W. Winiwarter, N2O release from agro-biofuel production 

negates global warming reduction by replacing fossil fuels, Atmos. Chem. Phys., 8 (2), 389-395, doi: 

10.5194/acp-8-389-2008, 2008.

Culbertson, J.A., J.M. Prins, E.P. Grimsrud, R.A. Rasmussen, M.A.K. Khalil, and M.J. Shearer, Observed 

trends for CF3-containing compounds in background air at Cape Meares, Oregon, Point Barrow, 

Alaska, and Palmer Station, Antarctica, Chemo-
1.89

ODSs and Related Chemicals

sphere, 55 (8), 1109-1119, doi: 10.1016/j.chemosphere.2003.11.002, 2004.

Daniel, J.S., and G.J.M Velders (Lead Authors), A.R. 

Douglass, P.M.D. Forster, D.A. Hauglustaine, 

I.S.A. Isaksen, L.J.M. Kuijpers, A. McCulloch, and 

T.J. Wallington, Halocarbon scenarios, ozone depletion potentials, and global warming potentials, 

Chapter 8 in Scientific Assessment of Ozone Depletion: 2006, Global Ozone Research and Monitoring 

Project—Report No. 50, 572 pp., World Meteorological Organization, Geneva, Switzerland, 2007.

Danilin, M.Y., and J.C. McConnell, Stratospheric effects of bromine activation on/in sulfate aerosol, 

J. Geophys. Res., 100 (D6), 11237-11243, doi: 

10.1029/95JD00999, 1995.

Danilin, M.Y., R.-L. Shia, M.K.W. Ko, D.K. Weisenstein, 

N.D. Sze, J.J. Lamb, T.W. Smith, P.D. Lohn, and 

M.J. Prather, Global stratospheric effects of the 

alumina emissions by solid-fueled rocket motors, 

J. Geophys. Res., 106 (D12), 12727-12738, doi: 

10.1029/2001JD900022, 2001.

Danilin, M.Y., P.J. Popp, R.L. Herman, M.K.W. Ko, M.N. 

Ross, C.E. Kolb, D.W. Fahey, L.M. Avallone, 

D.W. Toohey, B.A. Ridley, O. Schmid, J.C. Wilson, D.G. Baumgardner, R.R. Friedl, T.L. Thompson, and J.M. Reeves, Quantifying uptake of HNO3

and H2O by alumina particles in Athena-2 rocket 

plume, J. Geophys. Res., 108 (D4), 4141, doi: 

10.1029/2002JD002601, 2003.

Deeds, D.A., M.K. Vollmer, J.T. Kulongoski, B.R. Miller, 

J. Mühle, C.M. Harth, J.A. Izbicki, D.R. Hilton, and 

R.F. Weiss, Evidence for crustal degassing of CF4

and SF6 in Mojave Desert groundwaters, Geochim. 

Cosmochim. Acta, 72 (4), 999-1013, doi: 10.1016/j.

gca.2007.11.027, 2008.

Denman, K.L., G. Brasseur, A. Chidthaisong, P. Ciais, 

P.M. Cox, R.E. Dickinson, D. Hauglustaine, C. 

Heinze, E. Holland, D. Jacob, U. Lohmann, S. 

Ramachandran, P.L. da Silva Dias, S.C. Wofsy, 

and X. Zhang, Couplings between changes in the 

climate system and biogeochemistry, Chapter 7 in 

Climate Change 2007: The Physical Science Basis. 

Contribution of Working Group I to the Fourth Assessment Report of the Intergovernmental Panel on 

Climate Change, edited by S. Solomon, D. Qin, M. 

Manning, Z. Chen, M. Marquis, K.B. Averyt, M. 

Tignor, and H.L. Miller, 996 pp., Cambridge University Press, Cambridge, U.K., and New York, 

NY, U.S.A., 2007.

Derwent, R.G., M.E. Jenkin, S.M. Saunders, and M.J. 

Pilling, Photochemical ozone creation potentials 

for organic compounds in Northwest Europe calculated with a master chemical mechanism, Atmos. 

Environ., 32 (14-15), 2429-2441, doi: 10.1016/

S1352-2310(98)00053-3, 1998.

Derwent, R.G., M.E. Jenkin, N.R. Passant, and M.J. Pilling, Reactivity-based strategies for photochemical 

ozone control in Europe, Environ. Sci. Policy, 10

(5), 445-453, doi: 10.1016/j.envsci.2007.01.005, 

2007.

Dessens, O., G. Zeng, N. Warwick, and J. Pyle, Short-lived 

bromine compounds in the lower stratosphere; impact of climate change on ozone, Atmos. Sci. Lett.,

10 (3), 201-206, doi: 10.1002/asl.236, 2009.

Dillon, T.J., A. Horowitz, and J.N. Crowley, The atmospheric chemistry of sulphuryl fluoride, SO2F2, Atmos. Chem. Phys., 8 (6), 1547-1557, doi: 10.5194/

acp-8-1547-2008, 2008.

Dimmer, C.H., P.G. Simmonds, G. Nickless, and M.R. 

Bassford, Biogenic fluxes of halomethanes from Irish 

peatland ecosystems, Atmos. Environ., 35 (2), 321-

330, doi: 10.1016/S1352-2310(00)00151-5, 2001.

Dlugokencky, E.J., L. Bruhwiler, J.W.C. White, L.K. Emmons, P.C. Novelli, S.A. Montzka, K.A. Masarie, 

P.M. Lang, A.M. Crotwell, J.B. Miller, and L.V. 

Gatti, Observational constraints on recent increases 

in the atmospheric CH4 burden, Geophys. Res. Lett.,

36, L18803, doi: 10.1029/2009GL039780, 2009.

Donner, L.J., L.W. Horowitz, A.M. Fiore, C.J. Seman, 

D.R. Blake, and N.J. Blake, Transport of radon-222 

and methyl iodide by deep convection in the GFDL 

Global Atmospheric Model AM2, J. Geophys. Res., 

112, D17303, doi: 10.1029/2006JD007548, 2007.

Dorf, M., J.H. Butler, A. Butz, C. Camy-Peyret, M.P. 

Chipperfield, L. Kritten, S.A. Montzka, B. Simmes, 

F. Weidner, and K. Pfeilsticker, Long-term observations of stratospheric bromine reveal slow down 

in growth, Geophys. Res. Lett., 33, L24803, doi: 

10.1029/2006GL027714, 2006.

Dorf, M., A. Butz, C. Camy-Peyret, M.P. Chipperfield, L. 

Kritten, and K. Pfeilsticker, Bromine in the tropical 

troposphere and stratosphere as derived from balloonborne BrO observations, Atmos. Chem. Phys., 8 (23), 

7265-7271, doi: 10.5194/acp-8-7265-2008, 2008.

Douglass, A.R., R.S. Stolarski, M.R. Schoeberl, C.H. 

Jackman, M.L. Gupta, P.A. Newman, J.E. Nielsen, 

and E.L. Fleming, Relationship of loss, mean age 

of air and the distribution of CFCs to stratospheric circulation and implications for atmospheric 

lifetimes, J. Geophys. Res., 113, D14309, doi: 

10.1029/2007JD009575, 2008.

Drewer, J., M.R. Heal, K.V. Heal, and K.A. Smith, Temporal and spatial variation in methyl bromide flux 

from a salt marsh, Geophys. Res. Lett., 33, L16808, 

doi: 10.1029/2006GL026814, 2006.

Drewer, J., K.V. Heal, K.A. Smith, and M.R. Heal, 

Methyl bromide emissions to the atmosphere from 

temperate woodland ecosystems, Global Change 
1.90

Chapter 1

Biol., 14 (11), 2539-2547, doi: 10.1111/j.1365-

2486.2008.01676.x, 2008.

Duchatelet, P., E. Mahieu, R. Ruhnke, W. Feng, M. Chipperfield, P. Demoulin, P. Bernath, C.D. Boone, 

K.A. Walker, C. Servais, and O. Flock, An approach to retrieve information on the carbonyl 

fluoride (COF2) vertical distributions above Jungfraujoch by FTIR multi-spectrum multi-window 

fitting, Atmos. Chem. Phys., 9 (22), 9027-9042, doi: 

10.5194/acp-9-9027-2009, 2009.

Dunse, B.L., L.P. Steele, S.R. Wilson, P.J. Fraser, and 

P.B. Krummel, Trace gas emissions from Melbourne, Australia, based on AGAGE observations at Cape Grim, Tasmania, 1995–2000, Atmos. 

Environ., 39 (34), 6334-6344, doi: 10.1016/j.atmosenv.2005.07.014, 2005.

Engel, A., M. Strunk, M. Müller, H.-P. Haase, C. Poss, I. 

Levin, and U. Schmidt, Temporal development of 

total chlorine in the high-latitude stratosphere based 

on reference distributions of mean age derived from 

CO2 and SF6, J. Geophys. Res., 107 (D12), 4136, 

doi: 10.1029/2001JD000584, 2002.

FAO (Food and Agriculture Organization of the United 

Nations), FAOSTAT database, available at: faostat.

fao.org, 2009.

Feng, W., M.P. Chipperfield, M. Dorf, K. Pfeilsticker, 

and P. Ricaud, Mid-latitude ozone changes: Studies 

with a 3-D CTM forced by ERA-40 analyses, Atmos. Chem. Phys., 7 (9), 2357-2369, doi: 10.5194/

acp-7-2357-2007, 2007.

Ferretti, D.F., J.B. Miller, J.W.C. White, K.R. Lassey, 

D.C. Lowe, and D.M. Etheridge, Stable isotopes 

provide revised global limits of aerobic methane 

emissions from plants, Atmos. Chem. Phys., 7 (1), 

237-241, doi: 10.5194/acp-7-237-2007, 2007.

Fisher, B., and N. Nakicenovic, (Co-ordinating Lead 

Authors), K. Alfsen, J. Corfee Morlot, F. de la 

Chesnaye, J.-Ch. Hourcade, K. Jiang, M. Kainuma, 

E. La Rovere, A. Matysek, A. Rana, K. Riahi, R. 

Richels, S. Rose, D. van Vuuren, and R. Warren, 

Issues related to mitigation in the long term context, 

Chapter 3 in Climate Change 2007: Mitigation. 

Contribution of Working Group III to the Fourth 

Assessment Report of the Intergovernmental Panel 

on Climate Change, edited by B. Metz, O.R. Davidson, P.R. Bosch, R. Dave, and L.A. Meyer, Cambridge University Press, Cambridge, U.K. and New 

York, NY, U.S.A., 2007.

Forster, P., V. Ramaswamy, (Coordinating Lead Authors), P. Artaxo, T. Berntsen, R. Betts, D.W. Fahey, J. Haywood, J. Lean, D.C. Lowe, G. Myhre, 

J. Nganga, R. Prinn, G. Raga, M. Schulz, and R. 

Van Dorland, Changes in atmospheric constituents and in radiative forcing, Chapter 2 in Climate 

Change 2007: The Physical Science Basis. Contribution of Working Group I to the Fourth Assessment Report of the Intergovernmental Panel on 

Climate Change, edited by S. Solomon, D. Qin, 

M. Manning, Z. Chen, M. Marquis, K.B. Averyt, 

M. Tignor, and H.L. Miller, 996 pp., Cambridge 

University Press, Cambridge, U.K., and New York, 

NY, U.S.A., 2007.

Frank, H., E.H. Christoph, O. Holm-Hansen, and J.L. 

Bullister, Trifluoroacetate in ocean waters, Environ. Sci. Technol., 36 (1), 12-15, doi: 10.1021/

es0101532, 2002.

Frankenberg, C., J.F. Meirink, M. van Weele, U. Platt, and 

T. Wagner, Assessing methane emissions from global space-borne observations, Science, 308 (5724), 

1010-1014, doi: 10.1126/science.1106644, 2005.

Frankenberg, C., P. Bergamaschi, A. Butz, S. Houweling, 

J.F. Meirink, J. Notholt, A.K. Petersen, H. Schrijver, T. Warneke, and I. Aben, Tropical methane 

emissions: A revised view from SCIAMACHY onboard ENVISAT, Geophys. Res. Lett., 35, L15811, 

doi: 10.1029/2008GL034300, 2008.

Fraser, P.J., D.E. Oram, C.E. Reeves, S.A. Penkett, and 

A. McCulloch, Southern Hemispheric halon 

trends (1978–1998) and global halon emissions, 

J. Geophys. Res., 104 (D13), 15985-15999, doi: 

10.1029/1999JD900113, 1999.

Friedlingstein, P., P. Cox, R. Betts, L. Bopp, W. von Bloh, 

V. Brovkin, P. Cadule, S. Doney, M. Eby, I. Fung, 

G. Bala, J. John, C. Jones, F. Joos, T. Kato, M. 

Kawamiya, W. Knorr, K. Lindsay, H.D. Matthews, 

T. Raddatz, P. Rayner, C. Reick, E. Roeckner, K.-G. 

Schnitzler, R. Schnur, K. Strassmann, A.J. Weaver, 

C. Yoshikawa, and N. Zeng, Climate–carbon cycle 

feedback analysis: Results from the (C4MIP) model 

intercomparison, J. Clim., 19 (14), 3337-3353, doi: 

10.1175/JCLI3800.1, 2006.

Frische, M., K. Garofalo, T.H. Hansteen, R. Borchers, and 

J. Harnisch, The origin of stable halogenated compounds in volcanic gases, Environ. Sci. Pollut. Res.,

13 (6), 406-413, doi: 10.1065/espr2006.01.291, 2006.

Froidevaux, L., J.W. Waters, W.G. Read, P.S. Connell, 

D.E. Kinnison, and J.M. Russell III, Variations in 

the free chlorine content of the stratosphere (1991–

1997): Anthropogenic, volcanic, and methane influences, J. Geophys. Res., 105 (D4), 4471-4481, 2000.

Froidevaux, L., N.J. Livesey, W.G. Read, R.J. Salawitch, J.W. Waters, B. Drouin, I.A. MacKenzie, 

H.C. Pumphrey, P. Bernath, C. Boone, R. Nassar, S. Montzka, J. Elkins, D. Cunnold, and D. 

Waugh, Temporal decrease in upper atmospheric 

chlorine, Geophys. Res. Lett., 33, L23812, doi: 

10.1029/2006GL027600, 2006.

Froidevaux, L., Y.B. Jiang, A. Lambert, N.J. Livesey, 
1.91

ODSs and Related Chemicals

W.G. Read, J.W. Waters, R.A. Fuller, T.P. Marcy, 

P.J. Popp, R.S. Gao, D.W. Fahey, K.W. Jucks, 

R.A. Stachnik, G.C. Toon, L.E. Christensen, C.R. 

Webster, P.F. Bernath, C.D. Boone, K.A. Walker, 

H.C. Pumphrey, R.S. Harwood, G.L. Manney, M.J. 

Schwartz, W.H. Daffer, B.J. Drouin, R.E. Cofield, 

D.T. Cuddy, R.F. Jarnot, B.W. Knosp, V.S. Perun, W.V. Snyder, P.C. Stek, R.P. Thurstans, and 

P.A. Wagner, Validation of Aura Microwave Limb 

Sounder HCl measurements, J. Geophys. Res., 113, 

D15S25, doi: 10.1029/2007JD009025, 2008.

Fromm, M., E.P. Shettle, K.H. Fricke, C. Ritter, T. Trickl, 

H. Giehl, M. Gerding, J.E. Barnes, M. O’Neill, S.T. 

Massie, U. Blum, I.S. McDermid, T. Leblanc, and 

T. Deshler, Stratospheric impact of the Chisholm 

pyrocumulonimbus eruption: 2. Vertical profile 

perspective, J. Geophys. Res., 113, D08203, doi: 

10.1029/2007JD009147, 2008.

Fu, D., C.D. Boone, P.F. Bernath, K.A. Walker, R. Nassar, 

G.L. Manney, and S.D. McLeod, Global phosgene 

observations from the Atmospheric Chemistry Experiment (ACE) mission, Geophys. Res. Lett., 34

(17), L17815, doi: 10.1029/2007GL029942, 2007.

Fu, D., C.D. Boone, P.F. Bernath, D.K. Weisenstein, C.P. 

Rinsland, G.L. Manney, and K.A. Walker, First 

global observations of atmospheric COClF from 

the Atmospheric Chemistry Experiment mission, J. 

Quant. Spectrosc. Radiat. Transfer, 110 (12), 974-

985, doi: 10.1016/j.jqsrt.2009.02.018, 2009.

Fueglistaler, S., M. Bonazzola, P.H. Haynes, and T. Peter, Stratospheric water vapor predicted from the 

Lagrangian temperature history of air entering the 

stratosphere in the tropics, J. Geophys. Res., 110,

D08107, doi: 10.1029/2004JD005516, 2005.

Fueglistaler, S., A.E. Dessler, T.J. Dunkerton, I. 

Folkins, Q. Fu, and P.W. Mote, Tropical tropopause layer, Rev. Geophys., 47, RG1004, doi: 

10.1029/2008RG000267, 2009a.

Fueglistaler, S., B. Legras, A. Beljaars, J.-J. Morcrette, 

A. Simmons, A.M. Tompkins, and S. Uppala, The 

diabatic heat budget of the upper troposphere and 

lower/mid stratosphere in ECMWF reanalyses, 

Quart. J. Roy. Meteorol. Soc., 135 (638), 21-37, 

doi: 10.1002/qj.361, 2009b.

Gan, J., S.R. Yates, H.D. Ohr, and J.J. Sims, Production 

of methyl bromide by terrestrial higher plants, 

Geophys. Res. Lett., 25 (19), 3595-3598, doi: 

10.1029/98GL52697, 1998.

Gardiner, T., A. Forbes, M. de Mazière, C. Vigouroux, E. 

Mahieu, P. Demoulin, V. Velazco, J. Notholt, T. 

Blumenstock, F. Hase, I. Kramer, R. Sussmann, W. 

Stremme, J. Mellqvist, A. Strandberg, K. Ellingsen, 

and M. Gauss, Trend analysis of greenhouse gases 

over Europe measured by a network of groundbased remote FTIR instruments, Atmos. Chem. 

Phys., 8 (22), 6719-6727, doi: 10.5194/acp-8-6719-

2008, 2008.

Gebhardt, S., A. Colomb, R. Hofmann, J. Williams, and 

J. Lelieveld, Halogenated organic species over the 

tropical South American rainforest, Atmos. Chem. 

Phys., 8 (12), 3185-3197, doi: 10.5194/acp-8-3185-

2008, 2008.

Geller, L.S., J.W. Elkins, J.M. Lobert, A.D. Clarke, D.F. 

Hurst, J.H. Butler, and R.C. Meyer, Troposheric 

SF6: Observed latitudinal distribution and trends, 

derived emissions and interhemispheric exchange 

time, Geophys. Res. Lett., 24 (6), 675-678, doi: 

10.1029/97GL00523, 1997.

Gettelman, A., P.H. Lauritzen, M. Park, and J.E. Kay, Processes regulating short-lived species in the tropical 

tropopause layer, J. Geophys. Res., 114, D13303, 

doi: 10.1029/2009JD011785, 2009.

Gossé, S., L. Hespel, P. Gossart, and A. Delfour, Morphological characterization and particle sizing of 

alumina particles in solid rocket motor, J. Propul. 

Power, 22 (1), 127-135, 2006.

Greally, B.R., P.G. Simmonds, S. O’Doherty, A. McCulloch, B.R. Miller, P.K. Salameh, J. Mühle, T. Tanhua, C. Harth, R.F. Weiss, P.J. Fraser, P.B. Krummel, B.L. Dunse, L.W. Porter, and R.G. Prinn, 

Improved continuous in situ measurements of C1–

C3 PFCs, HFCs, HCFCs, CFCs and SF6 in Europe 

and Australia, J. Integr. Environ. Sci., 2 (2-3), 253-

261, doi: 10.1080/15693430500402614, 2005.

Greally, B.R., A.J. Manning, S. Reimann, A. McCulloch, 

J. Huang, B.L. Dunse, P.G. Simmonds, R.G. Prinn, 

P.J. Fraser, D.M. Cunnold, S. O’Doherty, L.W. 

Porter, K. Stemmler, M.K. Vollmer, C.R. Lunder, 

N. Schmidbauer, O. Hermansen, J. Arduini, P.K. 

Salameh, P.B. Krummel, R.H.J. Wang, D. Folini, 

R.F. Weiss, M. Maione, G. Nickless, F. Stordal, and 

R.G. Derwent, Observations of 1,1-difluoroethane 

(HFC-152a) at AGAGE and SOGE monitoring stations in 1994–2004 and derived global and regional 

emission estimates, J. Geophys. Res., 112, D06308, 

doi: 10.1029/2006JD007527, 2007.

Guo, H., A.J. Ding, T. Wang, I.J. Simpson, D.R. Blake, 

B. Barletta, S. Meinardi, F.S. Rowland, S.M. Saunders, T.M. Fu, W.T. Hung, and Y.S. Li, Source 

origins, modeled profiles, and apportionments of 

halogenated hydrocarbons in the greater Pearl River Delta region, southern China, J. Geophys. Res.,

114, D11302, doi: 10.1029/2008JD011448, 2009.

Guo, S., G.J.S. Bluth, W.I. Rose, I.M. Watson, and A.J. 

Prata, Re-evaluation of SO2 release of the 15 June 

1991 Pinatubo eruption using ultraviolet and infrared satellite sensors, Geochem. Geophys. Geosyst., 

5, Q04001, doi: 10.1029/2003GC000654, 2004.
1.92

Chapter 1

Hall, B.D., G.S. Dutton, and J.W. Elkins, The NOAA 

nitrous oxide standard scale for atmospheric observations, J. Geophys. Res., 112, D09305, doi: 

10.1029/2006JD007954, 2007.

Hamilton, J.T.G., W.C. McRoberts, F. Keppler, R.M. 

Kalin, and D.B Harper, Chloride methylation by 

plant pectin: An efficient environmentally significant process, Science, 301 (5630), 206-209, doi: 

10.1126/science.105036, 2003.

Happell, J.D., and D.W.R. Wallace, Removal of atmospheric CCl4 under bulk aerobic conditions in 

groundwater and soils, Environ. Sci. Technol., 32

(9), 1244-1252, doi: 10.1021/es970653o, 1998.

Happell, J.D., and M.P. Roche, Soils: A global sink of atmospheric carbon tetrachloride, Geophys. Res. Lett.,

30 (2), 1088, doi: 10.1029/2002GL015957, 2003.

Hardacre, C.J., E. Blei, and M.R. Heal, Growing season 

methyl bromide and methyl chloride fluxes at a 

sub-arctic wetland in Sweden, Geophys. Res. Lett.,

36, L12401, doi: 10.1029/2009GL038277, 2009.

Harnisch, J., Reactive fluorine compounds, Chapter 3 in 

The Handbook of Environmental Chemistry, Volume 4E: Reactive Halogen Compounds in the Atmosphere, edited by P. Fabian, and O.N. Singh, 

81-111, Springer-Verlag, Berlin Heidelberg, 1999.

Harnisch, J., Atmospheric perfluorocarbons: Sources and 

concentrations, in Non-CO2 Greenhouse Gases: 

Scientific Understanding, Control and Implementation, edited by J. van Ham, A.P.M. Baede, L.A. 

Meyer, and R. Ybema, 205-210, Kluwer Academic 

Publishers, Dordrecht, Netherlands, 2000.

Harnisch, J., R. Borchers, P. Fabian, H.W. Gäggeler, and U. 

Schotterer, Effect of natural tetrafluoromethane, Nature, 384 (6604), 32, doi: 10.1038/384032a0, 1996.

Haywood, J., and O. Boucher, Estimates of the direct and indirect radiative forcing due to tropospheric aerosols: 

A review, Rev. Geophys., 38 (4), 513-543, 2000.

Hegglin, M.I., D. Brunner, H. Wernli, C. Schwierz, O. 

Martius, P. Hoor, H. Fischer, U. Parchatka, N. 

Spelten, C. Schiller, M. Krebsbach, U. Weers, J. 

Staehelin, and Th. Peter, Tracing troposphere-tostratosphere transport above a mid-latitude deep 

convective system, Atmos. Chem. Phys., 4 (3), 741-

756, doi: 10.5194/acp-4-741-2004, 2004.

Hendrick, F., P.V. Johnston, M. De Mazière, C. Fayt, C. 

Hermans, K. Kreher, N. Theys, A. Thomas, and 

M. Van Roozendael, One-decade trend analysis of 

stratospheric BrO over Harestua (60°N) and Lauder 

(45°S) reveals a decline, Geophys. Res. Lett., 35, 

L14801, doi: 10.1029/2008GL034154, 2008.

Hense, I., and B. Quack, Modelling the vertical distribution of bromoform in the upper water column of the 

tropical Atlantic Ocean, Biogeosciences, 6 (4), 535-

544, doi: 10.5194/bg-6-535-2009, 2009.

Hirsch, A.I., A.M. Michalak, L.M. Bruhwiler, W. Peters, 

E.J. Dlugokencky, and P.P. Tans, Inverse modeling 

estimates of the global nitrous oxide surface flux 

from 1998–2001, Global Biogeochem. Cycles, 20, 

GB1008, doi: 10.1029/2004GB002443, 2006.

Hoffmann, L., M. Kaufmann, R. Spang, R. Müller, J.J. Remedios, D.P. Moore, C.M. Volk, T. von Clarmann, 

and M. Riese, Envisat MIPAS measurements of 

CFC-11: Retrieval, validation, and climatology, Atmos. Chem. Phys., 8 (13), 3671-3688, doi: 10.5194/

acp-8-3671-2008, 2008.

Hofmann, D.J., and S.A. Montzka, Recovery of the ozone 

layer: The Ozone Depleting Gas Index, EOS Transactions, 90 (1), 1-2, doi: 10.1029/2009EO010001, 2009.

Hofmann, D.J., J.H. Butler, E.J. Dlugokencky, J.W. Elkins, 

K. Masarie, S.A. Montzka, and P. Tans, The role 

of carbon dioxide in climate forcing from 1979 to 

2004: Introduction of the Annual Greenhouse Gas 

Index, Tellus, 58B, 614-619, doi: 10.1111/j.1600-

0889.2006.00201.x, 2006.

Hofmann, D., J. Barnes, M. O’Neill, M. Trudeau, and R. 

Neely, Increase in background stratospheric aerosol observed with lidar at Mauna Loa Observatory 

and Boulder, Colorado, Geophys. Res. Lett., 36, 

L15808, doi: 10.1029/2009GL039008, 2009.

Homan, C.D., C.M. Volk, A.C. Kuhn, A. Werner, J. Baehr, 

S. Viciani, A. Ulanovski, and F. Ravegnani, Tracer 

measurements in the tropical tropopause layer during the AMMA/SCOUT-O3 aircraft campaign, Atmos. Chem. Phys., 10 (8), 3615-3627, doi: 10.5194/

acp-10-3615-2010, 2010.

Hoor, P., C. Gurk, D. Brunner, M.I. Hegglin, H. Wernli, 

and H. Fischer, Seasonality and extent of extratropical TST derived from in-situ CO measurements 

during SPURT, Atmos. Chem. Phys., 4 (5), 1427–

1442, doi: 10.5194/acp-4-1427-2004, 2004.

Höpfner, M., J. Orphal, T. von Clarmann, G. Stiller, and 

H. Fischer, Stratospheric BrONO2 observed by 

MIPAS, Atmos. Chem. Phys., 9 (5), 1735-1746, 

doi: 10.5194/acp-9-1735-2009, 2009.

Hossaini, R., M.P. Chipperfield, B.M. Monge-Sanz, 

N.A.D. Richards, E. Atlas, and D.R. Blake, Bromoform and dibromomethane in the tropics: A 3-D 

model study of chemistry and transport, Atmos. 

Chem. Phys., 10 (2), 719-735, doi: 10.5194/acp-10-

719-2010, 2010.

Hu, L., S.A. Yvon-Lewis, Y. Liu, J.E. Salisbury, and 

J.E. O’Hern, Coastal emissions of methyl bromide and methyl chloride along the eastern Gulf 

of Mexico and the east coast of the United States, 

Global Biogeochem. Cycles, 24, GB1007, doi: 

10.1029/2009GB003514, 2010.

Huang, J., A. Golombek, R. Prinn, R. Weiss, P. Fraser, 

P. Simmonds, E.J. Dlugokencky, B. Hall, J. Elkins, 
1.93

ODSs and Related Chemicals

P. Steele, R. Langenfelds, P. Krummel, G. Dutton, 

and L. Porter, Estimation of regional emissions of 

nitrous oxide from 1997 to 2005 using multinetwork measurements, a chemical transport model, 

and an inverse method, J. Geophys. Res., 113, 

D17313, doi: 10.1029/2007JD009381, 2008.

Huhn, O., W. Roether, P. Beining, and H. Rose, Validity limits of carbon tetrachloride as an ocean tracer, 

Deep-Sea Res., 1, 48 (9), 2025-2049, doi: 10.1016/

S0967-0637(01)00004-8, 2001.

Hurley, M.D., J.C. Ball, T.J. Wallington, M.P. Sulbaek 

Andersen, O.J. Nielsen, D.A. Ellis, J.W. Martin, 

and S.A. Mabury, Atmospheric chemistry of nCxF2x+1CHO (x = 1, 2, 3, 4): Fate of n-CxF2x+1C(O) 

radicals, J. Phys. Chem. A, 110 (45), 12443-12447, 

doi: 10.1021/jp064029m, 2006.

Hurley, M.D., T.J. Wallington, M.S. Javadi, and O.J. 

Nielsen, Atmospheric chemistry of CF3CF=CH2: 

Products and mechanisms of Cl atom and OH radical initiated oxidation, Chem. Phys. Lett., 450 (4-6), 

263-267, doi: 10.1016/j.cplett.2007.11.051, 2008.

Hurst, D.F., J.C. Lin, P.A. Romashkin, B.C. Daube, C. 

Gerbig, D.M. Matross, S.C. Wofsy, B.D. Hall, and 

J.W. Elkins, Continuing global significance of emissions of Montreal Protocol-restricted halocarbons 

in the United States and Canada, J. Geophys. Res.,

111, D15302, doi: 10.1029/2005JD006785, 2006.

Inoue, Y., M. Kawasaki, T.J. Wallington, and M.D. Hurley, 

Atmospheric chemistry of CF3CH2CF2CH3 (HFC365mfc): Kinetics and mechanism of chlorine atom 

initiated oxidation, infrared spectrum, and global 

warming potential, Chem. Phys. Lett., 462, 4-6, 

164-168, doi: 10.1016/j.cplett.2008.07.054, 2008.

IPCC/TEAP (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, prepared by Working Groups I 

and III of the Intergovernmental Panel on Climate 

Change, and the Technical and Economic Assessment Panel, Cambridge University Press, Cambridge, U.K. and New York, NY, U.S.A., 2005.

James, R., M. Bonazzola, B. Legras, K. Surbled, and S. 

Fueglistaler, Water vapor transport and dehydration above convective outflow during Asian 

monsoon, Geophys. Res. Lett., 35, L20810, doi: 

10.1029/2008GL035441, 2008.

Javadi, M.S., R. Søndergaard, O.J. Nielsen, M.D. Hurley, 

and T.J. Wallington, Atmospheric chemistry of 

trans-CF3CH=CHF: Products and mechanisms of 

hydroxyl radical and chlorine atom initiated oxidation, Atmos. Chem. Phys., 8 (12), 3141-3147, doi: 

10.5194/acp-8-3141-2008, 2008.

Jenkin, M.E., Photochemical Ozone and PAN Creation 

Potentials: Rationalisation and Methods of Estimation, AEA Technology plc, AEAT Report 

4182/20150/003, Issue 1, National Environmental 

Technology Centre, U.K., 1998.

Jiang, X., W.L. Ku, R.-L. Shia, Q. Li, J.W. Elkins, R.G. 

Prinn, and Y.L. Yung, Seasonal cycle of N2O: 

Analysis of data, Global Biogeochem. Cycles, 21, 

GB1006, doi: 10.1029/2006GB002691, 2007.

Jourdain, J.L., G. Le Bras, and J. Combourieu, Kinetic 

study by electron paramagnetic resonance and 

mass spectrometry of the elementary reactions of 

phosphorus tribromide with H, O, and OH radicals, 

J. Phys. Chem., 86 (21), 4170-4175, doi: 10.1021/

j100218a016, 1982.

JPL 06-2: see Sander et al., 2006.

JPL 10-6: see Sander et al., 2010.

Kaminski, T., P.J. Rayner, M. Heimann, and I.G. Enting, 

On aggregation errors in atmospheric transport inversions, J. Geophys. Res., 106 (D5), 4703-4715, 

doi: 10.1029/2000JD900581, 2001.

Keene, W.C., M.A.K. Khalil, D.J. Erickson, A. McCulloch, T.E. Graedel, J.M. Lobert, M.L. Aucott, S.L. 

Gong, D.B. Harper, G. Kleiman, P. Midgley, R.M. 

Moore, C. Seuzaret, W.T. Sturges, C.M. Benkovitz, 

V. Koropalov, L.A. Barrie, and Y.F. Li, Composite global emissions of reactive chlorine from anthropogenic and natural sources: Reactive Chlorine 

Emissions Inventory, J. Geophys. Res., 104 (D7), 

8429-8440, doi: 10.1029/1998JD100084, 1999.

Keppler, F., R. Eiden, V. Niedan, J. Pracht, and H.F. 

Schöler, Halocarbons produced by natural oxidation 

processes during degradation of organic matter, Nature, 403, 298-301, doi: 10.1038/35002055, 2000.

Keppler, F., R.M. Kalin, D.B. Harper, W.C. McRoberts, 

and J.T.G. Hamilton, Carbon isotope anomaly in 

the major plant C1 pool and its global biogeochemical implications, Biogeosciences, 1 (2), 123-131, 

doi: 10.5194/bg-1-123-2004, 2004.

Keppler, F., D.B. Harper, T. Röckmann, R.M. Moore, and 

J.T.G. Hamilton, New insight into the atmospheric 

chloromethane budget gained using stable carbon 

isotope ratios, Atmos. Chem. Phys., 5 (9), 2403-

2411, doi: 10.5194/acp-5-2403-2005, 2005.

Keppler, F., J.T.G. Hamilton, M. Braß, and T. Röckmann, 

Methane emissions from terrestrial plants under 

aerobic conditions, Nature, 439, 187-191, doi: 

10.1038/nature04420, 2006.

Kerkweg, A., P. Jöckel, N. Warwick, S. Gebhardt, C.A.M. 

Brenninkmeijer, and J. Lelieveld, Consistent 

simulation of bromine chemistry from the marine 

boundary layer to the stratosphere – Part 2: Bromocarbons, Atmos. Chem. Phys., 8 (19), 5919-5939, 

doi: 10.5194/acp-8-5919-2008, 2008.
1.94

Chapter 1

Kim, J., S. Li, K.-R. Kim, A. Stohl, J. Mühle, S.-K. Kim, 

M.-K. Park, D.-J. Kang, G. Lee, C.M. Harth, 

P.K. Salameh, and R.F. Weiss, Regional atmospheric emissions determined from measurements 

at Jeju Island, Korea: Halogenated compounds 

from China, Geophys. Res. Lett., 37, L12801, doi: 

10.1029/2010GL043263, 2010.

Kindler, T.P., W.L. Chameides, P.H. Wine, D.M. Cunnold, F.N. Alyea, and J.A. Franklin, The fate of 

atmospheric phosgene and the stratospheric loadings of its parent compounds: CCl4, C2Cl4, C2HCl3, 

CH3CCl3, and CHCl3, J. Geophys. Res., 100 (D1), 

1235-1251, doi: 10.1029/94JD02518, 1995.

King, D.B., J.H. Butler, S.A. Yvon-Lewis, and S.A. Cotton, Predicting oceanic methyl bromide saturation 

from SST, Geophys. Res. Lett., 29 (24), 2199, doi: 

10.1029/2002GL016091, 2002.

Klimont, Z., J. Cofala, J. Xing, W. Wei, C. Zhang, S. 

Wang, J. Kejun, P. Bhandari, R. Mathur, P. Purohit, P. Rafaj, A. Chambers, M. Amann, and J. Hao, 

Projections of SO2, NOx and carbonaceous aerosols 

emissions in Asia, Tellus, 61B (4), 602-617, doi: 

10.1111/j.1600-0889.2009.00428.x, 2009.

Kloster, S., K.D. Six, J. Feichter, E. Maier-Reimer, E. 

Roeckner, P. Wetzel, P. Stier, and M. Esch, Response of dimethylsulfide (DMS) in the ocean and 

atmosphere to global warming, J. Geophys. Res., 

112, G03005, doi: 10.1029/2006JG000224, 2007.

Knorr, W., Is the airborne fraction of anthropogenic CO2

emissions increasing?, Geophys. Res. Lett., 36, 

L21710, doi: 10.1029/2009GL040613, 2009.

Ko, M.K.W., and G. Poulet (Lead Authors), D.R. Blake, 

O. Boucher, J.H. Burkholder, M. Chin, R.A. Cox, 

C. George, H.-F. Graf, J.R. Holton, D.J. Jacob, K.S. 

Law, M.G. Lawrence, P.M. Midgley, P.W. Seakins, 

D.E. Shallcross, S.E. Strahan, D.J. Wuebbles, and 

Y. Yokouchi, Very short-lived halogen and sulfur 

substances, Chapter 2 in Scientific Assessment of 

Ozone Depletion: 2002, Global Ozone Research and 

Monitoring Project–Report No. 47, World Meteorological Organization, Geneva, Switzerland, 2003.

Kovalenko, L.J., N.L. Livesey, R.J. Salawitch, C. CamyPeyret, M.P. Chipperfield, R.E. Cofield, M. Dorf, 

B.J. Drouin, L. Froidevaux, R.A. Fuller, F. Goutail, 

R.F. Jarnot, K. Jucks, B.W. Knosp, A. Lambert, I.A. 

MacKenzie, K. Pfeilsticker, J.-P. Pommereau, W.G. 

Read, M.L. Santee, M.J. Schwartz, W.V. Snyder, R. 

Stachnik, P.C. Stek, P.A. Wagner, and J.W. Waters, 

Validation of Aura Microwave Limb Sounder BrO 

observations in the stratosphere, J. Geophys. Res.,

112, D24S41, doi: 10.1029/2007JD008817, 2007.

Kremser, S., I. Wohltmann, M. Rex, U. Langematz, M. 

Dameris, and I. Wohltmann, Water vapour transport in the tropical tropopause region in coupled 

Chemistry-Climate Models and ERA-40 reanalysis 

data, Atmos. Chem. Phys., 9 (8), 2679-2694, doi: 

10.5194/acp-9-2679-2009, 2009.

Krüger, K., S. Tegtmeier, and M. Rex, Long-term climatology of air mass transport through the Tropical 

Tropopause Layer (TTL) during NH winter, Atmos. 

Chem. Phys., 8 (4), 813-823, doi: 10.5194/acp-8-

813-2008, 2008.

Krüger, K., S. Tegtmeier, and M. Rex, Variability of residence time in the Tropical Tropopause Layer during Northern Hemisphere winter, Atmos. Chem. 

Phys., 9 (18), 6717-6725, doi: 10.5194/acp-9-6717-

2009, 2009.

Kuell, V., D. Offermann, M. Jarisch, B. Schaeler, A. Engel, H. Claude, H.G.J. Smit, A. Ebel, and H. Feldmann, Tropopause region temperatures and CFC 11 

mixing ratios from CRISTA 2, J. Geophys. Res.,

110, D16104, doi: 10.1029/2004JD005592, 2005.

Kutsuna, S., L. Chen, T. Abe, J. Mizukado, T. Uchimaru, 

K. Tokuhashi, and A. Sekiya, Henry’s law constants of 2,2,2-trifluoroethyl formate, ethyl trifluoroacetate, and non-fluorinated analogous esters, Atmos. Environ., 39 (32), 5884-5892, doi: 10.1016/j.

atmosenv.2005.06.021, 2005.

Langbein, T., H. Sonntag, D. Trapp, A. Hoffmann, W. 

Malms, E.P. Roth, V. Mors, and R. Zellner, Volatile anaesthetics and the atmosphere: Atmospheric 

lifetimes and atmospheric effects of halothane, 

enflurane, isoflurane, desflurane and sevoflurane, 

British J. Anaesthesia, 82 (1), 66-73, 1999.

Laube, J.C., A. Engel, H. Bönisch, T. Möbius, D.R. Worton, W.T. Sturges, K. Grunow, and U. Schmidt, 

Contribution of very short-lived organic substances 

to stratospheric chlorine and bromine in the tropics—a case study, Atmos. Chem. Phys., 8 (23), 

7325-7334, doi: 10.5194/acp-8-7325-2008, 2008.

Laube, J.C., P. Martinerie, E. Witrant, T. Blunier, J. 

Schwander, C.A.M Brenninkmeijer, T.J. Schuck, 

M. Bolder, T. Röckmann, C. van der Veen, H. 

Bönisch, A. Engel, G.P. Mills, M.J. Newland, D.E. 

Oram, C.E. Reeves, and W.T. Sturges, Accelerating growth of HFC-227ea (1,1,1,2,3,3,3-heptafluoropropane) in the atmosphere, Atmos. Chem. Phys.,

10 (13), 5903-5910, doi: 10.5194/acp-10-5903-

2010, 2010.

Law, K.S., and W.T. Sturges (Lead Authors), 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, and S.A. Yvon-Lewis, Halogenated very short-lived substances, Chapter 2 in 

Scientific Assessment of Ozone Depletion: 2006,

Global Ozone Research and Monitoring Project—
1.95

ODSs and Related Chemicals

Report No.50, World Meteorological Organization, 

Geneva, Switzerland, 2007.

Le Calvé, S., G. Le Bras, and A. Mellouki, Temperature 

dependence for the rate coefficients of the reactions of the OH radical with a series of formates, J. 

Phys. Chem. A, 101 (30), 5489-5493, doi: 10.1021/

jp970554x, 1997.

Lee, B.-S., J.L. Bullister, and F.A. Whitney, Chlorofluorocarbon CFC-11 and carbon tetrachloride removal in 

Saanich Inlet, an intermittently anoxic basin, Mar. 

Chem., 66 (3-4), 171-185, doi: 10.1016/S0304-

4203(99)00039-0, 1999.

Lee-Taylor, J.M., and E.A. Holland, Litter decomposition as a potential natural source of methyl bromide, J. Geophys. Res., 105 (D7), 8857-8864, doi: 

10.1029/1999JD901112, 2000.

Lee-Taylor, J., and K.R. Redeker, Reevaluation of global 

emissions from rice paddies of methyl iodide and 

other species, Geophys. Res. Lett., 32, L15801, doi: 

10.1029/2005GL022918, 2005.

Leser, H., G. Hönninger, and U. Platt, MAX-DOAS measurements of BrO and NO2 in the marine boundary layer, Geophys. Res. Lett., 30 (10), 1537, doi: 

10.1029/2002GL015811, 2003.

Levin, I., T. Naegler, R. Heinz, D. Osusko, E. Cuevas, 

A. Engel, J. Ilmberger, R.L. Langenfelds, B. Neininger, C. v. Rohden, L.P. Steele, R. Weller, D.E. 

Worthy, and S.A. Zimov, The global SF6 source inferred from long-term high precision atmospheric 

measurements and its comparison with emission inventories, Atmos. Chem. Phys., 10 (6), 2655-2662, 

doi: 10.5194/acp-10-2655-2010, 2010.

Levine, J.G., P. Braesicke, N.R.P. Harris, and J.A. Pyle, 

Seasonal and inter-annual variations in troposphere-to-stratosphere transport from the tropical tropopause layer, Atmos. Chem. Phys., 8 (13), 

3689-3703, doi: 10.5194/acp-8-3689-2008, 2008.

Li, Y., K.O. Patten, D. Youn, and D.J. Wuebbles, Potential 

impacts of CF3I on ozone as a replacement for CF3Br 

in aircraft applications, Atmos. Chem. Phys., 6 (12), 

4559-4568, doi: 10.5194/acp-6-4559-2006, 2006.

Liang, Q., R.S. Stolarski, S.R. Kawa, J.E. Nielsen, J.M. 

Rodriguez, A.R. Douglass, J.M. Rodriguez, D.R. 

Blake, E.L. Atlas, and L.E. Ott, Finding the missing stratospheric Bry: A global modeling study of 

CHBr3 and CH2Br2, Atmos. Chem. Phys., 10 (5), 

2269-2286, doi: 10.5194/acp-10-2269-2010, 2010.

Liu, A.Q., and G.W.K. Moore, Lake-effect snowstorms 

over southern Ontario, Canada, and their associated 

synoptic-scale environment, Mon. Wea. Rev., 132

(11), 2595-2609, doi: 10.1175/MWR2796.1, 2004.

Liu, C., and E.J. Zipser, Global distribution of convection 

penetrating the tropical tropopause, J. Geophys. Res.,

110, D23104, doi: 10.1029/2005JD006063, 2005.

Liu, C., E.J. Zipser, and S.W. Nesbitt, Global distribution 

of tropical deep convection: Different perspectives 

from TRMM infrared and radar data, J. Clim., 20

(3), 489-503, doi: 10.1175/JCLI4023.1, 2007.

Liu, X.-F., Evidence of Biodegradation of Atmospheric 

Carbon Tetrachloride in Soils: Field and Microcosm Studies, Ph.D. thesis, 139 pp., Columbia University, New York, NY, U.S.A., 2006.

Livesey, N.J., L.J. Kovalenko, R.J. Salawitch, I.A. MacKenzie, M.P. Chipperfield, W.G. Read, R.F. Jarnot, 

and J.W. Waters, EOS Microwave Limb Sounder 

observations of upper stratospheric BrO: Implications for total bromine, Geophys. Res. Lett., 33, 

L20817, doi: 10.1029/2006GL026930, 2006.

Low, J.C., N.Y. Wang, J. Williams, and R.J. Cicerone, Measurements of ambient atmospheric C2H5Cl and other 

ethyl and methyl halides at coastal California sites 

and over the Pacific Ocean, J. Geophys. Res., 108 

(D19), 4608, doi: 10.1029/2003JD003620, 2003.

Lu, X.-L., G.-P. Yang, G.-S. Song, and L. Zhang, Distributions and fluxes of methyl chloride and methyl 

bromide in the East China Sea and the Southern 

Yellow Sea in autumn, Mar. Chem., 118 (1-2), 75-

84, doi: 10.1016/j.marchem.2009.11.002, 2010.

Luderer, G., J. Trentmann, K. Hungershöfer, M. Herzog, 

M. Fromm, and M.O. Andreae, Small-scale mixing 

processes enhancing troposphere-to-stratosphere 

transport by pyro-cumulonimbus storms, Atmos. 

Chem. Phys., 7 (23), 5945-5957, doi: 10.5194/acp7-5945-2007, 2007.

Luecken, D.J., R.L. Waterland, S. Papasavva, K.N. Taddonio, W.T. Hutzel, J.P. Rugh, and S.O. Andersen, Ozone and TFA impacts in North America 

from degradation of 2,3,3,3-tetrafluoropropene 

(HFO-1234yf), a potential greenhouse gas replacement, Environ. Sci. Technol., 44 (1), 343-348, doi: 

10.1021/es902481f, 2010.

Mahieu, E., P. Duchatelet, R. Zander, P. Demoulin, C. 

Servais, C.P. Rinsland, M.P. Chipperfield, and 

M. De Mazière, The evolution of inorganic chlorine above the Jungfraujoch station: An update, in 

Ozone Vol. II, Proceedings of the XX Quadrennial 

Ozone Symposium, 1-8 June 2004, Kos, Greece, edited by C.S. Zerefos, 997-998, Int. Ozone Comm., 

Athens, Greece, 2004.

Mahieu, E., P. Duchatelet, P. Demoulin, K.A. Walker, 

E. Dupuy, L. Froidevaux, C. Randall, V. Catoire, 

K. Strong, C.D. Boone, P.F. Bernath, J.-F. Blavier, T. Blumenstock, M. Coffey, M. De Mazière, 

D. Griffith, J. Hannigan, F. Hase, N. Jones, K.W. 

Jucks, A. Kagawa, Y. Kasai, Y. Mebarki, S. Mikuteit, R. Nassar, J. Notholt, C.P. Rinsland, C. Robert, 

O. Schrems, C. Senten, D. Smale, J. Taylor, C. Tétard, G.C. Toon, T. Warneke, S.W. Wood, R. Zan-
1.96

Chapter 1

der, and C. Servais, Validation of ACE-FTS v2.2 

measurements of HCl, HF, CCl3F and CCl2F2 using 

space-, balloon- and ground-based instrument observations, Atmos. Chem. Phys., 8 (20), 6199-6221, 

doi: 10.5194/acp-8-6199-2008, 2008.

Maione, M., J. Arduini, G. Mangani, and A. Geniali, 

Evaluation of an automatic sampling gas chromatographic-mass spectrometric instrument for continuous monitoring of trace anthropogenic gases, Int. 

J. Environ. Anal. Chem., 84 (4), 241–253, doi: 

10.1080/03067310310001626740, 2004.

Maione, M., U. Giostra, J. Arduini, L. Belfiore, F. Furlani, A. Geniali, G. Mangani, M.K. Vollmer, and 

S. Reimann, Localization of source regions of selected hydrofluorocarbons combining data collected at two European mountain stations, Sci. Total 

Environ., 391 (2-3), 232-240, doi: 10.1016/j.scitotenv.2007.10.023, 2008.

Makide, Y., and F.S. Rowland, Tropospheric concentrations of methylchloroform, CH3CCl3, in January 

1978 and estimates of the atmospheric residence 

times for hydrohalocarbons, Proc. Natl. Acad. Sci.,

78 (10), 5933-5937, 1981.

Manley, S.L., N.-Y. Wang, M.L. Walser, and R.J. Cicerone, Coastal salt marshes as global methyl halide 

sources from determinations of intrinsic production 

by marsh plants, Global Biogeochem. Cycles, 20, 

GB3015, doi: 10.1029/2005GB002578, 2006.

Manley, S.L., N.-Y. Wang, M.L. Walser, and R.J. Cicerone, Methyl halide emissions from greenhousegrown mangroves, Geophys. Res. Lett., 34, L01806, 

doi: 10.1029/2006GL027777, 2007.

Manning, A.J., D.B. Ryall, R.G. Derwent, P.G. Simmonds, and S. O’Doherty, Estimating European 

emissions of ozone-depleting and greenhouse gases 

using observations and a modeling back-attribution 

technique, J. Geophys. Res., 108 (D14), 4405, doi: 

10.1029/2002JD002312, 2003.

Marcy, T.P., P.J. Popp, R.S. Gao, D.W. Fahey, E.A. Ray, 

E.C. Richard, T.L. Thompson, E.L. Atlas, M. 

Loewenstein, S.C. Wofsy, S. Park, E.M. Weinstock, W.H. Swartz, and M.J. Mahoney, Measurements of trace gases in the tropical tropopause 

layer, Atmos. Environ., 41 (34), 7253-7261, doi: 

10.1016/j.atmosenv.2007.05.032, 2007.

Martin, J.W., S.A. Mabury, C.S. Wong, F. Noventa, K.R. 

Solomon, M. Alaee, and D.C.G. Muir, Airborne 

haloacetic acids, Environ. Sci. Technol., 37 (13), 

2889-2897, doi: 10.1021/es026345u, 2003.

Martin, M., D. Pöhler, K. Seitz, R. Sinreich, and U. Platt, 

BrO measurements over the Eastern North-Atlantic, Atmos. Chem. Phys., 9 (24), 9545-9554, doi: 

10.5194/acp-9-9545-2009, 2009.

Martinerie, P., E. Nourtier-Mazauric, J.-M. Barnola, W.T. 

Sturges, D.R. Worton, E. Atlas, L.K. Gohar, K.P. 

Shine, and G.P. Brasseur, Long-lived halocarbon 

trends and budgets from atmospheric chemistry 

modelling constrained with measurements in polar 

firn, Atmos. Chem. Phys., 9 (12), 3911-3934, doi: 

10.5194/acp-9-3911-2009, 2009.

Martinsson, B.G., H.N. Nguyen, C.A.M. Brenninkmeijer, 

A. Zahn, J. Heintzenberg, M. Hermann, and P.F.J. 

van Velthoven, Characteristics and origin of lowermost stratospheric aerosol at northern midlatitudes 

under volcanically quiescent conditions based on 

CARIBIC observations, J. Geophys. Res., 110, 

D12201, doi: 10.1029/2004JD005644, 2005.

McCulloch, A., M.L. Aucott, T.E. Graedel, G. Kleiman, 

P.M. Midgley, and Y.-F. Li, Industrial emissions 

of trichloroethene, tetrachloroethene, and dichloromethane: Reactive Chlorine Emissions Inventory, J. Geophys. Res., 104 (D7), 8417-8427, doi: 

10.1029/1999JD900011, 1999.

McCulloch, A., Chloroform in the environment: Occurrence, sources, sinks and effects, Chemosphere, 50 (10), 1291-1308, doi: 10.1016/S0045-

6535(02)00697-5, 2003.

McLinden, C.A., C.S. Haley, N.D. Lloyd, F. Hendrick, 

A. Rozanov, B.-M. Sinnhuber, F. Goutail, D.A. 

Degenstein, E.J. Llewellyn, C.E. Sioris, M. Van 

Roozendael, J.P. Pommereau, W. Lotz, and J.P. 

Burrows, Odin/OSIRIS observations of stratospheric BrO: Retrieval methodology, climatology, 

and inferred Bry, J. Geophys. Res., 115, D15308, 

doi: 10.1029/2009JD012488, 2010.

Mead, M.I., I.R. White, G. Nickless, K.-Y. Wang, and D.E. 

Shallcross, An estimation of the global emission of 

methyl bromide from rapeseed (Brassica napus) 

from 1961 to 2003, Atmos. Environ., 42 (2), 337-

345, doi: 10.1016/j.atmosenv.2007.09.020, 2008a.

Mead, M.I., M.A.H. Khan, G. Nickless, B.R. Greally, D. 

Tainton, T. Pitman, and D.E. Shallcross, Leaf cutter 

ants: A possible missing source of biogenic halocarbons, Environ. Chem., 5 (1), 5-10, doi: 10.1071/

EN07068, 2008b.

Mead, M.I., M.A.H. Khan, I.R. White, G. Nickless, and 

D.E. Shallcross, Methyl halide emission estimates 

from domestic biomass burning in Africa, Atmos. 

Environ., 42 (21), 5241-5250, doi: 10.1016/j.atmosenv.2008.02.066, 2008c.

Mébarki, Y., V. Catoire, N. Huret, G. Berthet, C. Robert, 

and G. Poulet, More evidence for very short-lived 

substance contribution to stratospheric chlorine 

inferred from HCl balloon-borne in situ measurements in the tropics, Atmos. Chem. Phys., 10 (2), 

397-409, doi: 10.5194/acp-10-397-2010, 2010.

Meier, R.R., J.M.C. Plane, M.H. Stevens, L.J. Paxton, A.B. Christensen, and G. Crowley, Can mo-
1.97

ODSs and Related Chemicals

lecular diffusion explain Space Shuttle plume 

spreading?, Geophys. Res. Lett., 37, L08101, doi: 

10.1029/2010GL042868, 2010.

Meirink, J.F., P. Bergamaschi, C. Frankenberg, M.T.S. 

d’Amelio, E.J. Dlugokencky, L.V. Gatti, S. Houweling, J.B. Miller, T. Röckmann, M.G. Villani, 

and M.C. Krol, Four-dimensional variational data 

assimilation for inverse modelling of atmospheric 

methane emissions: Analysis of SCIAMACHY 

observations, J. Geophys. Res., 113, D17301, doi: 

10.1029/2007JD009740, 2008.

Mikuteit, S., Trendbestimmung stratosphärischer Spurengase mit Hilfebodengebundener FTIR-Messungen, 

dissertation, FZK Report No. 7385, Forschungszentrum Karlsruhe, Germany, available: http://

www-imk.fzk.de/asf/ftir/refs.htm, 2008.

Miller, B.R., R.F. Weiss, P.K. Salameh, T. Tanhua, B.R. 

Greally, J. Mühle, and P.G. Simmonds, Medusa: A 

sample preconcentration and GC/MS detector system for in situ measurements of atmospheric trace 

halocarbons, hydrocarbons, and sulfur compounds, 

Anal. Chem., 80 (5), 1536-1545, doi: 10.1021/

ac702084k, 2008.

Miller, B.R, M. Rigby, L.J.M. Kuijpers, P.B. Krummel, 

L.P. Steele, M. Leist, P.J. Fraser, A. McCulloch, 

C. Harth, P. Salameh, J. Mühle, R.F. Weiss, R.G. 

Prinn, R.H.J. Wang, S. O’Doherty, B.R. Greally, 

and P.G. Simmonds, HFC-23 (CHF3) emission 

trend response to HCFC-22 (CHClF2) production 

and recent HFC-23 emission abatement measures, 

Atmos. Chem. Phys., 10 (16), 7875-7890, doi: 

10.5194/acp-10-7875-2010, 2010.

Millet, D.B., E.L. Atlas, D.R. Blake, N.J. Blake, G.S. 

Diskin, J.S. Holloway, R.C. Hudman, S. Meinardi, 

T.B. Ryerson, and G.W. Sachse, Halocarbon emissions from the United States and Mexico and their 

global warming potential, Environ. Sci. Technol.,

43 (4), 1055-1060, doi: 10.1021/es802146j, 2009.

Møgelberg, T.E., J. Platz, O.J. Nielsen, J. Sehested, and T.J. 

Wallington, Atmospheric chemistry of HFC-236fa: 

Spectrokinetic investigation of the CF3CHO2·CF3

radical, its reactions with NO, and the fate of the 

CF3CHO·CF3 radical, J. Phys. Chem., 99 (15), 

5373-5378, doi: 10.1021/j100015a021, 1995.

Molina, L.T., P.J. Wooldridge, and M.J. Molina, Atmospheric reactions and ultraviolet and infrared absorptivities of nitrogen trifluoride, Geophys. Res. Lett., 

22 (14), 1873-1876, doi: 10.1029/95GL01669, 1995.

Montzka, S.A., and P.J. Fraser (Lead Authors), J.H. Butler, P.S. Connell, D.M. Cunnold, J.S. Daniel, R.G. 

Derwent, S. Lal, A. McCulloch, D.E. Oram, C.E. 

Reeves, E. Sanhueza, L.P. Steele, G.J.M. Velders, 

R.F. Weiss, and R.J. Zander, Controlled substances 

and other source gases, Chapter 1 in Scientific Assessment of Ozone Depletion: 2002, Global Ozone 

Research and Monitoring Project—Report No. 47, 

Geneva, Switzerland, 2003.

Montzka, S.A., R.C. Myers, J.H. Butler, J.W. Elkins, 

L.T. Lock, A.D. Clarke, and A.H. Goldstein, 

Observations of HFC-134a in the remote troposphere, Geophys. Res. Lett., 23 (2), 169-172, doi: 

10.1029/95GL03590, 1996.

Montzka, S.A., J.H. Butler, J.W. Elkins, T.M. Thompson, A.D. Clarke, and L.T. Lock, Present and future trends in the atmospheric burden of ozonedepleting halogens, Nature, 398, 690-694, doi: 

10.1038/19499, 1999.

Montzka, S.A., C.M. Spivakovsky, J.H. Butler, J.W. Elkins, L.T. Lock, and D.J. Mondeel, New observational constraints for atmospheric hydroxyl on 

global and hemispheric scales, Science, 288 (5465), 

500-503, doi: 10.1126/science.288.5465.500, 2000.

Montzka, S.A., J.H. Butler, B.D. Hall, D.J. Mondeel, and 

J.W. Elkins, A decline in tropospheric organic 

bromine, Geophys. Res. Lett., 30 (15), 1826, doi: 

10.1029/2003GL017745, 2003.

Montzka, S.A., M. Aydin, M. Battle, J.H. Butler, E.S. 

Saltzman, B.D. Hall, A.D. Clarke, D. Mondeel, and 

J.W. Elkins, A 350-year atmospheric history for 

carbonyl sulfide inferred from Antarctic firn air and 

air trapped in ice, J. Geophys. Res., 109, D22302, 

doi: 10.1029/2004JD004686, 2004.

Montzka, S.A., P. Calvert, B.D. Hall, J.W. Elkins, T.J. 

Conway, P.P. Tans, and C. Sweeney, On the 

global distribution, seasonality, and budget of atmospheric carbonyl sulfide (COS) and some similarities to CO2, J. Geophys. Res., 112, D09302, doi: 

10.1029/2006JD007665, 2007.

Montzka, S.A., J.S. Daniel, J. Cohen, and K. Vick, Current 

trends, mixing ratios, and emissions of ozone-depleting substances and their substitutes, in Trends 

in Emissions of Ozone-Depleting Substances,

Ozone Layer Recovery, and Implications for Ultraviolet Radiation Exposure, Synthesis and Assessment Product 2.4, report by the U.S. Climate 

Change Science Program and the Subcommittee on 

Global Change Research, edited by A.R. Ravishankara, M.J. Kurylo, and C.A. Ennis, 29-78, Department of Commerce, NOAA National Climatic Data 

Center, Asheville, NC, 2008.

Montzka, S.A., B.D. Hall, and J.W. Elkins, Accelerated increases observed for hydrochlorofluorocarbons since 

2004 in the global atmosphere, Geophys. Res. Lett.,

36, L03804, doi: 10.1029/2008GL036475, 2009.

Montzka, S.A., L. Kuijpers, M.O. Battle, M. Aydin, 

K.R. Verhulst, E.S. Saltzman, and D.W. Fahey, 

Recent increases in global HFC-23 emissions, 

Geophys. Res. Lett., 37, L02808, doi: 10.1029/ 
1.98

Chapter 1

2009GL041195, 2010.

Moore, D.P., and J.J. Remedios, Growth rates of stratospheric HCFC-22, Atmos. Chem. Phys., 8 (1), 73-

82, doi: 10.5194/acp-8-73-2008, 2008.

Moore, R.M., A photochemical source of methyl chloride 

in saline waters, Environ. Sci. Technol., 42 (6), 

1933-1937, doi: 10.1021/es0719201, 2008.

Moore, R.M., A. Gut, and M.O. Andreae, A pilot study 

of methyl chloride emissions from tropical woodrot fungi, Chemosphere, 58 (2), 221-225, doi: 

10.1016/j.chemosphere.2004.03.011, 2005.

Mühle, J., J. Huang, R.F. Weiss, R.G. Prinn, B.R. Miller, P.K. Salameh, C.M. Harth, P.J. Fraser, L.W. 

Porter, B.R. Greally, S. O’Doherty, and P.G. 

Simmonds, Sulfuryl fluoride in the global atmosphere, J. Geophys. Res., 114, D05306, doi: 

10.1029/2008JD011162, 2009.

Mühle, J., A.L. Ganesan, B.R. Miller, P.K. Salameh, 

C.M. Harth, B.R. Greally, M. Rigby, L.W. Porter, L.P. Steele, C.M. Trudinger, P.B. Krummel, 

S. O’Doherty, P.J. Fraser, P.G. Simmonds, R.G. 

Prinn, and R.F. Weiss, Perfluorocarbons in the 

global atmosphere: Tetrafluoromethane, hexafluoroethane, and octafluoropropane, Atmos. Chem. 

Phys., 10 (11), 5145-5164, doi: 10.5194/acp-10-

5145-2010, 2010.

Murphy, D.M., D.J. Cziczo, P.K. Hudson, and D.S. 

Thomson, Carbonaceous material in aerosol particles in the lower stratosphere and tropopause 

region, J. Geophys. Res., 112, D04203, doi: 

10.1029/2006JD007297, 2007.

Naik, V., A.K. Jain, K.O. Patten, and D.J. Wuebbles, Consistent sets of atmospheric lifetimes and radiative 

forcings on climate for CFC replacements: HCFCs 

and HFCs, J. Geophys. Res., 105 (D5), 6903-6914, 

doi: 10.1029/1999JD901128, 2000.

Neuman, J.A., J.B. Nowak, L.G. Huey, J.B. Burkholder, 

J.E. Dibb, J.S. Holloway, J. Liao, J. Peischl, J.M. 

Roberts, T.B. Ryerson, E. Scheuer, H. Stark, R.E. 

Stickel, D.J. Tanner, and A. Weinheimer, Bromine 

measurements in ozone depleted air over the Arctic 

Ocean, Atmos. Chem. Phys., 10 (14), 6503-6514, 

doi: 10.5194/acp-10-6503-2010, 2010.

Nevison, C.D., N.M. Mahowald, R.F. Weiss, and R.G. 

Prinn, Interannual and seasonal variability in atmospheric N2O, Global Biogeochem. Cycles, 21, 

GB3017, doi: 10.1029/2006GB002755, 2007.

Newman, P.A. and J.A. Pyle (Lead Authors), J. Austin, 

G.O. Braathen, P.O. Canziani, K.S. Carslaw, P.M. 

de F. Forster, S. Godin-Beekmann, B.M. Knudsen, 

K. Kreher, H. Nakane, S. Pawson, V. Ramaswamy, 

M. Rex, R.J. Salawitch, D.T. Shindell, A. Tabazadeh, and D.W. Toohey, Polar stratospheric ozone: 

Past and future, Chapter 3 in Scientific Assessment 

of Ozone Depletion: 2002, Global Ozone Research 

and Monitoring Project—Report No. 47, Geneva, 

Switzerland, 2003.

Newman, P.A., J.S. Daniel, D.W. Waugh, and E.R. Nash, A 

new formulation of equivalent effective stratospheric 

chlorine (EESC), Atmos. Chem. Phys., 7 (17), 4537-

4522, doi: 10.5194/acp-7-4537-2007, 2007.

Nielsen, O.J., B.F. Scott, C. Spencer, T.J. Wallington, and 

J.C. Ball, Trifluoroacetic acid in ancient freshwater, 

Atmos. Environ., 35 (16), 2799-2801, doi: 10.1016/

S1352-2310(01)00148-0, 2001.

Nielsen, O.J., M.S. Javadi, M.P. Sulbaek Andersen, M.D. 

Hurley, T.J. Wallington, and R. Singh, Atmospheric chemistry of CF3CF=CH2: Kinetics and mechanisms of gas-phase reactions with Cl atoms, OH 

radicals, and O3, Chem. Phys. Lett., 439 (1-3),18-

22, doi: 10.1016/j.cplett.2007.03.053, 2007.

Notholt, J., B.P. Luo, S. Fueglistaler, D. Weisenstein, 

M. Rex, M.G. Lawrence, H. Bingemer, I. Wohltmann, T. Corti, T. Warneke, R. von Kuhlmann, 

and T. Peter, Influence of tropospheric SO2 emissions on particle formation and the stratospheric 

humidity, Geophys. Res. Lett., 32, L07810, doi: 

10.1029/2004GL022159, 2005.

O’Brien, L.M., N.R.P. Harris, A.D. Robinson, B. Gostlow, N. Warwick, X. Yang, and J.A. Pyle, Bromocarbons in the tropical marine boundary layer at the 

Cape Verde Observatory – measurements and modelling, Atmos. Chem. Phys., 9 (22), 9083-9099, doi: 

10.5194/acp-9-9083-2009, 2009.

O’Doherty, S., D.M. Cunnold, A. Manning, B.R. Miller, 

R.H.J. Wang, P.B. Krummel, P.J. Fraser, P.G. Simmonds, A. McCulloch, R.F. Weiss, P. Salameh, 

L.W. Porter, R.G. Prinn, J. Huang, G. Sturrock, 

D. Ryall, R.G. Derwent, and S.A. Montzka, Rapid 

growth of hydrofluorocarbon 134a and hydrochlorofluorocarbons 141b, 142b, and 22 from Advanced 

Global Atmospheric Gases Experiment (AGAGE) 

observations at Cape Grim, Tasmania, and Mace 

Head, Ireland, J. Geophys. Res., 109, D06310, doi: 

10.1029/2003JD004277, 2004.

O’Doherty, S., D.M. Cunnold, B.R. Miller, J. Mühle, A. 

McCulloch, P.G. Simmonds, A.J. Manning, S. 

Reimann, M.K. Vollmer, B.R. Greally, R.G. Prinn, 

P.J. Fraser, L.P. Steele, P.B. Krummel, B.L. Dunse, 

L.W. Porter, C.R. Lunder, N. Schmidbauer, O. Hermansen, P.K. Salameh, C.M. Harth, R.H.J. Wang, 

and R.F. Weiss, Global and regional emissions of 

HFC-125 (CHF2CF3) from in situ and air archive 

atmospheric observations at AGAGE and SOGE 

observatories, J. Geophys. Res., 114, D23304, doi: 

10.1029/2009JD012184, 2009.

Oram, D.E., Trends of Long-Lived Anthropogenic Halocarbons in the Southern Hemisphere and Model 
1.99

ODSs and Related Chemicals

Calculations of Global Emissions, Ph.D. thesis, 

Univ. of E. Anglia, Norwich, U.K., 249 pp., 1999.

Oram, D.E., W.T. Sturges, S.A. Penkett, A. McCulloch, and 

P.J. Fraser, Growth of fluoroform (CHF3, HFC-23) 

in the background atmosphere, Geophys. Res. Lett.,

25 (1), 35-38, doi: 10.1029/97GL03483, 1998.

Orkin, V.L., F. Louis, R.E. Huie, and M.J. Kurylo, Photochemistry of bromine-containing fluorinated 

alkenes: Reactivity toward OH and UV spectra, 

J. Phys. Chem. A, 106 (43), 10195-10199, doi: 

10.1021/jp014436s, 2002.

Orkin, V.L., L.E. Martynova, and A.N. Ilichev, High-accuracy measurements of OH reaction rate constants 

and IR absorption spectra: CH2=CF-CF3 and transCHF=CH-CF3, J. Phys. Chem. A., 114 (19), 5967-

5979, doi: 10.1021/jp9092817, 2010.

Oyaro, N., S.R. Sellevåg, and C.J. Nielsen, Study of the 

OH and Cl-initiated oxidation, IR absorption crosssection, radiative forcing, and global warming 

potential of four C4-hydrofluoroethers, Environ. 

Sci. Technol., 38 (21), 5567-5576, doi: 10.1021/

es0497330, 2004.

Oyaro, N., S.R. Sellevåg, and C.J. Nielsen, Atmospheric 

chemistry of hydrofluoroethers: Reaction of a series of hydrofluoroethers with OH radicals and Cl 

atoms, atmospheric lifetimes, and global warming 

potentials, J. Phys. Chem. A, 109 (2), 337-346, doi: 

10.1021/jp047860c, 2005.

Packer, M., Algal capture of carbon dioxide; biomass 

generation as a tool for greenhouse gas mitigation 

with reference to New Zealand energy strategy 

and policy, Energy Policy, 37 (9), 3428-3437, doi: 

10.1016/j.enpol.2008.12.025, 2009.

Palmer, C.J., and C.J. Reason, Relationships of surface bromoform concentrations with mixed layer 

depth and salinity in the tropical oceans, Global 

Biogeochem. Cycles, 23, GB2014, doi: 10.1029/ 

2008GB003338, 2009.

Papadimitriou, V.C., R.K. Talukdar, R.W. Portmann, A.R. 

Ravishankara, and J.B. Burkholder, CF3CF=CH2

and (Z)-CF3CF=CHF: Temperature dependent 

OH rate coefficients and global warming potentials, Phys. Chem. Chem. Phys., 10, 808-820, doi: 

10.1039/B714382F, 2008a.

Papadimitriou, V.C., R.W. Portmann, D.W. Fahey, J. 

Mühle, R.F. Weiss, and J.B. Burkholder, Experimental and theoretical study of the atmospheric 

chemistry and global warming potential of SO2F2, 

J. Phys. Chem. A, 112 (49), 12657-12666, doi: 

10.1021/jp806368u, 2008b.

Park, M., W.J. Randel, L.K. Emmons, P.F. Bernath, K.A. 

Walker, and C.D. Boone, Chemical isolation in 

the Asian monsoon anticyclone observed in Atmospheric Chemistry Experiment (ACE-FTS) data, 

Atmos. Chem. Phys., 8 (3), 757-764, doi: 10.5194/

acp-8-757-2008, 2008.

Peters, W., A.R. Jacobson, C. Sweeney, A.E. Andrews, 

T.J. Conway, K. Masarie, J.B. Miller, L.M.P. Bruhwiler, G. Pétron, A.I. Hirsch, D.E.J. Worthy, G.R. 

van der Werf, J.T. Randerson, P.O. Wennberg, M.C. 

Krol, and P.P. Tans, An atmospheric perspective on 

North American carbon dioxide exchange: CarbonTracker, Proc. Natl. Acad. Sci., 104 (48), 18925-

18930, doi: 10.1073/pnas.0708986104, 2007.

Peylin, P., D. Baker, J. Sarmiento, P. Ciais, and P. Bousquet, Influence of transport uncertainty on annual mean and seasonal inversions of atmospheric 

CO2 data, J. Geophys. Res., 107 (D19), 4385, doi: 

10.1029/2001JD000857, 2002.

Ploeger, F., P. Konopka, G. Günther, J.-U. Grooß, and 

R. Müller, Impact of the vertical velocity scheme 

on modeling transport in the tropical tropopause layer, J. Geophys. Res., 115, D03301, doi: 

10.1029/2009JD012023, 2010.

Porter, E., J. Wenger, J. Treacy, H. Sidebottom, A. Mellouki, S. Téton, and G. LeBras, Kinetic studies on 

the reactions of hydroxyl radicals with diethers and 

hydroxyethers, J. Phys. Chem. A, 101 (32), 5770-

5775, doi: 10.1021/jp971254i, 1997.

Portmann, R.W., and S. Solomon, Indirect radiative forcing of the ozone layer during the 21st 

century, Geophys. Res. Lett., 34, L02813, doi: 

10.1029/2006GL028252, 2007.

Prather, M.J., Timescales in atmospheric chemistry: 

CH3Br, the ocean, and ozone depletion potentials, 

Global Biogeochem. Cycles, 11 (3), 393-400, doi: 

10.1029/97GB01055, 1997.

Prather, M., and D. Ehhalt (Co-ordinating Lead Authors), 

F. Dentener, R. Derwent, E. Dlugokencky, E. Holland, I. Isaksen, J. Katima, V. Kirchhoff, P. Matson, 

P. Midgley, and M. Wang (Lead Authors), Atmospheric chemistry and greenhouse gases, Chapter 

4 in Climate Change 2001: The Scientific Basis.

Contribution of Working Group I to the Third Assessment Report of the Intergovernmental Panel on 

Climate Change, edited by J.T. Houghton, Y. Ding, 

D.J. Griggs, M. Noguer, P.J. van der Linden, X. 

Dai, K. Maskell, and C.A. Johnson, 881 pp., Cambridge University Press, Cambridge, U.K., 2001.

Prather, M.J., and J. Hsu, NF3, the greenhouse gas missing 

from Kyoto, Geophys. Res. Lett., 35, L12810, doi: 

10.1029/2008GL034542, 2008.

Prather, M.J., and J. Hsu, Correction to “NF3, the greenhouse gas missing from Kyoto,” Geophys. Res. Lett., 

37, L11807, doi: 10.1029/2010GL043831, 2010.

Prinn, R.G., and R. Zander (Lead Authors), D.M. Cunnold, J.W. Elkins, A. Engel, P.J. Fraser, M.R. Gunson, M.K.W. Ko, E. Mahieu, P.M. Midgley, J.M. 
1.100

Chapter 1

Russell III, C.M. Volk, and R.F. Weiss, Long-lived 

ozone-related compounds, Chapter 1 in Scientific 

Assessment of Ozone Depletion: 1998, Global 

Ozone Research and Monitoring Project–Report 

No. 44, World Meteorological Organization, Geneva, Switzerland, 1999.

Prinn, R., D. Cunnold, R. Rasmussen, P. Simmonds, F. 

Alyea, A. Crawford, P. Fraser, and R. Rosen, Atmospheric emissions and trends of nitrous oxide 

deduced from 10 years of ALE-GAGE data, J. Geophys. Res., 95 (D11), 18369-18385, doi: 10.1029/

JD095iD11p18369, 1990.

Prinn, R.G., R.F. Weiss, P.J. Fraser, P.G. Simmonds, D.M. 

Cunnold, F.N. Alyea, S. O’Doherty, P. Salameh, 

B.R. Miller, J. Huang, R.H.J. Wang, D.E. Hartley, 

C. Harth, L.P. Steele, G. Sturrock, P.M. Midgley, 

and A. McCulloch, A history of chemically and radiatively important gases in air deduced from ALE/

GAGE/AGAGE, J. Geophys. Res., 105 (D14), 

17751-17792, doi: 10.1029/2000JD900141, 2000.

Prinn, R.G., J. Huang, R.F. Weiss, D.M. Cunnold, P.J. Fraser, P.G. Simmonds, A. McCulloch, C. Harth, S. Reimann, P. Salameh, S. O’Doherty, R.H.J. Wang, L.W. 

Porter, B.R. Miller, and P.B. Krummel, Evidence for 

variability of atmospheric hydroxyl radicals over 

the past quarter century, Geophys. Res. Lett., 32, 

L07809, doi: 10.1029/2004GL022228, 2005.

Pyle, J.A., N. Warwick, X. Yang, P.J. Young, and G. 

Zeng, Climate/chemistry feedbacks and biogenic 

emissions, Phil. Trans. R. Soc. A, 365 (1856), 1727-

1740, doi: 10.1098/rsta.2007.2041, 2007.

Quack, B., E. Atlas, G. Petrick, and D.W.R. Wallace, 

Bromoform and dibromomethane above the Mauritanian upwelling: Atmospheric distributions and 

oceanic emissions, J. Geophys. Res., 112, D09312, 

doi: 10.1029/2006JD007614, 2007.

Ramaswamy, V. (Co-ordinating Lead Author), O. 

Boucher, J. Haigh, D. Hauglustaine, J. Haywood, 

G. Myhre, T. Nakajima, G.Y. Shi, and S. Solomon (Lead Authors), Radiative forcing of climate 

change, Chapter 6 in Climate Change 2001: The 

Scientific Basis. Contribution of Working Group 

I to the Third Assessment Report of the Intergovernmental Panel on Climate Change, edited by J.T. 

Houghton, Y. Ding, D.J. Griggs, M. Noguer, P.J. 

van der Linden, X. Dai, K. Maskell, and C.A. Johnson, 881 pp., Cambridge University Press, Cambridge, U.K., 2001. 

Randel, W.J., M. Park, L. Emmons, D. Kinnison, P. Bernath, K.A. Walker, C. Boone, and H. Pumphrey, 

Asian monsoon transport of pollution to the stratosphere, Science, 328 (5978), 611-613, doi: 10.1126/

science.1182274, 2010.

Ravishankara, A.R., S. Solomon, A.A. Turnipseed, and 

R.F. Warren, Atmospheric lifetimes of long-lived 

halogenated species, Science, 259 (5092), 194-199, 

doi: 10.1126/science.259.5092.194, 1993.

Ravishankara, A.R., J.S. Daniel, and R.W. Portmann, Nitrous 

oxide (N2O): The dominant ozone-depleting substance 

emitted in the 21st century, Science, 326 (5949), 123-

125, doi: 10.1126/science.1176985, 2009.

Read, K.A., A.S. Mahajan, L.J. Carpenter, M.J. Evans, 

B.V.E. Faria, D.E. Heard, J.R. Hopkins, J.D. Lee, 

S.J. Moller, A.C. Lewis, L. Mendes, J.B. McQuaid, 

H. Oetjen, A. Saiz-Lopez, M.J. Pilling, and J.M.C. 

Plane, Extensive halogen-mediated ozone destruction over the tropical Atlantic Ocean, Nature, 453, 

1232-1235, doi: 10.1038/nature07035, 2008.

Read, W.G., L. Froidevaux, and J.W. Waters, Microwave limb sounder measurement of stratospheric 

SO2 from the Mt. Pinatubo Volcano, Geophys. 

Res. Lett., 20 (12), 1299-1302, doi: 10.1029/ 

93GL00831, 1993.

Redeker, K.R., and R.J. Cicerone, Environmental controls over methyl halide emissions from rice paddies, Global Biogeochem. Cycles, 18, GB1027, doi: 

10.1029/2003GB002092, 2004.

Reeves, C.E., W.T. Sturges, G.A. Sturrock, K. Preston, 

D.E. Oram, J. Schwander, R. Mulvaney, J.-M. Barnola, and J. Chappellaz, Trends of halon gases in 

polar firn air: Implications for their emission distributions, Atmos. Chem. Phys., 5 (8), 2055-2064, doi: 

10.5194/acp-5-2055-2005, 2005.

Reichenau, T.G., and G. Esser, Is interannual fluctuation of atmospheric CO2 dominated by combined 

effects of ENSO and volcanic aerosols?, Global 

Biogeochem. Cycles, 17 (4), 1094, doi: 10.1029/ 

2002GB002025, 2003.

Reimann, S., D. Schaub, K. Stemmler, D. Folini, M. 

Hill, P. Hofer, B. Buchmann, P.G. Simmonds, 

B.R. Greally, and S. O’Doherty, Halogenated 

greenhouse gases at the Swiss high alpine site of 

Jungfraujoch (3580 m asl): Continuous measurements and their use for regional European source 

allocation, J. Geophys. Res., 109, D05307, doi: 

10.1029/2003JD003923, 2004.

Reimann, S., M.K. Vollmer, D. Folini, M. Steinbacher, 

M. Hill, B. Buchmann, R. Zander, and E. Mahieu, 

Observations of long-lived anthropogenic halocarbons at the high-Alpine site of Jungfraujoch (Switzerland) for assessment of trends and European 

sources, Sci. Total Environ., 391 (2-3), 224-231, 

doi: 10.1016/j.scitotenv.2007.10.022, 2008.

Rhew, R.C., and T. Abel, Measuring simultaneous production and consumption fluxes of methyl chloride 

and methyl bromide in annual temperate grasslands, Environ. Sci. Technol., 41 (22), 7837-7843, 

doi: 10.1021/es0711011, 2007.
1.101

ODSs and Related Chemicals

Rhew, R.C., B.R. Miller, and R.F. Weiss, Natural methyl bromide and methyl chloride emissions from 

coastal salt marshes, Nature, 403, 292-295, doi: 

10.1038/35002043, 2000.

Rhew, R.C., B.R. Miller, M.K. Vollmer, and R.F. Weiss, 

Shrubland fluxes of methyl bromide and methyl 

chloride, J. Geophys. Res., 106 (D18), 20875-

20882, doi: 10.1029/2001JD000413, 2001.

Rhew, R.C., L. �Ø stergaard, E.S. Saltzman, and M.F. Yanofsky, Genetic control of methyl halide production in Arabidopsis, Current Biology, 13 (20), 

1809-1813, doi: 10.1016/j.cub.2003.09.055, 2003.

Rhew, R.C., Y.A. Teh, and T. Abel, Methyl halide and 

methane fluxes in the northern Alaskan coastal tundra, J. Geophys. Res., 112, G02009, doi: 

10.1029/2006JG000314, 2007.

Rhew, R.C., B.R. Miller, and R.F. Weiss, Chloroform, 

carbon tetrachloride and methyl chloroform 

fluxes in southern California ecosystems, Atmos. 

Environ., 42 (30), 7135-7140, doi: 10.1016/j.atmosenv.2008.05.038, 2008.

Rhew, R.C., C. Chen, Y.A. Teh, and D. Baldocchi, Gross 

fluxes of methyl chloride and methyl bromide 

in a California oak-savanna woodland, Atmos. 

Environ., 44 (16), 2054-2061, doi: 10.1016/j.atmosenv.2009.12.014, 2010.

Ricaud, P., B. Barret, J.-L. Attié, E. Motte, E. Le Flochmoën, H. Teyssèdre, V.-H. Peuch, N. Livesey, A. 

Lambert, and J.-P. Pommereau, Impact of land convection on troposphere-stratosphere exchange in 

the tropics, Atmos. Chem. Phys., 7 (21), 5639-5657, 

doi: 10.5194/acp-7-5639-2007, 2007.

Rice, A.L., C.L. Butenhoff, M.J. Shearer, D. Teama, 

T.N. Rosenstiel, and M.A.K. Khalil, Emissions 

of anaerobically produced methane by trees, 

Geophys. Res. Lett., 37, L03807, doi: 10.1029/ 

2009GL041565, 2010.

Richter, A., F. Wittrock, A. Ladstätter-Weißßßßßßß enmayer, and J.P. Burrows, GOME measurements of 

stratospheric and tropospheric BrO, Adv. Space 

Res., 29 (11), 1667-1672, doi: 10.1016/S0273-

1177(02)00123-0, 2002.

Rigby, M., R.G. Prinn, P.J. Fraser, P.G. Simmonds, 

R.L. Langenfelds, J. Huang, D.M. Cunnold, L.P. 

Steele, P.B. Krummel, R.F. Weiss, S. O’Doherty, 

P.K. Salameh, H.J. Wang, C.M. Harth, J. Mühle, 

and L.W. Porter, Renewed growth of atmospheric 

methane, Geophys. Res. Lett., 35, L22805, doi: 

10.1029/2008GL036037, 2008.

Rinsland, C.P., E. Mahieu, R. Zander, N.B. Jones, M.P. 

Chipperfield, A. Goldman, J. Anderson, J.M. Russell III, P. Demoulin, J. Notholt, G.C. Toon, J.-F.

Blavier, B. Sen, R. Sussmann, S.W. Wood, A. 

Meier, D.W.T. Griffith, L.S. Chiou, F.J. Murcray, 

T.M. Stephen, F. Hase, S. Mikuteit, A. Schulz, and 

T. Blumenstock, Long-term trends of inorganic 

chlorine from ground-based infrared solar spectra: Past increases and evidence for stabilization, 

J. Geophys. Res., 108 (D8), 4252, doi: 10.1029/ 

2002JD003001, 2003.

Rinsland, C.P., C. Boone, R. Nassar, K. Walker, P. Bernath, E. Mahieu, R. Zander, J.C. McConnell, and 

L. Chiou, Trends of HF, HCl, CCl2F2, CCl3F, 

CHClF2 (HCFC-22), and SF6 in the lower stratosphere from Atmospheric Chemistry Experiment 

(ACE) and Atmospheric Trace Molecule Spectroscopy (ATMOS) measurement near 30°N latitude, Geophys. Res. Lett., 32 (16), L16S03, doi: 

10.1029/2005GL022415, 2005.

Rinsland, C.P., R. Nassar, C.D. Boone, P. Bernath, L. 

Chiou, D.K. Weisenstein, E. Mahieu, and R. Zander, Spectroscopic detection of COClF in the tropical and mid-latitude lower stratosphere, J. Quant. 

Spectrosc. Radiat. Transfer, 105 (3), 467-475, doi: 

10.1016/j.jqsrt.2006.11.013, 2007.

Rinsland, C.P., L. Chiou, C. Boone, P. Bernath, and E. 

Mahieu, First measurements of the HCFC-142b 

trend from atmospheric chemistry experiment 

(ACE) solar occultation spectra, J. Quant. Spectrosc. Radiat. Transfer, 110 (18), 2127-2134, doi: 

10.1016/j.jqsrt.2009.05.011, 2009.

Robson, J.I., L.K. Gohar, M.D. Hurley, K.P. Shine, and 

T.J. Wallington, Revised IR spectrum, radiative efficiency and global warming potential of nitrogen 

trifluoride, Geophys. Res. Lett., 33, L10817, doi: 

10.1029/2006GL026210, 2006.

Rödenbeck, C., S. Houweling, M. Gloor, and M. Heimann, CO2 flux history 1982–2001 inferred from 

atmospheric data using a global inversion of atmospheric transport, Atmos. Chem. Phys., 3 (6), 1919-

1964, doi: 10.5194/acp-3-1919-2003, 2003.

Römpp, A., O. Klemm, W. Fricke, and H. Frank, Haloacetates in fog and rain, Environ. Sci. Technol., 35

(7), 1294-1298, doi: 10.1021/es0012220, 2001.

Rontu Carlon, N., D.K. Papanastasiou, E.L. Fleming, C.H. 

Jackman, P.A. Newman, and J.B. Burkholder, UV 

absorption cross sections of nitrous oxide (N2O) 

and carbon tetrachloride (CCl4) between 210 and 

350 K and the atmospheric implications, Atmos. 

Chem. Phys., 10 (13), 6137-6149, doi: 10.5194/

acp-10-6137-2010, 2010.

Ross, M.N., M.Y. Danilin, D.K. Weisenstein, and 

M.K.W. Ko, Ozone depletion caused by NO and 

H2O emissions from hydrazine-fueled rockets, 

J. Geophys. Res., 109, D21305, doi: 10.1029/ 

2003JD004370, 2004.

Ross, M., D. Toohey, M. Peinemann, and P. Ross, Limits 

on the space launch market related to stratospheric 
1.102

Chapter 1

ozone depletion, Astropolitics, 7 (1), 50-82, doi: 

10.1080/14777620902768867, 2009.

Rowland, F.S., S.C. Tyler, D.C. Montague, and Y. Makide, 

Dichlorodifluoromethane, CCl2F2, in the Earth’s atmosphere, Geophys. Res. Lett., 9 (4), 481-484, doi: 

10.1029/GL009i004p00481, 1982.

Rudolph, J., A. Khedim, R. Koppmann, and B. Bonsang, 

Field study of the emissions of methyl chloride and 

other halocarbons from biomass burning in Western Africa, J. Atmos. Chem., 22 (1-2), 67-80, doi: 

10.1007/BF00708182, 1995.

Rudolph, J., K. von Czapiewski, and R. Koppmann, 

Emissions of methyl chloroform (CH3CCl3) from 

biomass burning and the tropospheric methyl chloroform budget, Geophys. Res. Lett., 27 (13), 1887-

1890, doi: 10.1029/1999GL011178, 2000.

Saito, T., and Y. Yokouchi, Diurnal variation in methyl 

halide emission rates from tropical ferns, Atmos. 

Environ., 40 (16), 2806-2811, doi: 10.1016/j.atmosenv.2006.01.016, 2006.

Saito, T., and Y. Yokouchi, Stable carbon isotope ratio 

of methyl chloride emitted from glasshouse-grown 

tropical plants and its implication for the global 

methyl chloride budget, Geophys. Res. Lett., 35, 

L08807, doi: 10.1029/2007GL032736, 2008.

Saito, T., Y. Yokouchi, Y. Kosugi, M. Tani, E. Philip, and T. Okuda, Methyl chloride and isoprene 

emissions from tropical rain forest in Southeast Asia, Geophys. Res. Lett., 35, L19812, doi: 

10.1029/2008GL035241, 2008.

Saiz-Lopez, A., J.M.C. Plane, and J.A. Shillito, Bromine oxide in the mid-latitude marine boundary layer, Geophys. Res. Lett., 31, L03111, doi: 

10.1029/2003GL018956, 2004.

Salawitch, R.J., D.K. Weisenstein, L.J. Kovalenko, C.E. 

Sioris, P.O. Wennberg, K. Chance, M.K.W. Ko, 

and C.A. McLinden, Sensitivity of ozone to bromine in the lower stratosphere, Geophys. Res. Lett.,

32, L05811, doi: 10.1029/2004GL021504, 2005.

Salawitch, R.J., T. Canty, T. Kurosu, K. Chance, Q. Liang, 

A. da Silva, S. Pawson, J.E. Nielsen, J.M. Rodriguez, P.K. Bhartia, X. Liu, L.G. Huey, J. Liao, R.E. 

Stickel, D.J. Tanner, J.E. Dibb, W.R. Simpson, D. 

Donohoue, A. Weinheimer, F. Flocke, D. Knapp, D. 

Montzka, J.A. Neuman, J.B. Nowak, T.B. Ryerson, 

S. Oltmans, D.R. Blake, E.L. Atlas, D.E. Kinnison, 

S. Tilmes, L.L. Pan, F. Hendrick, M. Van Roozendael, K. Kreher, P.V. Johnston, R.S. Gao, B. Johnson, T.P. Bui, G. Chen, R.B. Pierce, J.H. Crawford, 

and D.J. Jacob, A new interpretation of total column BrO during Arctic spring, Geophys. Res. Lett., 

37, L21805, doi:10.1029/2010GL043798, 2010.

Saltzman, E.S., M. Aydin, W.J. De Bruyn, D.B. King, 

and S.A. Yvon-Lewis, Methyl bromide in preindustrial air: Measurements from an Antarctic ice core, J. Geophys. Res., 109, D05301, doi: 

10.1029/2003JD004157, 2004.

Saltzman, E.S., M. Aydin, C. Tatum, and M.B. Williams, 

2,000-year record of atmospheric methyl bromide 

from a South Pole ice core, J. Geophys. Res., 113, 

D05304, doi: 10.1029/2007JD008919, 2008.

Sander, S.P., R.R. Friedl, D.M. Golden, M.J. Kurylo, 

G.K. Moortgat, H. Keller-Rudek, P.H. Wine, A.R. 

Ravishankara, C.E. Kolb, M.J. Molina, B.J. Finlayson-Pitts, R.E. Huie, and V.L. Orkin, Chemical 

Kinetics and Photochemical Data for Use in Atmospheric Studies, Evaluation Number 15, JPL Publication 06-02, Jet Propulsion Laboratory, Pasadena, 

Calif., 2006. 

Sander, S.P., J. Abbatt, J.R. Barker, J.B. Burkholder, R.R. 

Friedl, D.M. Golden, R.E. Huie, C.E. Kolb, M.J. 

Kurylo, G.K. Moortgat, V.L. Orkin, and P.H. Wine, 

Chemical Kinetics and Photochemical Data for 

Use in Atmospheric Studies, Evaluation Number 

17, JPL Publication 10-6, Jet Propulsion Laboratory, Pasadena, Calif., available: http://jpldataeval.

jpl.nasa.gov/, 2010.

Schauffler, S.M., E.L. Atlas, D.R. Blake, F. Flocke, R.A. 

Lueb, J.M. Lee-Taylor, V. Stroud, and W. Travnicek, 

Distributions of brominated organic compounds in 

the troposphere and lower stratosphere, J. Geophys. 

Res., 104 (D17), 21513-21536, 1999.

Schauffler, S.M., E.L. Atlas, S.G. Donnelly, A. Andrews, 

S.A. Montzka, J.W. Elkins, D.F. Hurst, P.A. Romashkin, G.S. Dutton, and V. Stroud, Chlorine budget and partitioning during the Stratospheric Aerosol and Gas Experiment (SAGE) III Ozone Loss and 

Validation Experiment (SOLVE), J. Geophys. Res.,

108, 4173, doi: 10.1029/2001JD002040, 2003.

Schiller. C., J.-U. Grooβ, P. Konopka, F. Plöger, F.H. 

Silva dos Santos, and, N. Spelten, Hydration and 

dehydration at the trpical tropopause, Atmos. Chem. 

Phys., 9 (24), 9647-9660, doi: 10.5194/acp-9-9647-

2009, 2009.

Schmid, O., J.M. Reeves, J.C. Wilson, C. Wiedinmyer, 

C.A. Brock, D.W. Toohey, L.M. Avallone, A.M. 

Gates, and M.N. Ross, Size-resolved particle 

emission indices in the stratospheric plume of an 

Athena II rocket, J. Geophys. Res., 108, 4250, doi: 

10.1029/2002JD002486, 2003.

Schmittner, A., A. Oschlies, H.D. Matthews, and E.D. Galbraith, Future changes in climate, ocean circulation, 

ecosystems, and biogeochemical cycling simulated 

for a business-as-usual CO2 emission scenario until year 4000 AD, Global Biogeochem. Cycles, 22, 

GB1013, doi: 10.1029/2007GB002953, 2008.

Schofield, R., K. Kreher, B.J. Connor, P.V. Johnston, 

A. Thomas, D. Shooter, M.P. Chipperfield, C.D. 
1.103

ODSs and Related Chemicals

Rodgers, and G.H. Mount, Retrieved tropospheric 

and stratospheric BrO columns over Lauder, New 

Zealand, J. Geophys. Res., 109, D14304, doi: 

10.1029/2003JD004463, 2004.

Schofield, R., P.V. Johnston, A. Thomas, K. Kreher, B.J. 

Connor, S. Wood, D. Shooter, M.P. Chipperfield, 

A. Richter, R. von Glasow, and C.D. Rodgers, Tropospheric and stratospheric BrO columns over Arrival Heights, Antarctica, 2002, J. Geophys. Res., 

111, D22310, doi: 10.1029/2005JD007022, 2006.

Shine, K.P., L.K. Gohar, M.D. Hurley, G. Marston, D. 

Martin, P.G. Simmonds, T.J. Wallington, and M. 

Watkins, Perfluorodecalin: Global warming potential and first detection in the atmosphere, Atmos. Environ., 39 (9), 1759-1763, doi: 10.1016/j.

atmosenv.2005.01.001, 2005.

Shorter, J.H., C.E. Kolb, P.M. Crill, R.A. Kerwin, R.W. 

Talbot, M.E. Hines, and R.C. Harriss, Rapid degradation of atmospheric methyl bromide in soils, Nature, 377, 717-719, doi: 10.1038/377717ao, 1995.

Simmonds, P.G., R.G. Derwent, A.J. Manning, P.J. Fraser, P.B. Krummel, S. O’Doherty, R.G. Prinn, D.M. 

Cunnold, B.R. Miller, H.J. Wang, D.B. Ryall, L.W. 

Porter, R.F. Weiss, and P.K. Salameh, AGAGE Observations of methyl bromide and methyl chloride 

at Mace Head, Ireland, and Cape Grim, Tasmania, 

1998–2001, J. Atmos. Chem., 47 (3), 243-269, doi: 

10.1023/B:JOCH.0000021136.52340.9c, 2004.

Simmonds, P.G., A.J. Manning, D.M. Cunnold, A. McCulloch, S. O’Doherty, R.G. Derwent, P.B. Krummel, P.J. Fraser, B. Dunse, L.W. Porter, R.H.J. 

Wang, B.R. Greally, B.R. Miller, P. Salameh, R.F. 

Weiss, and R.G. Prinn, Global trends, seasonal 

cycles, and European emissions of dichloromethane, trichloroethene, and tetrachloroethene from 

the AGAGE observations at Mace Head, Ireland, 

and Cape Grim, Tasmania, J. Geophys. Res., 111, 

D18304, doi: 10.1029/2006JD007082, 2006.

Simpson, I.J., S. Meinardi, N.J. Blake, F.S. Rowland, and D.R. Blake, Long-term decrease in the 

global atmospheric burden of tetrachloroethene 

(C2Cl4), Geophys. Res. Lett., 31, L08108, doi: 

10.1029/2003GL019351, 2004.

Simpson, I.J., N.J. Blake, D.R. Blake, S. Meinardi, M.P.S. 

Andersen, and F.S. Rowland, Strong evidence for 

negligible methyl chloroform (CH3CCl3) emissions 

from biomass burning, Geophys. Res. Lett., 34, 

L10805, doi: 10.1029/2007GL029383, 2007.

Sinnhuber, B.-M., and I. Folkins, Estimating the contribution of bromoform to stratospheric bromine and its 

relation to dehydration in the tropical tropopause 

layer, Atmos. Chem. Phys., 6 (12), 4755-4761, doi: 

10.5194/acp-6-4755-2006, 2006.

Sinnhuber, B.-M., D.W. Arlander, H. Bovensmann, J.P. 

Burrows, M.P. Chipperfield, C.-F. Enell, U. Frieß, 

F. Hendrick, P.V. Johnston, R.L. Jones, K. Kreher, 

N. Mohamed-Tahrin, R. Müller, K. Pfeilsticker, U. 

Platt, J.-P. Pommereau, I. Pundt, A. Richter, A.M. 

South, K.K. Tørnkvist, M. Van Roozendael, T. 

Wagner, and F. Wittrock, Comparison of measurements and model calculations of stratospheric bromine monoxide, J. Geophys. Res., 107, 4398, doi: 

10.1029/2001JD000940, 2002.

Sinnhuber, B.-M., A. Rozanov, N. Sheode, O.T. Afe, A. 

Richter, M. Sinnhuber, F. Wittrock, J.P. Burrows, 

G.P. Stiller, T. von Clarmann, and A. Linden, 

Global observations of stratospheric bromine monoxide from SCIAMACHY, Geophys. Res. Lett., 32, 

L20810, doi: 10.1029/2005GL023839, 2005.

Sinnhuber, B.-M., N. Sheode, M. Sinnhuber, M.P. Chipperfield, and W. Feng, The contribution of anthropogenic bromine emissions to past stratospheric 

ozone trends: A modelling study, Atmos. Chem. 

Phys., 9 (8), 2863-2871, doi: 10.5194/acp-9-2863-

2009, 2009.

Sioris, C.E., L.J. Kovalenko, C.A. McLinden, R.J. Salawitch, M. Van Roozendael, F. Goutail, M. Dorf, 

K. Pfeilsticker, K. Chance, C. von Savigny, X. 

Liu, T.P. Kurosu, J.-P. Pommereau, H. Bösch, and 

J. Frerick, Latitudinal and vertical distribution of 

bromine monoxide in the lower stratosphere from 

Scanning Imaging Absorption Spectrometer for 

Atmospheric Chartography limb scattering measurements, J. Geophys. Res., 111, D14301, doi: 

10.1029/2005JD006479, 2006.

Siskind, D.E., L. Froidevaux, J.M. Russell, and J. Lean, 

Implications of upper stratospheric trace constituent changes observed by HALOE for O3 and ClO 

from 1992 to 1995, Geophys. Res. Lett., 25 (18), 

3513-3516, 1998.

Sive, B.C., R.K. Varner, H. Mao, D.R. Blake, O.W. Wingenter, and R. Talbot, A large terrestrial source of 

methyl iodide, Geophys. Res. Lett., 34, L17808, 

doi: 10.1029/2007GL030528, 2007.

Solomon, P., J. Barrett, T. Mooney, B. Connor, A. Parrish, and D.E. Siskind, Rise and decline of active 

chlorine in the stratosphere, Geophys. Res. Lett., 

33, L18807, doi: 10.1029/2006GL027029, 2006.

Søndergaard, R., O.J. Nielsen, M.D. Hurley, T.J. Wallington, and R. Singh, Atmospheric chemistry of 

trans-CF3CH=CHF: Kinetics of the gas-phase reactions with Cl atoms, OH radicals, and O3, Chem. 

Phys. Lett., 443 (4-6), 199-204, doi: 10.1016/j.

cplett.2007.06.084, 2007.

Sorokin, V.I., N.P. Gritsan, and A.I. Chichinin, Collisions 

of O(1D) with HF, F2, XeF2, NF3, and CF4: Deactivation and reaction, J. Chem. Phys., 108 (21), 

8995-9003, doi: 10.1063/1.476346, 1998.
1.104

Chapter 1

SPARC (Stratospheric Processes And their Role in Climate), SPARC Assessment of Stratospheric Aerosol 

Properties, edited by L. Thomason and Th. Peter, 

World Climate Research Program Report 124, 

SPARC Report 4, WMO/TD- No. 1295, 346 pp., 

Verrières le Buisson, France, 2006.

SPARC CCMVal (Stratospheric Processes And their Role 

in Climate), SPARC Report on the Evaluation of 

Chemistry-Climate Models, edited by V. Eyring, 

T.G. Shepherd, and D.W. Waugh, SPARC Report 

No. 5, WCRP-132, WMO/TD-No. 1526, 478 pp., 

available: http://www.atmosp.physics.utoronto.ca/

SPARC/ccmval_final/index.php, 2010.

Spivakovsky, C.M., J.A. Logan, S.A. Montzka, Y.J. Balkanski, M. Foreman-Fowler, D.B.A. Jones, L.W. 

Horowitz, A.C. Fusco, C.A.M. Brenninkmeijer, 

M.J. Prather, S.C. Wofsy, and M.B. McElroy, 

Three-dimensional climatological distribution of 

tropospheric OH: Update and evaluation, J. Geophys. Res., 105 (D7), 8931-8980, 2000.

Stemmler, K., D. Folini, S. Ubl, M.K. Vollmer, S. Reimann, S. O’Doherty, B.R. Greally, P.G. Simmonds, 

and A.J. Manning, European emissions of HFC365mfc, a chlorine-free substitute for the foam 

blowing agents HCFC-141b and CFC-11, Environ. 

Sci. Technol., 41 (4), 1145-1151, doi: 10.1021/

es061298h, 2007.

Stephens, B.B., K.R. Gurney, P.P. Tans, C. Sweeney, W. 

Peters, L. Bruhwiler, P. Ciais, M. Ramonet, P. Bousquet, T. Nakazawa, S. Aoki, T. Machida, G. Inoue, 

N. Vinnichenko, J. Lloyd, A. Jordan, M. Heimann, 

O. Shibistova, R.L. Langenfelds, L.P. Steele, R.J. 

Francey, and A.S. Denning, Weak northern and 

strong tropical land carbon uptake from vertical 

profiles of atmospheric CO2, Science, 316, 1732-

1735, doi: 10.1126/science.1137004, 2007.

Stern, D.I., Global sulfur emissions from 1850 to 2000, 

Chemosphere, 58 (2), 163-175, doi: 10.1016/j.chemosphere.2004.08.022, 2005.

Stevens, M.H., R.R. Meier, X. Chu, M.T. DeLand, and 

J.M.C. Plane, Antarctic mesospheric clouds formed 

from space shuttle exhaust, Geophys. Res. Lett., 32, 

L13810, doi: 10.1029/2005GL023054, 2005.

Stiller, G.P., T. von Clarmann, M. Höpfner, N. Glatthor, 

U. Grabowski, S. Kellmann, A. Kleinert, A. Linden, M. Milz, T. Reddmann, T. Steck, H. Fischer, 

B. Funke, M. López-Puertas, and A. Engel, Global 

distribution of mean age of stratospheric air from 

MIPAS SF6 measurements, Atmos. Chem. Phys., 8 

(3), 677-695, doi: 10.5194/acp-8-677-2008, 2008.

Stohl, A., S. Eckhardt, C. Forster, P. James, and N. Spichtinger, On the pathways and timescales of intercontinental air pollution transport, J. Geophys. Res., 

107, 4684, doi: 10.1029/2001JD001396, 2002.

Stohl, A., P. Seibert, J. Arduini, S. Eckhardt, P. Fraser, 

B.R. Greally, C. Lunder, M. Maione, J. Mühle, S. 

O’Doherty, R.G. Prinn, S. Reimann, T. Saito, N. 

Schmidbauer, P.G. Simmonds, M.K. Vollmer, R.F. 

Weiss, and Y. Yokouchi, An analytical inversion 

method for determining regional and global emissions of greenhouse gases: Sensitivity studies and application to halocarbons, Atmos. Chem. Phys., 9 (5), 

1597-1620, doi: 10.5194/acp-9-1597-2009, 2009.

Stohl, A., J. Kim, S. Li, S. O’Doherty, J. Mühle, P.K. Salameh T. Saito, M.K. Vollmer, D. Wan, R.F. Weiss, 

B. Yao, Y. Yokouchi, and L.X. Zhou, Hydrochlorofluorocarbon and hydrofluorocarbon emissions in 

East Asia determined by inverse modeling, Atmos. 

Chem. Phys., 10 (8), 3545-3560, doi: 10.5194/acp10-3545-2010, 2010.

Sturrock, G.A., L.W. Porter, P.J. Fraser, In situ measurement of CFC replacement chemicals and other 

halocarbons at Cape Grim: The AGAGE CG-MS 

program, in Baseline.Atmospheric Program Australia 1997-98, edited by N.W. Tindale, N. Derek, and R.J. Francey, Bureau of Meteorology and 

CSIRO Atmospheric Research, Melbourne, 2001.

Sturrock, G.A., D.M. Etheridge, C.M. Trudinger, P.J. Fraser, and A.M. Smith, Atmospheric histories of halocarbons from analysis of Antarctic firn air: Major 

Montreal Protocol species, J. Geophys. Res., 107, 

4765, doi: 10.1029/2002JD002548, 2002.

Sulbaek Andersen, M.P., A. Toft, O.J. Nielsen, M.D. Hurley, T.J. Wallington, H. Chishima, K. Tonokura, 

S.A. Mabury, J.W. Martin, and D.A. Ellis, Atmospheric chemistry of perfluorinated aldehyde 

hydrates (n-CxF2x+1CH(OH)2, x = 1, 3, 4): Hydration, dehydration, and kinetics and mechanism 

of Cl atom and OH radical initiated oxidation, J. 

Phys. Chem. A, 110 (32), 9854-9860, doi: 10.1021/

jp060404z, 2006.

Sulbaek Andersen, M.P., E.J.K. Nilsson, O.J. Nielsen, 

M.S. Johnson, M.D. Hurley, and T.J. Wallington, 

Atmospheric chemistry of trans-CF3CH=CHCl: 

Kinetics of the gas-phase reactions with Cl atoms, 

OH radicals, and O3, J. Photochem. Photobio. A: 

Chemistry, 199 (1), 92-97, doi: 10.1016/j.jphotochem.2008.05.013, 2008.

Sulbaek Andersen, M.P., M.D. Hurley, and T.J. Wallington, Kinetics of the gas phase reactions of chlorine 

atoms and OH radicals with CF3CBr=CH2 and 

CF3CF2CBr=CH2, Chem. Phys. Lett., 482 (1-3), 20-

23, doi: 10.1016/j.cplett.2009.09.056, 2009a.

Sulbaek Andersen, M.P., D.R. Blake, F.S. Rowland, 

M.D. Hurley, and T.J. Wallington, Atmospheric 

chemistry of sulfuryl fluoride: Reaction with OH 

radicals, Cl atoms and O3, atmospheric lifetime, IR 

spectrum, and global warming potential, Environ. 
1.105

ODSs and Related Chemicals

Sci. Technol., 43 (4), 1067-1070, doi: 10.1021/

es802439f, 2009b.

Suntharalingam, P., A.J. Kettle, S.M. Montzka, and D.J. 

Jacob, Global 3-D model analysis of the seasonal 

cycle of atmospheric carbonyl sulfide: Implications 

for terrestrial vegetation uptake, Geophys. Res. Lett., 

35, L19801, doi: 10.1029/2008GL034332, 2008.

Tanhua, T., and K.A. Olsson, Removal and bioaccumulation of anthropogenic, halogenated transient tracers 

in an anoxic fjord, Mar. Chem., 94 (1-4), 27-41, 

doi: 10.1016/j.marchem.2004.07.009, 2005.

Taniguchi, N., T.J. Wallington, M.D. Hurley, A.G. Guschin, L.T. Molina, and M.J. Molina, Atmospheric 

chemistry of C2F5C(O)CF(CF3)2: Photolysis and 

reaction with Cl atoms, OH radicals, and ozone, J. 

Phys. Chem. A, 107 (15), 2674-2679, doi: 10.1021/

jp0220332, 2003.

Tarnocai, C., J.G. Canadell, E.A.G. Schuur, P. Kuhry, 

G. Mazhitova, and S. Zimov, Soil organic carbon 

pools in the northern circumpolar permafrost region, Global Biogeochem. Cycles, 23, GB2023, 

doi: 10.1029/2008GB003327, 2009.

Teh, Y.A., R.C. Rhew, A. Atwood, and T. Abel, Water, temperature, and vegetation regulation of 

methyl chloride and methyl bromide fluxes 

from a shortgrass steppe ecosystem, Global 

Change Biol., 14 (1), 77-91, doi: 10.1111/j.1365-

2486.2007.01480.x, 2008.

Teh, Y.A., O. Mazéas, A.R. Atwood, T. Abel, and R.C. 

Rhew, Hydrologic regulation of gross methyl chloride and methyl bromide uptake from Alaskan Arctic tundra, Global Change Biol., 15 (2), 330-345, 

doi: 10.1111/j.1365-2486.2008.01749.x, 2009.

Theys, N., M. Van Roozendael, F. Hendrick, C. Fayt, 

C. Hermans, J.-L. Barray, F. Goutail, J.-P. Pommereau, and M. De Mazière, Retrieval of stratospheric and tropospheric BrO columns from multiaxis DOAS measurements at Reunion Island (21°S, 

56°E), Atmos. Chem. Phys., 7 (18), 4733-4749, doi: 

10.5194/acp-7-4733-2007, 2007.

Thomas, V.M., J.A. Bedford, and R.J. Cicerone, Bromine 

emissions from leaded gasoline, Geophys. Res. 

Lett., 24 (11), 1371-1374, 1997.

Tokuhashi, K., A. Takahashi, M. Kaise, and S. Kondo, Rate constants for the reactions of OH radicals with CH3OCF2CHFCl, CHF2OCF2CHFCl, 

CHF2OCHClCF3, and CH3CH2OCF2CHF2, J. 

Geophys. Res., 104 (D15), 18681-18688, doi: 

10.1029/1999JD900278, 1999.

Tokuhashi, K., L. Chen, S. Kutsuna, T. Uchimaru, M. Sugie, and A. Sekiya, Environmental assessment of 

CFC alternatives: Rate constants for the reactions 

of OH radicals with fluorinated compounds, J. 

Fluor. Chem., 125 (11), 1801-1807, doi: 10.1016/j.

jfluchem.2004.09.013, 2004.

Trentmann, J., G. Luderer, T. Winterrath, M.D. Fromm, 

R. Servranckx, C. Textor, M. Herzog, H.-F. Graf, 

and M.O. Andreae, Modeling of biomass smoke injection into the lower stratosphere by a large forest 

fire (Part I): Reference simulation, Atmos. Chem. 

Phys., 6 (12), 5247-5260, doi: 10.5194/acp-6-5247-

2006, 2006.

Trudinger, C.M., D.M. Etheridge, G.A. Sturrock, P.J. Fraser, P.B. Krummel, and A. McCulloch, Atmospheric 

histories of halocarbons from analysis of Antarctic 

firn air: Methyl bromide, methyl chloride, chloroform, and dichloromethane, J. Geophys. Res., 109, 

D22310, doi: 10.1029/2004JD004932, 2004.

UNEP (United Nations Environment Programme), 2006 

Report of the Halons Technical Options Committee 

2006 Assessment, edited by D. Catchpole and D. 

Verdonik, Nariobi, Kenya, 2007a.

UNEP (United Nations Environment Programme), 2006 

Report of the Methyl Bromide Technical Options 

Committee 2006 Assessment, coordinated by M. 

Pizano, I. Porter, M. Marcotte, M. Besri, and J. 

Banks, Nariobi, Kenya, 2007b.

UNEP (United Nations Environment Programme), data 

reported to the ozone secretariat, available online 

at: http://ozone.unep.org/Data_Reporting/, 2010.

UNEP/TEAP (United Nations Environment Programme/

Technology and Economic Assessment Panel), 

1998 Assessment Report of the Technology and 

Economic Assessment Panel (TEAP), 286 pp., Nariobi, Kenya, available at: http://ozone.unep.org/teap/

Reports/TEAP_Reports/TEAPAS98.pdf, 1998.

UNEP/TEAP (United Nations Environment Programme/

Technology and Economic Assessment Panel), 

Task Force on Emissions Discrepancies Report, 

co-chaired by P. Ashford, L. Kuijpers, and S. 

Montzka, Nairobi, Kenya, 2006.

UNEP/TEAP (United Nations Environment Programme/

Technology and Economic Assessment Panel), 

Report of the Technology and Economic Assessment Panel (TEAP), Progress report, Volume 2, 

coordinated by L. Kuipers and M. Seki, Nairobi, 

Kenya, 2010.

van der Werf, G.R., J.T. Randerson, G.J. Collatz, L. Giglio, P.S. Kasibhatla, A.F. Arellano Jr., S.C. Olsen, 

and E.S. Kasischke, Continental-scale partitioning 

of fire emissions during the 1997 to 2001 El Niño/

La Niña period, Science, 303 (5654), 73-76, doi: 

10.1126/science.1090753, 2004.

Van Roozendael, M., T. Wagner, A. Richter, I. Pundt, 

D.W. Arlander, J.P. Burrows, M. Chipperfield, 

C. Fayt, P.V. Johnston, J.-C. Lambert, K. Kreher, 

K. Pfeilsticker, U. Platt, J.-P. Pommereau, B.-M. 

Sinnhuber, K.K. Tørnkvist, and F. Wittrock, In-
1.106

Chapter 1

tercomparison of BrO measurements from ERS-2 

GOME, ground-based and balloon platforms, Adv. 

Space Res., 29 (11), 1661-1666, 2002.

Varner, R.K., P.M. Crill, and R.W. Talbot, Wetlands: 

A potentially significant source of atmospheric 

methyl bromide and methyl chloride, Geophys. 

Res. Lett., 26 (16), 2433-2436, 1999a.

Varner, R.K., P.M. Crill, R.W. Talbot, and J.H. Shorter, 

An estimate of the uptake of atmospheric methyl 

bromide by agricultural soils, Geophys. Res. Lett., 

26 (6), 727-730, 1999b.

Varner, R.K., M.L. White, C.H. Mosedale, and P.M. 

Crill, Production of methyl bromide in a temperate forest soil, Geophys. Res. Lett., 30, 1521, doi: 

10.1029/2002GL016592, 2003.

Vecchi, G.A., and B.J. Soden, Global warming and the 

weakening of the tropical circulation, J. Climate, 

20 (17), 4316-4340, 2007.

Velders, G.J.M., D.W. Fahey, J.S. Daniel, M. McFarland, 

and S.O. Andersen, The large contribution of projected HFC emissions to future climate forcing, 

Proc. Natl. Acad. Sci., 106 (27), 10949-10954, 2009.

Vigano, I., H. van Weelden, R. Holzinger, F. Keppler, A. 

McLeod, and T. Röckmann, Effect of UV radiation 

and temperature on the emission of methane from 

plant biomass and structural components, Biogeosciences, 5 (3), 937-947, doi: 10.5194/bg-5-937-

2008, 2008.

Volk, C.M., J.W. Elkins, D.W. Fahey, G.S. Dutton, J.M. 

Gilligan, M. Loewenstein, J.R. Podolske, K.R. 

Chan, and M.R. Gunson, Evaluation of source gas 

lifetimes from stratospheric observations, J. Geophys. Res., 102 (D21), 25543-25564, 1997.

Vollmer, M.K., S. Reimann, D. Folini, L.W. Porter, and 

L.P. Steele, First appearance and rapid growth of 

anthropogenic HFC-245fa (CHF2CH2CF3) in the 

atmosphere, Geophys. Res. Lett., 33, L20806, doi: 

10.1029/2006GL026763, 2006.

Vollmer, M.K., L.X. Zhou, B.R. Greally, S. Henne, B. 

Yao, S. Reimann, F. Stordal, D.M. Cunnold, X.C. 

Zhang, M. Maione, F. Zhang, J. Huang, and P.G. 

Simmonds, Emissions of ozone-depleting halocarbons from China, Geophys. Res. Lett., 36, L15823, 

doi: 10.1029/2009GL038659, 2009.

Wallace, D.W.R., P. Beining, and A. Putzka, Carbon tetrachloride and chlorofluorocarbons in the South 

Atlantic Ocean, 19°S, J. Geophys. Res., 99 (C4), 

7803-7819, 1994.

Wallington, T.J., W.F. Schneider, D.R. Worsnop, O.J. 

Nielsen, J. Sehested, W.J. DeBruyn, and J.A. 

Shorter, The environmental impact of CFC replacements – HFCs and HCFCs, Environ. Sci. 

Technol., 28 (7), 320A-326A, doi: 10.1021/

es00056a002, 1994.

Wallington, T.J., M.D. Hurley, J.M. Fracheboud, J.J. Orlando, G.S. Tyndall, J. Sehested, T.E. Møgelberg, 

and O.J. Nielsen, Role of excited CF3CFHO radicals in the atmospheric chemistry of HFC-134a, J. 

Phys. Chem., 100 (46), 18116-18121, doi: 10.1021/

jp9624764, 1996.

Wallington, T.J., M.D. Hurley, J. Xia, D.J. Wuebbles, S. 

Sillman, A. Ito, J.E. Penner, D.A. Ellis, J. Martin, S.A. Mabury, O.J. Nielsen, and M.P. Sulbaek 

Andersen, Formation of C7F15COOH (PFOA) and 

other perfluorocarboxylic acids during the atmospheric oxidation of 8:2 fluorotelomer alcohol, Environ. Sci. Technol., 40 (3), 924-930, doi: 10.1021/

es051858x, 2006.

Wallington, T.J., M.P. Sulbaek Andersen, and O.J. 

Nielsen, Estimated photochemical ozone creation 

potentials (POCPs) of CF3CF=CH2 (HFO-1234yf) 

and related hydrofluoroolefins (HFOs), Atmos. 

Environ., 44 (11), 1478-1481, doi: 10.1016/j.atmosenv.2010.01.040, 2010.

Wan, D., J. Xu, J. Zhang, X. Tong, and J. Hu, Historical and projected emissions of major halocarbons 

in China, Atmos. Environ., 43 (36), 5822-5829, doi: 

10.1016/j.atmosenv.2009.07.052, 2009.

Wang, J., R. Li, Y. Guo, P. Qin, and S. Sun, The flux 

of methyl chloride along an elevational gradient of a coastal salt marsh, Eastern China, Atmos. 

Environ., 40 (34), 6592-6605, doi: 10.1016/j.atmosenv.2006.05.065, 2006.

Wang, P.K., The thermodynamic structure atop a penetrating convective thunderstorm, Atmos. Res., 83, 254-

262, doi: 10.1016/j.atmosres.2005.08.010, 2007.

Wang, P.K., M. Setvák, W. Lyons, W. Schmid, and H.-M. 

Lin, Further evidences of deep convective vertical 

transport of water vapor through the tropopause, 

Atmos. Res., 94 (3), 400-408, doi: 10.1016/j.atmosres.2009.06.018, 2009.

Warwick, N.J., J.A. Pyle, G.D. Carver, X. Yang, N.H. Savage, F.M. O’Connor, and R.A. Cox, Global modeling of biogenic bromocarbons, J. Geophys. Res.,

111, D24305, doi: 10.1029/2006JD007264, 2006.

Waugh, D., and T. Hall, Age of stratospheric air: Theory, 

observations, and models, Rev. Geophys., 40, 1010, 

doi: 10.1029/2000RG000101, 2002.

Weiss, R.F., J. Mühle, P.K. Salameh, and C.M. 

Harth, Nitrogen trifluoride in the global atmosphere, Geophys. Res. Lett., 35, L20821, doi: 

10.1029/2008GL035913, 2008.

White, M.L., R.K. Varner, P.M. Crill, and C.H. Mosedale, 

Controls on the seasonal exchange of CH3Br in 

temperate peatlands, Global Biogeochem. Cycles,

19, GB4009, doi: 10.1029/2004GB002343, 2005.

Williams, J., V. Gros, E. Atlas, K. Maciejczyk, A. Batsaikhan, H.F. Schöler, C. Forster, B. Quack, N. 
1.107

ODSs and Related Chemicals

Yassaa, R. Sander, and R. Van Dingenen, Possible 

evidence for a connection between methyl iodide 

emissions and Saharan dust, J. Geophys. Res., 112, 

D07302, doi: 10.1029/2005JD0006702, 2007.

Wilson, E.W., Jr., W.A. Hamilton, and H.R. Mount, Rate 

constants for the reactions of hydroxyl radical with 

several fluoroethers by the relative rate method, J. 

Phys. Chem. A, 111 (9), 1610-1617, doi: 10.1021/

jp068355d, 2007.

Wishkerman, A., S. Gebhardt, C.W. McRoberts, J.T.G. 

Hamilton, J. Williams, and F. Keppler, Abiotic 

methyl bromide formation from vegetation, and 

its strong dependence on temperature, Environ.

Sci. Tech., 42 (18), 6837-6842, doi: 10.1021/

es800411j, 2008.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 2006, Global 

Ozone Research and Monitoring Project–Report 

No. 50, 572 pp., Geneva, Switzerland, 2007.

Worton, D.R., W.T. Sturges, J. Schwander, R. Mulvaney, 

J.-M. Barnola, and J. Chappellaz, 20th century 

trends and budget implications of chloroform and 

related tri- and dihalomethanes inferred from firn 

air, Atmos. Chem. Phys., 6 (10), 2847-2863, doi: 

10.5194/acp-6-2847-2006, 2006.

Worton, D.R., W.T. Sturges, L.K. Gohar, K.P. Shine, P. 

Martinerie, D.E. Oram, S.P. Humphrey, P. Begley, L. Gunn, J.-M. Barnola, J. Schwander, and R. 

Mulvaney, Atmospheric trends and radiative forcings of CF4 and C2F6 inferred from firn air, Environ. Sci. Technol., 41 (7), 2184-2189, doi: 10.1021/

es061710t, 2007.

Wuebbles, D.J., D. Youn, K. Patten, D. Wang, and M. Martinez-Aviles, Metrics for ozone and climate: Threedimensional modeling studies of ozone depletion 

potentials and indirect global warming potentials, in 

Twenty Years of Ozone Decline, Proceedings of the 

Symposium for the 20th Anniversary of the Montreal 

Protocol, edited by C. Zerefos, G. Contopoulos, and 

G. Skalkeas, 297-326, Springer, Netherlands, doi: 

10.1007/978-90-481-2469-5_23, 2009.

Xiao, X., Optimal Estimation of the Surface Fluxes of 

Chloromethanes Using a 3-D Global Atmospheric 

Chemical Transport Model, Ph.D. thesis, Massachusetts Institute of Technology, Cambridge, MA, 

U.S.A., 210 pp., 2008.

Xiao, X., R.G. Prinn, P.J. Fraser, P.G. Simmonds, R.F. 

Weiss, S. O’Doherty, B.R. Miller, P.K. Salameh, 

C.M. Harth, P.B. Krummel, L.W. Porter, J. Mühle, 

B.R. Greally, D. Cunnold, R. Wang, S.A. Montzka, 

J.W. Elkins, G.S. Dutton, T.M. Thompson, J.H. 

Butler, B.D. Hall, S. Reimann, M.K. Vollmer, F. 

Stordal, C. Lunder, M. Maione, J. Arduini, and Y. 

Yokouchi, Optimal estimation of the surface fluxes of methyl chloride using a 3-D global chemical transport model, Atmos. Chem. Phys., 10 (12), 

5515-5533, doi: 10.5194/acp-10-5515-2010, 2010a.

Xiao, X., R.G. Prinn, P.J. Fraser, R.F. Weiss, P.G. Simmonds, S. O’Doherty, B.R. Miller, P.K. Salameh, 

C.M. Harth, P.B. Krummel, A. Golombek, L.W. 

Porter, J.W. Elkins, G.S. Dutton, B.D. Hall, L.P. 

Steele, R.H.J. Wang, and D.M. Cunnold, Atmospheric three-dimensional inverse modeling of regional industrial emissions and global oceanic uptake of carbon tetrachloride, Atmos. Chem. Phys.,

10 (21), 10421-10434, doi:10.5194/acp-10-10421-

2010, 2010b.

Yang, E.-S., D.M. Cunnold, M.J. Newchurch, R.J. Salawitch, M.P. McCormick, J.M. Russell, J.M. Zawodny, and S.J. Oltmans, First stage of Antarctic 

ozone recovery, J. Geophys. Res., 113, D20308, 

doi: 10.1029/2007JD009675, 2008.

Yang, X., R.A. Cox, N.J. Warwick, J.A. Pyle, G.D. Carver, F.M. O’Connor, and N.H. Savage, Tropospheric 

bromine chemistry and its impacts on ozone: A 

model study, J. Geophys. Res., 110, D23311, doi: 

10.1029/2005JD006244, 2005.

Yokouchi, Y., M. Ikeda, Y. Inuzuka, and T. Yukawa, Strong 

emission of methyl chloride from tropical plants, 

Nature, 416, 163-165, doi: 10.1038/416163a, 2002.

Yokouchi, Y., F. Hasebe, M. Fujiwara, H. Takashima, 

M. Shiotani, N. Nishi, Y. Kanaya, S. Hashimoto, 

P. Fraser, D. Toom-Sauntry, H. Mukai, and Y. 

Nojiri, Correlations and emission ratios among 

bromoform, dibromochloromethane, and dibromomethane in the atmosphere, J. Geophys. Res., 110, 

D23309, doi: 10.1029/2005JD006303, 2005.

Yokouchi, Y., S. Taguchi, T. Saito, Y. Tohjima, H. Tanimoto, and H. Mukai, High frequency measurements 

of HFCs at a remote site in east Asia and their implications for Chinese emissions, Geophys. Res. Lett.,

33, L21814, doi: 10.1029/2006GL026403, 2006.

Yokouchi, Y., T. Saito, C. Ishigaki, and M. Aramoto, 

Identification of methyl chloride-emitting plants 

and atmospheric measurements on a subtropical island, Chemosphere, 69 (4), 549-553, doi: 10.1016/j.

chemosphere.2007.03.028, 2007.

Yokouchi, Y., K. Osada, M. Wada, F. Hasebe, M. Agama, 

R. Murakami, H. Mukai, Y. Nojiri, Y. Inuzuka, D. 

Toom-Sauntry, and P. Fraser, Global distribution 

and seasonal concentration change of methyl iodide 

in the atmosphere, J. Geophys. Res., 113, D18311, 

doi: 10.1029/2008JD009861, 2008.

Young, C.J., M.D. Hurley, T.J. Wallington, and S.A. 

Mabury, Atmospheric chemistry of CF3CF2H and 

CF3CF2CF2CF2H: Kinetics and products of gasphase reactions with Cl atoms and OH radicals, 

infrared spectra, and formation of perfluorocar-
1.108

Chapter 1

boxylic acids, Chem. Phys. Lett., 473 (4-6), 251-

256, doi: 10.1016/j.cplett.2009.04.001, 2009.

Yujing, M., and A. Mellouki, Rate constants for the reactions of OH with chlorinated propanes, Phys. 

Chem. Chem. Phys., 3, 2614-2617, doi: 10.1039/

b102971c, 2001.

Yvon-Lewis, S.A., and J.H. Butler, Effect of oceanic uptake on atmospheric lifetimes of selected 

trace gases, J. Geophys. Res., 107, 4414, doi: 

10.1029/2001JD001267, 2002.

Yvon-Lewis, S.A., E.S. Saltzman, and S.A. Montzka, Recent trends in atmospheric methyl bromide: Analysis of post-Montreal Protocol variability, Atmos. 

Chem. Phys., 9 (16), 5963-5974, doi: 10.5194/acp9-5963-2009, 2009.

Zander, R., E. Mahieu, P. Demoulin, P. Duchatelet, G. Roland, C. Servais, M. De Mazière, and C.P. Rinsland, 

Evolution of a dozen non-CO2 greenhouse gases 

above central Europe since the mid-1980s, Environ. 

Sci., 2 (2-3), 295-303, 2005.

Zander, R., E. Mahieu, P. Demoulin, P. Duchatelet, G. 

Roland, C. Servais, M. De Mazière, S. Reimann, 

and C.P. Rinsland, Our changing atmosphere: Evidence based on long-term infrared solar observations at the Jungfraujoch since 1950, Sci. Tot. 

Environ., 391 (2-3), 184-195, doi: 10.1016/j.scitotenv.2007.10.018, 2008.

Zellner, R., G. Bednarek, A. Hoffmann, J.P. Kohlmann, V. 

Mörs, and H. Saathoff, Rate and mechanism of the 

atmospheric degradation of 2H-heptafluoropropane 

(HFC-227), Ber. Bunsenges. Phys. Chem., 98 (2), 

141-146, doi: 10.1002/bbpc.19940980202, 1994.

Zeng, G., J.A. Pyle, and P.J. Young, Impact of climate 

change on tropospheric ozone and its global budgets, Atmos. Chem. Phys., 8 (2), 369-387, doi: 

10.5194/acp-8-369-2008, 2008.

Zhao, Z., P.L. Laine, J.M. Nicovich, and P.H. Wine, Reactive and nonreactive quenching of O(1

D) by the 

potent greenhouse gases SO2F2, NF3, and SF5CF3, 

Proc. Natl. Acad. Sci., 107 (15), 6610-6615, doi: 

10.1073/pnas.0911228107, 2010.

Zhou, Y., H. Mao, R.S. Russo, D.R. Blake, O.W. Wingenter, K.B. Haase, J. Ambrose, R.K. Varner, R. 

Talbot, and B.C. Sive, Bromoform and dibromomethane measurements in the seacoast region of 

New Hampshire, 2002–2004, J. Geophys. Res., 

113, D08305, doi: 10.1029/2007JD009103, 2008. 
Chapter 2

Stratospheric Ozone and Surface

Ultraviolet Radiation

Coordinating Lead Authors:

A. Douglass

V. Fioletov

Lead Authors:

S. Godin-Beekmann

R. Müller

R.S. Stolarski

A. Webb

Coauthors:

A. Arola

J.B. Burkholder

J.P. Burrows

M.P. Chipperfield

R. Cordero

C. David

P.N. den Outer

S.B. Diaz

L.E. Flynn

M. Hegglin

J.R. Herman

P. Huck

S. Janjai

I.M. Jánosi

J.W. Krzyścin

Y. Liu

J. Logan

K. Matthes

R.L. McKenzie

N.J. Muthama

I. Petropavlovskikh

M. Pitts

S. Ramachandran

M. Rex

R.J. Salawitch

B.-M. Sinnhuber

J. Staehelin

S. Strahan

K. Tourpali

J. Valverde-Canossa

C. Vigouroux

Contributors:

G.E. Bodeker

T. Canty

H. De Backer

P. Demoulin

U. Feister

S.M. Frith

J.-U. Grooß

F. Hase

J. Klyft

T. Koide

M.J. Kurylo

D. Loyola

C.A. McLinden

I.A. Megretskaia

P.J. Nair

M. Palm

D. Papanastasiou

L.R. Poole

M. Schneider

R. Schofield

H. Slaper

W. Steinbrecht

S. Tegtmeier

Y. Terao

S. Tilmes

D.I. Vyushin

M. Weber

E.-S. Yang

Chapter 2

STRATOSPHERIC OZONE AND SURFACE ULTRAVIOLET RADIATION

Contents

SCIENTIFIC SUMMARY .............................................................................................................................................1

INTRODUCTION ..........................................................................................................................................................5

2.1 OZONE OBSERVATIONS ....................................................................................................................................5

2.1.1 State of Science in 2006 ............................................................................................................................5

2.1.2 Update on Methods Used to Evaluate the Effect of ODSs on Ozone .......................................................5

2.1.3 Update on Total Ozone Changes...............................................................................................................6

2.1.3.1 Measurements.............................................................................................................................6

2.1.3.2 Total Ozone Changes and Trends...............................................................................................8

2.1.4 Update on Ozone Profile Changes..........................................................................................................10

2.1.4.1 Measurements...........................................................................................................................10

2.1.4.2 Ozone Profile Changes .............................................................................................................12

2.2 POLAR OZONE....................................................................................................................................................17

2.2.1 State of Science in 2006 ..........................................................................................................................17

2.2.2 Polar Ozone Chemistry ...........................................................................................................................18

2.2.2.1 Laboratory Studies of the ClOOCl UV Absorption Spectrum .................................................19

2.2.2.2 Field Observations of Chlorine Partitioning.............................................................................21

2.2.2.3 Other Issues Related to Polar Ozone Chemistry.......................................................................25

2.2.3 Polar Stratospheric Cloud Processes.......................................................................................................25

2.2.3.1 New Observational Data Sets ...................................................................................................26

2.2.3.2 PSC Composition......................................................................................................................27

2.2.3.3 PSC Forcing Mechanisms.........................................................................................................27

2.2.3.4 Use of Proxies to Represent PSC Processes.............................................................................28

2.2.4 Arctic Polar Temperatures and Ozone ....................................................................................................29

2.2.5 Antarctic Polar Temperatures and Ozone ...............................................................................................30

2.2.6 The Onset of Antarctic Ozone Depletion................................................................................................31

2.3 SURFACE ULTRAVIOLET RADIATION .........................................................................................................31

2.3.1 State of Science in 2006 ..........................................................................................................................31

Box 2-1 Radiation Amplification Factor for Erythemal Irradiance........................................................32

2.3.2 Update on Factors Affecting UV Radiation............................................................................................33

2.3.2.1 Ozone Effects............................................................................................................................33

2.3.2.2 Other Influences on UV............................................................................................................34

2.3.3 Ground-Based and Satellite UV Data .....................................................................................................35

2.3.3.1 Ground-Based UV Measurements............................................................................................35

2.3.3.2 Ground-Based UV Reconstruction ...........................................................................................35

2.3.3.3 UV Estimates from Satellite Observations...............................................................................36

2.3.4 Long-Term Changes in UV.....................................................................................................................37

2.3.4.1 Ground-Based Observations.....................................................................................................37

2.3.4.2 Reconstructed UV Data ............................................................................................................38

2.3.4.3 Satellite Estimates of Irradiance Changes.................................................................................40

2.3.4.4 Consistency of UV Estimates from Observations, Reconstructions, and Satellite Data ..........41

2.4 INTERPRETATION OF OBSERVED OZONE CHANGES...............................................................................41

2.4.1 State of Science in 2006 ..........................................................................................................................41
2.4.2 Updates to Kinetic and Photochemical Data...........................................................................................43

2.4.2.1 Update from JPL 2002 to JPL 2006..........................................................................................43

2.4.2.2 Updates since JPL 2006............................................................................................................43

2.4.3 The Distribution and Variability of Stratospheric Ozone and Their Representation in Models.............44

2.4.3.1 Annual Cycle and Natural Variability ......................................................................................44

2.4.3.2 Solar Cycle................................................................................................................................44

2.4.3.3 Volcanic and Aerosol Effects ...................................................................................................46

2.4.3.4 Evaluation of Simulated Transport ...........................................................................................47

2.4.3.5 Evaluation of the Chemical Mechanism and Its Implementation.............................................49

2.4.3.6 Evaluation of Simulations of the Upper Troposphere/Lower Stratosphere..............................49

2.4.4 Recovery Detection and Attribution........................................................................................................50

2.4.4.1 Dynamical Contributions to Apparent Trend ...........................................................................50

2.4.4.2 Greenhouse Gas Effects on Ozone Trends ...............................................................................50

2.4.4.3 Polar Loss and Dilution to Midlatitudes...................................................................................51

2.4.5 Simulation of Ozone Changes for the Last Three Decades.....................................................................51

2.4.5.1 Total Ozone Trends...................................................................................................................51

2.4.5.2 Midlatitude Profile Trends........................................................................................................53

2.4.5.3 Tropical Profile Trends.............................................................................................................55

2.4.5.4 Polar Trends..............................................................................................................................55

2.4.5.5 Trends and Recovery in the Upper Troposphere and Lower Stratosphere...............................58

REFERENCES .............................................................................................................................................................59
2.1

Stratospheric Ozone and Surface UV

SCIENTIFIC SUMMARY

Global Ozone Observations and Interpretation

As a result of the Montreal Protocol, ozone is expected to recover from the effect of ozone-depleting substances 

(ODSs) as their abundances decline in the coming decades. The 2006 Assessment showed that globally averaged column 

ozone ceased to decline around 1996, meeting the criterion for the first stage of recovery. Ozone is expected to increase 

as a result of continued decrease in ODSs (second stage of recovery). This chapter discusses recent observations of ozone 

and ultraviolet radiation in the context of their historical records. Natural variability, observational uncertainty, and 

stratospheric cooling necessitate a long record in order to attribute an ozone increase to decreases in ODSs. Table S2-1 

summarizes ozone changes since 1980.

The primary tools used in this Assessment for prediction of ozone are chemistry-climate models (CCMs). These 

CCMs are designed to represent the processes determining the amount of stratospheric ozone and its response to changes 

in ODSs and greenhouse gases. Eighteen CCMs have been recently evaluated using a variety of process-based comparisons to measurements. The CCMs are further evaluated here by comparison of trends calculated from measurements with 

trends calculated from simulations designed to reproduce ozone behavior during an observing period.

Total Column Ozone

•	 Average total ozone values in 2006–2009 have remained at the same level for the past decade, about 3.5%

and 2.5% below the 1964–1980 averages respectively for 90°S–90°N and 60°S–60°N. Average total ozone from 

CCM simulations behaves in a manner similar to observations between 1980 and 2009. The average column ozone 

for 1964–1980 is chosen as a reference for observed changes for two reasons: 1) reliable ground-based observations sufficient to produce a global average are available in this period; 2) a significant trend is not discernible in the 

observations during this period.

•	 Southern Hemisphere midlatitude (35°S–60°S) annual mean total column ozone amounts over the period

2006–2009 have remained at the same level as observed during 1996–2005, approximately 6% below the

1964–1980 average. Simulations by CCMs also show declines of the same magnitude between 1980 and 1996, and 

minimal change after 1996, thus both observations and simulations are consistent with the expectations of the impact 

of ODSs on southern midlatitude ozone.

•	 Northern Hemisphere midlatitude (35°N–60°N) annual mean total column ozone amounts over the period

2006–2009 have remained at the same level as observed during 1998–2005, approximately 3.5% below

the 1964–1980 average. A minimum about 5.5% below the 1964–1980 average was reached in the mid-1990s. 

Simulations by CCMs agree with these measurements, again showing the consistency of data with the expected 

impact of ODSs. The simulations also indicate that the minimum in the mid-1990s was primarily caused by the ozone 

response to effects of volcanic aerosols from the 1991 eruption of Mt. Pinatubo.

•	 The latitude dependence of simulated total column ozone trends generally agrees with that derived from

measurements, showing large negative trends at Southern Hemisphere mid and high latitudes and Northern

Hemisphere midlatitudes for the period of ODS increase. However, in the tropics the statistically significant 

range of trends produced by CCMs (−1.5 to −4 Dobson units per decade (DU/decade)) does not agree with the trend 

obtained from measurements (+0.3 ± 1 DU/decade).

Ozone Profiles

•	 Northern Hemisphere midlatitude (35°N–60°N) ozone between 12 and 15 km decreased between 1979 and

1995, and increased between 1996 and 2009. The increase since the mid-1990s is larger than the changes expected 

from the decline in ODS abundances.
2.2

Chapter 2

•	 Northern Hemisphere midlatitude (35°N–60°N) ozone between 20 and 25 km declined during 1979–1995 and has

since ceased to decline. Observed increases between 1996 and 2008 are statistically significant at some locations but 

not globally.

•	 Northern Hemisphere midlatitude (35°N–60°N) ozone between 35 and 45 km measured using a broad range of

ground-based and satellite instruments ceased to decline after the mid-1990s, consistent with the leveling off

of ODS abundances. All data sets show a small ozone increase since that time, with varying degrees of statistical 

significance but this increase cannot presently be attributed to ODS decrease because of observational uncertainty, 

natural ozone variability, and stratospheric cooling. CCMs simulate the ozone response to changes in ODSs and 

increases in greenhouse gases; analysis of CCM results suggests that longer observational records are required to 

separate these effects from each other and from natural variability.

•	 In the midlatitude upper stratosphere (35–45 km) of both hemispheres, the profile ozone trends derived from

most CCMs from 1980 to 1996 agree well with trends deduced from measurements. The agreement in both magnitude and shape of the ozone trends provides evidence that increases in ODSs between 1980 and 1996 are primarily 

responsible for the observed behavior.

•	 In the tropical lower stratosphere, all simulations show a negative ozone trend just above the tropopause,

centered at about 18–19 km (70–80 hectoPascals, hPa), due to an increase in upwelling. The simulated trends in 

the lower tropical stratosphere are consistent with trends deduced for 1985–2005 from Stratospheric Aerosol and Gas 

Experiment (SAGE II) satellite data, although uncertainties in the SAGE II trends are large. The near-zero trend in 

tropical total ozone measurements is inconsistent with the negative trend found in the integrated SAGE I + SAGE II 

stratospheric profiles. The tropospheric ozone column does not increase enough to resolve this discrepancy.

Table S2-1. Summary of ozone changes estimated from observations.

Column Ozone 12–15 km 20–25 km 35–45 km Comment

Data Sources Ground-based, 

satellite Ozonesondes Ozonesondes,satellites, FTIR

Satellites, 

Umkehrs, FTIR 

Northern

midlatitudes

1980–1996

Declined by about 

6%

Declined by 

about 9%

Declined by 

about 7%

Declined by 

about 10%

1992–1996 

column and lower 

stratosphere data 

affected by Mt. 

Pinatubo

Northern

midlatitudes

1996–2009

Increased from 

the minimum 

values by about 

2% by 1998 

and remained at 

the same level 

thereafter 

Increased by 

about 6%

Increased by

about 2.5%

Increased by 

1 to 2%, but 

uncertainties are 

large

Southern

midlatitudes

1980–1996

Declined by 6% No information Declined by 

about 7%

Declined by 

about 10%

Southern

midlatitudes

1996–2009

Remained at 

approximately the 

same level 

No statistically 

significant 

changes

No statistically 

significant 

changes

Increased by 

1 to 3%, but 

uncertainties are 

large
2.3

Stratospheric Ozone and Surface UV

Polar Ozone Observations and Interpretation

•	 The Antarctic ozone hole continued to appear each spring from 2006 to 2009. This is expected because decreases 

in stratospheric chlorine and bromine have been moderate over the last few years. Analysis shows that since 1979 

the abundance of total column ozone in the Antarctic ozone hole has evolved in a manner consistent with the time 

evolution of ODSs. Since about 1997 the ODS amounts have been nearly constant and the depth and magnitude of 

the ozone hole have been controlled by variations in temperature and dynamics. The October mean column ozone 

within the vortex has been about 40% below 1980 values for the past fifteen years.

•	 Arctic winter and spring ozone loss has varied between 2007 and 2010, but remained in a range comparable

to the values that have prevailed since the early 1990s. Chemical loss of about 80% of the losses observed in the 

record cold winters of 1999/2000 and 2004/2005 has occurred in recent cold winters.

•	 Recent laboratory measurements of the chlorine monoxide dimer (ClOOCl) dissociation cross section and

analyses of observations from aircraft and satellites have reaffirmed the fundamental understanding that

polar springtime ozone depletion is caused primarily by the ClO + ClO catalytic ozone destruction cycle, with

significant contributions from the BrO + ClO cycle.

•	 Polar stratospheric clouds (PSCs) over Antarctica occur more frequently in early June and less frequently in

September than expected based on the previous satellite PSC climatology. This result is obtained from measurements by a new class of satellite instruments that provide daily vortex-wide information concerning PSC composition 

and occurrence in both hemispheres. The previous satellite PSC climatology was developed from solar occultation 

instruments that have limited daily coverage.

•	 Calculations constrained to match observed temperatures and halogen levels produce Antarctic ozone losses

that are close to those derived from data. Without constraints, CCMs simulate many aspects of the Antarctic 

ozone hole, however they do not simultaneously produce the cold temperatures, isolation from middle latitudes, deep 

descent, and high amounts of halogens in the polar vortex. Furthermore, most CCMs underestimate the Arctic ozone 

loss that is derived from observations, primarily because the simulated northern winter vortices are too warm.

Ultraviolet Radiation

Ground-based measurements of solar ultraviolet (UV) radiation (wavelength 280–400 nanometers) remain limited 

both spatially and in duration. However, there have been advances both in reconstructing longer-term UV records from 

other types of ground-based measurements and in satellite UV retrievals. Where these UV data sets coincide, long-term 

changes agree, even though there may be differences in instantaneous, absolute levels of UV.

•	 Ground-based UV reconstructions and satellite UV retrievals, supported in the later years by direct groundbased UV measurements, show that erythemal (“sunburning”) irradiance over midlatitudes has increased

since the late 1970s, in qualitative agreement with the observed decrease in column ozone. The increase in 

satellite-derived erythemal irradiance over midlatitudes during 1979–2008 is statistically significant, while there are 

no significant changes in the tropics. Satellite estimates of UV are difficult to interpret over the polar regions.

•	 In the Antarctic, large ozone losses produce a clear increase in surface UV radiation. Ground-based measurements show that the average spring erythemal irradiance for 1990–2006 is up to 85% greater than the modeled 

irradiance for 1963–1980, depending on site. The Antarctic spring erythemal irradiance is approximately twice that 

measured in the Arctic for the same season.

•	 Clear-sky UV observations from unpolluted sites in midlatitudes show that since the late 1990s, UV irradiance

levels have been approximately constant, consistent with ozone column observations over this period.
2.4

Chapter 2

•	 Surface UV levels and trends have also been significantly influenced by clouds and aerosols, in addition to

stratospheric ozone. Daily measurements under all atmospheric conditions at sites in Europe and Japan show that 

erythemal irradiance has continued to increase in recent years due to net reductions in the effects of clouds and aerosols. In contrast, in southern midlatitudes, zonal and annual average erythemal irradiance increases due to ozone 

decreases since 1979 have been offset by almost a half due to net increases in the effects of clouds and aerosols. 
2.5

Stratospheric Ozone and Surface UV

INTRODUCTION

This chapter presents information on several topics 

but is conceptually organized around a single question: is 

the Montreal Protocol working? This chapter is focused 

on the observational record and interpretation thereof up 

to the present, and consolidates information found in three 

separate chapters in the previous Assessment (WMO, 

2007): Chapter 3 (“Global Ozone: Past and Present,” 

Chipperfield and Fioletov et al., 2007), Chapter 4 (“Polar 

Ozone: Past and Present,” Newman and Rex et al., 2007), 

and Chapter 7 (“Surface Ultraviolet Radiation: Past, Present and Future,” Bais and Lubin et al., 2007). There are 

four sections in this chapter: Ozone Observations, Polar 

Ozone, Surface Ultraviolet Radiation, and Interpretation 

of Observed Ozone Changes. Each section begins with a 

summary of WMO (2007) followed by a combination of 

updates to the observational records and longer discussion 

of new discoveries and observations.

2.1 OZONE OBSERVATIONS

2.1.1 State of Science in 2006

The long-term changes in global ozone were reviewed in Chapter 3 of WMO (2007). From the analysis 

of data from multiple sources, it was shown that the global 

mean total column ozone values for the period 2002–2005 

had stabilized to values similar to those observed in 1998–

2001, at approximately 3.5% below the 1964–1980 average 

values. Differences between the Northern and Southern 

Hemispheres (NH and SH) were noted, with ozone average values respectively 3% and 5.5% below their pre-1980 

average values. The time series behavior of total ozone 

column was also shown to be different in both hemispheres 

during the 1990s. Ozone showed a minimum in the NH 

around 1993 followed by an increase, while it decreased 

through the late 1990s in the SH and leveled off in about 

2000. Seasonal differences between ozone changes over 

midlatitude regions in both hemispheres were also noticed. 

The changes with respect to the pre-1980 values were larger in spring in the NH, while no seasonal dependence was 

found in the SH. Over the tropics, no change in column 

ozone values was found, which was consistent with the 

findings of WMO (2003).

Because total ozone column was no longer decreasing in most observations, several methods were discussed 

in WMO (2003) and WMO (2007) for the evaluation of 

ozone trends. Previous Assessments had described longterm ozone changes due to chemical destruction by ozonedepleting substances (ODSs) in terms of linear trends 

estimated using multiple regression analysis. Because the 

change in ODSs after the mid-1990s was no longer linear 

with time, other methods were proposed, e.g., the piecewise linear trend model in which different linear fits are 

used before and after a turning point, and the fit to the 

equivalent effective stratospheric chlorine (EESC) function (see Chapter 1 of this Assessment). Such methods 

have been used in most recent studies on ozone trends and 

are also discussed in the present Assessment.

Regarding changes in the vertical ozone distribution, satellite and ground-based measurements showed 

that in the upper stratosphere, the ozone decrease had 

stopped and ozone values were relatively constant since 

1995. Similar stabilization was found in the lower stratosphere between 20 and 25 kilometers (km) altitude. In the 

lowermost stratosphere below 15 km altitude in the NH, a 

significant increase was found from 1996, after the strong 

decrease observed between 1979 and 1995. This change 

in the lowermost stratosphere had a substantial impact on 

the total ozone column. Such an ozone increase was not 

observed in the SH. The lowermost stratosphere is defined and discussed in more detail in Section 2.4 below.

All studies in WMO (2007) pointed out the stabilization of ozone both in total column and in the vertical 

distribution at various levels, various locations, and at the 

global scale. They concurred that the first stage of recovery (i.e., slowing of ozone decline attributable to ODS 

changes) had already occurred and that the second stage 

(i.e., onset of ozone increase) was expected to become evident within the next two decades.

2.1.2 Update on Methods Used to Evaluate 

the Effect of ODSs on Ozone

As discussed in previous Assessments, the longterm and short-term variability of ozone in the stratosphere is generally estimated using multi-regression statistical models that quantify the relationship between ozone 

and different explanatory variables describing natural or 

anthropogenic forcings (e.g., SPARC, 1998). The longterm trend components representing the effect of ODSs 

are extracted simultaneously with other regression terms 

and autocorrelated noise.

To describe the long-term trend in ozone that is 

related to ODSs, the equivalent effective stratospheric 

chlorine (EESC) (see Section 1.4.4 of Chapter 1) is commonly used as a proxy in statistical models (Stolarski 

et al., 2006; Dhomse et al., 2006; Brunner et al., 2006; 

Randel and Wu, 2007; Wohltmann et al., 2007; Mäder 

et al., 2007; Vyushin et al., 2007; Harris et al., 2008). 

Statistical methods are used to quantify the relationship 

between ozone changes and EESC, and verify whether the 

EESC-related term is statistically significant. Analysis 

of the residuals, for example, by the Cumulative Sum 

of Residuals (CUSUM) technique (Reinsel et al., 2002; 
2.6

Chapter 2

Newchurch et al., 2003) then can be used to check that a 

statistical model with the EESC term adequately describes 

the observed ozone changes.

The EESC depends on latitude and altitude. Moreover, the present estimates of EESC are different from 

those used, for example, in WMO (2003), as discussed in 

Section 1.4.4 of Chapter 1. As a result, all of the ozone 

trend studies mentioned above did not use the same EESC 

function. While a particular shape of the EESC curve has 

little effect on the EESC-based trend estimates in the past 

(they all represent a linear decline during the 1980s and 

early 1990s with leveling off thereafter), the shape of an 

EESC function will have more impact on the estimated 

trend value as time moves from the EESC turning point. 

The shape of the EESC curve is particularly important for 

the detectability of future trends (Vyushin et al., 2010). 

On the other hand, once the EESC shape is specified, the 

sensitivity of ozone to EESC obtained from this type of 

statistical analysis varies little as a result of small differences in the length of record.

A statistically significant EESC-related term can be 

used as evidence of the ODS-related destruction of ozone. 

The EESC was a linear function of time in the 1980s and 

thus the EESC fit to ozone can be expressed in terms of 

linear changes at that time, with results reported in ozone 

changes (% or Dobson units (DU)) per decade. WMO 

(2007, Section 3.2.1) discussed ozone trends in terms of 

EESC. Adding four more years to 25-year-long observation records discussed in WMO (2007) does not change 

the trend estimates for that period significantly. Similarly, the EESC is nearly a linear function in the 2000s 

and therefore the expected rate of ozone increase during 

the declining phase of the EESC can be expressed in % or 

DU per decade.

WMO (2007) concluded that the first stage of the 

ozone recovery, i.e., 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, 

had already occurred. The second stage of the ozone recovery or the onset of ozone increases (turnaround) is 

identified as the occurrence of statistically significant 

increases in ozone above previous minimum values due to 

declining EESC.

The ozone increase after the minimum can be 

estimated by fitting the data with a linear function or by 

calculating a piecewise linear trend (PWLT) with a turning point near the EESC maximum (Reinsel et al., 2005; 

Miller et al., 2006; Vyushin et al., 2007; S.-K. Yang et 

al., 2009). The slope estimated from ozone data during 

the declining phase of EESC should agree with the slope 

expected from the EESC fit if the ozone increase is indeed 

related to the EESC decline.

As discussed by WMO (2007, Section 3.4.2) a sizable fraction of the long-term ozone changes, particularly 

over northern mid and high latitudes, can be related to 

dynamical processes. Estimation of ozone trends requires 

a proper accounting for the effect of these processes on 

ozone. One approach is to add more terms to the statistical model used for trend calculations using a purely 

statistical approach and letting the regression model find 

the best proxies (e.g., Mäder et al., 2007) or by adding 

proxies based on possible physical processes that cause 

the ozone changes (e.g., Wohltmann et al., 2007). However the physical mechanisms underlying these additional 

terms are often not well understood, and therefore it is difficult to account for them properly in a statistical model. 

This issue is addressed in detail in Section 2.4. Another 

approach is to consider the contribution from dynamical 

processes as noise. This results in a larger uncertainty in 

the trend estimates and also requires an additional analysis 

of the autocorrelation function of the residuals (Vyushin 

et al., 2007). In both approaches, the eleven-year solar 

activity cycle and the quasi-biennial oscillation (QBO) are 

typically included in the statistical model because these 

oscillations are located in a narrow frequency range.

2.1.3 Update on Total Ozone Changes

2.1.3.1 Measurements

Ground-Based Measurements

Dobson, Brewer, and filter instruments provide 

long-term ground-based total ozone time series. The 

instrumental precision of well maintained Dobson and 

Brewer instruments was recently estimated by Scarnato et 

al. (2010) to be respectively 0.5% and 0.15% (1-sigma). 

When comparing ground-based total ozone measurements 

with satellite overpass data, the standard deviation of 

monthly differences was on average about 1.5% and within 

0.6–2.6% for 90% of Dobson and Brewer network stations 

and on average about 2% and within 1.5–3.5% for 90% of 

stations equipped with filter instruments M-124 (Fioletov 

et al., 2008). The agreement between various instruments 

can be further improved as new ozone absorption cross 

sections are adopted (Scarnato et al., 2009). A recently 

established committee is presently addressing the issue of 

ozone cross sections used in ground-based and satellite 

measurements (see http://igaco-o3.fmi.fi/ACSO/). Since 

the end of the 1980s, other instruments have been implemented for the monitoring of total ozone. Long-term and 

regular ground-based Fourier transform infrared (FTIR) 

measurements are performed at many stations around the 

world and these data were used to assess ozone trends 

over Western Europe from 79°N to 28°N (Vigouroux et 

al., 2008). The precision of FTIR ozone total columns 
2.7

Stratospheric Ozone and Surface UV

is about 4%, but it has been demonstrated that it can 

reach 1 DU in some conditions (Schneider et al., 2008). 

No calibration is needed, but the instrumental line shape 

must be known in order to avoid introducing a bias in the 

ozone retrievals. UV-Visible spectrophotometers such as 

the System d’ Analyse par Observation Zenitale (SAOZ) 

instruments (Pommereau and Goutail, 1988) retrieve total 

ozone as well as nitrogen dioxide (NO2) column amounts 

from zenith sky measurements using Differential Optical 

Absorption Spectroscopy (DOAS). A new version of the 

zenith-sky retrieval algorithm using improved air mass 

factors was recently introduced. SAOZ observations were 

used here in addition to Dobson, Brewer, and filter instrument data to form the ground-based zonal mean data set as 

described by Fioletov et al. (2002). This data set with the 

list of contributed stations is available from http://woudc.

org/data_e.html.

Satellite Measurements

Satellite instruments have observed the total ozone 

distributions at the global scale since 1970, when the 

Nimbus 4 satellite was launched with the Backscatter 

Ultraviolet (BUV) instrument onboard. To date, the longest total ozone records are provided by the series of Total 

Ozone Mapping Spectrometer (TOMS) and Solar Backscatter Ultraviolet 2 (SBUV/2) instruments. Since 2004, 

the TOMS total ozone record has been taken over by the 

Ozone Monitoring Instrument (OMI), an instrument on 

the Aura satellite. TOMS, BUV, and SBUV/2 data presented here are retrieved with the version 8 algorithm 

(Bhartia et al., 2004; Flynn, 2007). There are two operationally available OMI satellite total ozone column data 

products, based on the OMI-TOMS and the OMI-DOAS 

retrieval algorithms, but outputs of the OMI-TOMS algorithm agree better with the most accurate ground-based 

measurements than those for the OMI-DOAS algorithm 

(Balis et al., 2007). The TOMS algorithm uses only two 

wavelengths (317.5 and 331.2 nanometers (nm)) to derive 

total ozone (four other wavelengths are used for diagnostics and error correction). The version 8.5 OMI algorithm 

is similar to the TOMS version 8 algorithm and is used to 

process OMI data presented here.

In order to obtain long-term total ozone records, 

several data sets merging various satellite ozone records 

have been constructed. The TOMS+OMI+SBUV(/2) 

merged ozone data set (MOD) (Stolarski and Frith, 2006), 

used in WMO (2007), has been updated through December of 2009. The input now includes version 8.5 data from 

OMI and version 8.0 data from NOAA-17 SBUV/2. Data 

from 1970 through 1972 have also been added from the 

Nimbus 4 BUV experiment in 1970–1977. The merged 

ozone data set (MOD) can be obtained at http://acdb-ext.

gsfc.nasa.gov/Data_services/merged/.

Version 8 ozone retrievals from Nimbus 7 SBUV, 

and NOAA-9, -11, -14, -16, -17, and -18 SBUV/2 instruments were used in a NOAA cohesive SBUV(/2) total 

ozone data set (S.K. Yang et al., 2009) available at ftp://

ftp.cpc.ncep.noaa.gov/long/SBUV_v8_Cohesive.

The European instruments Global Ozone Monitoring 

Experiment (GOME) on the European Remote Sensing Satellite (ERS-2) (1995–2003, global coverage), Scanning Imaging Absorption Spectrometer for Atmospheric Cartography (SCIAMACHY) on the Environmental Satellite 

(Envisat; 2002–present), and GOME-2 on Meteorological 

Operational satellite (MetOp)-A (2006–present) apply the 

DOAS algorithm technique in the continuous 325–335 nm 

wavelength range (Burrows et al., 1999) to retrieve total 

ozone estimates. Different types of DOAS algorithms have 

been developed: WFDOAS (Coldewey-Egbers et al., 2005), 

TOGOMI/TOSOMI (Eskes et al., 2005), and SDOAS/

GDOAS/GDP (Van Roozendael et al., 2006). By comparing to Brewer/Dobsons and other satellite data, all algorithms applied to GOME were shown to be in good agreement (Weber et al., 2005; Balis et al., 2007; Fioletov et al., 

2008). Overall good agreement was also found in the comparison of SCIAMACHY total ozone to ground data and 

other satellite data over more than six years (Lerot et al., 

2009). However, a downward drift of total ozone from 

SCIAMACHY with respect to GOME and other correlative 

data has been identified that is independent of the algorithm 

used (Lerot et al., 2009; Loyola et al., 2009a). GOME-2 has 

almost three years of total ozone data. First validation 

results have been reported (Antón et al., 2009). A merged 

data set from GOME, SCIAMACHY, and GOME-2 by successive scaling of SCIAMACHY and GOME-2 monthlymean zonal mean data to GOME is described in Loyola et 

al. (2009a). They report that a scaling of +2 to +3% was 

required to match GOME-2 to the GOME data record.

The GOME-SCIAMACHY data is based on combined GOME, SCIAMACHY, and GOME-2 records, 

with SCIAMACHY and GOME-2 records adjusted using 

a stable record of the GOME instrument (although with a 

limited coverage after 2003). While multiple versions of 

the data processing algorithm and merged data sets exist 

(Weber et al., 2007; Loyola et al., 2009a), they produce 

nearly identical records of zonal monthly-mean ozone values.

Measurements from four TOMS instruments, 

GOME, four SBUV(/2) instruments, and OMI are used 

to produce the New Zealand National Institute of Water 

and Atmospheric Research (NIWA) combined total 

ozone data set (Bodeker et al., 2005; Müller et al., 2008). 

Offsets and drifts between all of the satellite-based data 

sets are removed through intercomparisons with the 
2.8

Chapter 2

Dobson and Brewer ground-based network. The NIWA 

data set is available from http://www.bodekerscientific.

com/data/ozone.

2.1.3.2 Total Ozone Changes and Trends

The quasi-global (60°S–60°N) ozone record from 

the MOD is shown in Figure 2-1. The annual variation 

and an 11-year periodical component are evident from the 

plot and are discussed in detail in WMO, 2007 (Chipperfield and Fioletov et al., 2007). The total ozone deviations 

for the 60°S–60°N, 90°S–90°N, 25°S–25°N, 35°N–60°N, 

and 35°S–60°S latitude belts are shown in Figure 2-2. The 

approach used in Fioletov et al. (2002) and WMO (2007) 

is again used here. Five data sets of 5°-wide zonal averages of total ozone values are analyzed in this Assessment. Area-weighted annual averages are calculated for 

different latitude belts and for the globe. All panels of 

Figure 2-2 indicate that average total ozone deviations in 

2006–2009 display very little change as compared to the 

2002–2005 values reported in WMO (2007). The global 

and 60°S–60°N averages were about 3.5% and 2.5% below the 1964–1980 average values, respectively. The total 

column ozone for 1964–1980 is chosen as a reference for 

observed changes for two reasons: (1) reliable groundbased observations sufficient to produce a global average 

are available in this period; and (2) a significant trend is 

not discernible in the observations during this period. In 

midlatitude regions of both hemispheres, ozone values in 

the NH and SH stabilized at respectively about 3.5% and 

6% lower than the 1964–1980 average, with little sign of 

increase in recent years.

Several authors have examined the zonally averaged total ozone data and find statistically significant positive trends since the second half of the 1990s. S.-K. Yang 

et al. (2009) find a positive trend of about 1.2 ± 0.8%/decade 

for the period 1996–2007 in the averaged 50°S–50°N 

SBUV(/2) satellite data using the PWLT model. Using 

Dobson total ozone measurements, Angell and Free (2009) 

find positive trends in the same regions after application of 

5-year running linear trends to the smoothed individual 

station ground-based data. They used 11-year running 

means to minimize the 11-year solar and QBO effects in 

the ozone time series. It should be mentioned however, 

that the positive trend in 50°S–50°N region is largely associated with an ozone increase in the tropical belt related 

to relatively low ozone values there in the mid-1990s and 

relatively high values during the recent solar activity minimum. Loyola et al. (2009b) analyzed the merged 

GOME(/2)+SCIAMACHY data set as well the MOD set 

for the period from June 1995 to April 2009. They report 

a statistically significant positive linear trend between 5°S 

and 30°N for both satellite data sets. All these findings 

seem to contradict previous estimates of the number of 

years required to detect statistically significant ozone 

trend expected from the decline of ODSs (Weatherhead et 

al., 2000; Vyushin et al., 2007). These studies predicted 

that statistically significant ozone trends will be detectable first at southern midlatitudes but that this will not be 

possible earlier than 2015–2020.

Comparison of the PWLT (or linear trend) estimates with results based on the EESC fit, shows that these 

recent positive ozone trends are larger than those expected 

from the decline in ODSs. As mentioned above, knowing 

the EESC decreasing rate after the turning point in the late 

1990s, the corresponding linear term in total ozone regression can be compared to positive trends in PWLT models. 

Figure 2-3 (updated Figures 8 and 9 of Vyushin et al., 

2007) illustrates the ozone zonal trends by PWLT and 

EESC models with the solar and QBO terms applied to the 

MOD set for the periods 1979–2008, with the turning 

point for the PWLT in 1996. Figure 2-3 shows the rate of 

ozone increase based on the EESC fit for the period corresponding to the declining phase of EESC and the estimates for the linear trend after the turning point of the 

PWLT. The gray areas indicate 95% confidence intervals 

for the PWLT estimate. The two trends are fairly similar 

in southern middle and high latitudes, although the uncertainties on the observed trends encompass zero. In northern middle and high latitudes, however, the observed 

linear trend is roughly four times the EESC-predicted 

trend and is actually statistically significant over northern 

middle and low latitudes according to the PWLT estimate 

of the noise. In these regions, the ODS decrease induces a 

positive trend but it is overwhelmed by large dynamically 

driven variations. This result is confirmed by several 

60oS-60oN Area-Weighted Average

1970 1980 1990 2000 2010

Year

270

280

290

300

310

Total Ozone (DU)

Figure 2-1. Quasi-global (60°N–60°S) average of 

total ozone distribution (Dobson units) for the period 

1970–2009 from the BUV/TOMS/SBUV(/2) merged 

ozone data set.
2.9

Stratospheric Ozone and Surface UV

authors, who indicate that the EESC decrease since the 

mid-1990s is not a major contributor to the recent increase in ozone (Reinsel et al., 2005; Dhomse et al., 

2006; Wohltmann et al., 2007; Harris et al., 2008).

On a regional scale, Krzyścin and Borkowski 

(2008) evaluate the ozone trend variability over Europe 

using 10-year blocks of reconstructed total ozone time 

series since 1950. Statistically significant negative trends 

of 1 to 5%/decade are found almost over the whole of 

Europe only in the period 1985–1994. Trends up to −3%/

decade appeared over small areas in earlier periods when 

the anthropogenic forcing on the ozone layer was weak. 

Vigouroux et al. (2008) provide total ozone trends from 

homogenized FTIR measurements in European stations, 

Ground-based dataset

NASA TOMS/OMI/SBUV merged dataset

GOME/SCIAMACHY merged dataset

NOAA SBUV merged dataset

NIWA Assimilated dataset

Figure 2-2. Annual mean area-weighted total ozone deviations from the 1964–1980 means for the latitude 

bands 90°S–90°N, 60°S–60°N, 25°S–25°N, 35°N–60°N, and 35°S–60°S, estimated from different global data 

sets: ground-based (black), NASA TOMS/OMI/SBUV(/2) merged satellite data set (red), National Institute of 

Water and Atmospheric Research (NIWA) assimilated data set (magenta), NOAA SBUV(/2) (blue), and GOME/

SCIAMACHY merged total ozone data (green). Each data set 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. Deviations are expressed as percentages of the groundbased time average for the period 1964–1980. Figure updated from Chapter 3 of WMO, 2007.
2.10

Chapter 2

over the 1995–2004 period. These trends have been updated for the 1995–2009 period for the present Assessment and are summarized in Table 2-1. Because the time 

series are too short to employ the multi-regression models 

described in Section 2.1.2, a bootstrap resampling method 

was used, which allows for non-normally distributed data 

and gives an independent evaluation of the uncertainty in 

the trend value (Gardiner et al., 2008). The total column 

trends are close to zero and not significant at all stations 

except at Kiruna, where the trend is significantly positive.

2.1.4 Update on Ozone Profile Changes

2.1.4.1 Measurements

Ground-Based Measurements

Ozonesondes, Dobson and Brewer spectrometers 

using the Umkehr method, lidars, and microwave instruments provide long-term measurements of ozone vertical 

distribution. Various recent studies have focused on 

assessing the quality and stability of ozonesonde data. Differences ranging from 5 to 10% were found between data 

obtained with sondes produced from different manufacturers or with different sensing solutions (Thompson et al., 

2007; Smit et al., 2007; Kivi et al., 2007; Deshler et al., 

2008; Stübi et al., 2008). The Umkehr method retrieves 

ozone profiles from Dobson and Brewer measurements 

with a vertical resolution of about 5 km in the stratosphere. 

Current Umkehr data are retrieved using the UMK04 algorithm already released for the previous Assessment 

(Petropavlovskikh et al., 2005a, 2005b). Lidars and microwave spectrometers provide range-resolved measurements 

from the lower stratosphere to about 50 km for the lidar and 

higher for the microwave. Lidar measurements are characterized by a higher vertical resolution than microwave 

measurements, but they require clear skies so fewer measurements are obtained. A detailed description of the characteristics of ozonesondes, Umkehr, lidar, and microwave 

measurements in terms of accuracy and vertical resolution 

can be found in WMO (2007) and SPARC (1998).

A new ozone profile data source has been introduced for this Assessment, based on the inversion of FTIR 

measurements (Hase, 2000; Pougatchev et al., 1995). The 

Latitude

60S 30S Eq. 30N 60N

Total Ozone Anomaly (DU)

40−45S 

1980 1985 1990 1995 2000 2005 2010

−50 −30 −10 10 −40 −20 0

40−45N 

1980 1985 1990 1995 2000 2005 2010

60S 30S Eq. 30N 60N

Trend (DU/year) Trend (DU/year)

-0.5 0 0.5 1 1.5 2 2.5

PWLT

PWLT

EESC

EESC

1996-2008

1979-1995

 -2 -1.5 -1 -0.5 0

Figure 2-3. (top) The EESC-based 

linear trend in ozone (Dobson units/

year) calculated for the increasing 

part of EESC (yellow solid circles 

connected by the yellow line) is compared to the first (declining) slope of 

the PWLT fit for the period 1979–1995 

(blue diamonds connected by the solid 

blue line) with 95% confidence intervals (gray shading) for the PWLT fit. 

(middle) The EESC-based linear trend 

in ozone (DU/year) calculated for the 

declining part of EESC (yellow solid 

circles connected by the yellow line) 

is compared to the second (increasing) slope of the PWLT fit for the period 1996–2008 (blue diamonds connected by the solid blue line) with 95% 

confidence intervals (gray shading) for 

the PWLT fit. (bottom) Monthly and 

zonal mean total ozone anomalies 

for the 40°S–45°S and 40°N–45°N 

latitude bands obtained by filtering out 

the seasonal cycle, QBO, and solar 

flux, together with the EESC (yellow) 

and PWLT (blue) fits. Updated from 

Vyushin et al. (2007). The MOD set 

was used.
2.11

Stratospheric Ozone and Surface UV

inversion is based on the optimal estimation method (Rodgers, 2000) and leads to 4–5 degrees of freedom in the 

whole column (Barret et al., 2002). Therefore, in addition to total column ozone, FTIR measurements can provide ozone profile data in 4 vertical layers (approximately 

ground–10 km; 10–18 km; 18–27 km; and 27–42 km). 

The precision of these four ozone partial columns is about 

9.5%, 6.5%, 8.5%, and 6.0%, respectively (Vigouroux et 

al., 2008).

Satellite Measurements and Merged Data Sets

The second Stratospheric Aerosol and Gas Experiment (SAGE II) and Halogen Occultation Experiment 

(HALOE) satellite instruments that provided long and stable global observations of the ozone vertical distribution 

using the solar occultation technique ceased operation in 

2005. Their long observational records, which were used 

in the previous Assessments, cover the periods 1984–2005 

and 1991–2005 respectively. The longest ozone profile 

record based on a single instrument type still in operation is now provided by the series of SBUV and SBUV/2 

instruments in operation since 1978. Retrieved profiles 

however have much lower vertical resolution than SAGE 

II or HALOE. The data are retrieved with the version 8 

algorithm also used for total ozone retrieval (Bhartia et al., 

2004; Flynn, 2007).

Data availability and quality remain the key issues 

in assessing changes in ozone profiles from satellite data. 

Jones et al. (2009) estimated that the smallest detectable 

linear trend in the midlatitude upper stratosphere from 

accurate but sparse SAGE occultation data was 

~2.9%/decade (for the 1979–1997 period), while a trend 

of 1.5%/decade could be detected if the SAGE time series 

are combined with HALOE and much more frequent 

SBUV(/2) nadir observations. The SBUV(/2) record is 

comprised of data from multiple instruments. Biases 

between some of these instruments are comparable with 

long-term ozone changes (e.g., Terao and Logan, 2007; 

Fioletov, 2009) that make the combined record difficult 

to use for the trend estimates.

In the last decade, several satellite instruments providing ozone profile measurements according to various 

measuring techniques have been launched onboard various 

satellite platforms, e.g., Odin (launched in 2001), Envisat 

(2002), SCISAT (2003), Aura (2004), and MetOp-A 

(2006), but their records are too short to contribute to this 

analysis on their own. Evaluation of the impact of the 

stabilization and subsequent decrease of ODS abundances 

in the stratosphere thus requires the merging of multiple 

data sets.

Merging the various satellite ozone profile data sets 

into a single, homogeneous data record suitable for trend 

studies is a challenge since each record is subject to its 

own instrument effects (noise, systematic errors, degradation, aging) and sampling issues (vertical and horizontal 

sampling, resolution, repeat time). Several such data sets 

have been introduced, based on single instrument type, 

such as SBUV and SBUV(/2) (Frith et al., 2004) or multiple instruments, e.g., sondes and solar occultation instruments (Randel and Wu, 2007; Hassler et al., 2008), or satellite instruments using different measurement techniques 

(Jones et al., 2009). McLinden et al. (2009) provide 

another data set based on SAGE and SBUV(/2) data spanning 1979–2005 where drifts in individual SBUV instruments and inter-SBUV biases are corrected using SAGE I 

and II by calculating differences between coincident 

SAGE-SBUV(/2) measurements. In this way the daily, 

near-global coverage of SBUV(/2) is combined with the 

stability and precision of SAGE to provide a homogeneous ozone record. Another approach is used by Jones et 

al. (2009): in order to remove biases between individual 

instrument records, ozone anomalies (deviations from the 

annual cycle) in overlapping periods are compared and 

then corrected for the difference. The data quality issue 

for trend estimates has become particularly important after 

Table 2-1. Annual ozone trends and uncertainties (95% confidence limits), in %/decade, for partial and 

total columns. The measurements at Ny-Ålesund and Kiruna are restricted to the March–September and 

January–November period, respectively. Updated from Vigouroux et al. (2008).

Ozone Trend (%/decade)------------------

FTIR Station Latitude Period 10–18 km 18–27 km 27–42 km Total

Column

Ny-Ålesund 79°N 1995–2009 0.0 ± 5.4 −0.8 ± 2.8 5.7 ± 2.7 0.0 ± 2.6

Kiruna 68°N 1996–2009 −1.0 ± 3.9 4.2 ± 2.3 11.4 ± 2.5 2.8 ± 2.2

Harestua 60°N 1995–2009 −11.4 ± 6.2 4.9 ± 2.3 7.2 ± 2.6 0.3 ± 2.7

Jungfraujoch 47°N 1995–2009 −1.9 ± 3.6 0.4 ± 0.9 0.8 ± 0.9 −0.1 ± 1.1

Izaña 28°N 1999–2009 −3.6 ± 4.2 1.8 ± 1.1 1.0 ± 1.2 0.5 ± 1.1
2.12

Chapter 2

August 2005, when the SAGE II instrument stopped its 

operation, ending its long and stable data record. Because 

both SAGE II and HALOE ceased operations in 2005, few 

trend analyses using satellite ozone profile data were performed since the WMO (2007) report.

2.1.4.2 Ozone Profile Changes

Profile Trends in Altitude and Pressure Coordinates

As discussed in the previous Ozone Assessments, 

care must be taken when comparing trends in ozone derived from data in different geophysical units and/or different vertical coordinate systems (WMO, 2007). This is 

due to simultaneous trends in temperature that impact the 

air density directly and the altitude of a pressure surface 

indirectly. Rosenfield et al. (2005) demonstrated using a 

two-dimensional model that trends in upper stratospheric 

ozone may differ by 1 to 2%/decade depending on the 

units and vertical coordinate of the time series. Terao and 

Logan (2007) show differences up to 4%/decade between 

SAGE trends in altitude and pressure coordinates if 

National Centers for Environmental Prediction (NCEP) 

temperature reanalysis data are used for the conversion. 

An analysis of SBUV(/2) and SAGE ozone time series 

suggests that this difference can be as much as 4% in the 

upper stratosphere if SAGE trends calculated in number 

density versus altitude are compared to SBUV(/2) partial 

pressure versus pressure trends (Figure 2-4). However, if 

SAGE data are converted to the same units as SBUV(/2) 

using temperature data with proper temperature trends 

(Randel et al., 2009) and are adjusted to match SBUV(/2) 

vertical resolution as was done in the SAGE-corrected 

SBUV data set (McLinden et al., 2009), the ozone trends 

derived from SAGE and SBUV(/2) are consistent at the 

1–2%/decade level, roughly that of the trend uncertainties.

Ozone Changes in the Upper Stratosphere

The upper stratosphere (35–45 km) is the region 

where the effects of ODSs are expected to be the easiest to quantify, since the destruction of ozone there is 

mainly due to processes linked to homogeneous chemistry (WMO, 1999). Most studies performed in that region 

show a strong and statistically significant decline (6–8% 

per decade) for the period up to the mid-1990s and a nearzero or slightly positive trend thereafter (e.g., Randel and 

Wu, 2007; Steinbrecht et al., 2009; Jones et al., 2009; 

McLinden et al., 2009).

The ozone variability in the upper stratosphere 

(35–45 km) was examined by Steinbrecht et al. (2009) 

from various satellites and five ground-based lidar stations 

located in northern midlatitudes, tropical latitudes, and 

southern midlatitudes (Figure 2-5). This study extends 

results mentioned in WMO (2007) and includes an evaluation of temperature variability over the same locations. 

The new analysis confirms that the upper stratospheric 

ozone decline apparent from 1979 until the mid-1990s has 

stopped and ozone has stabilized since 1995–1996, depending on the latitude. Tatarov et al. (2009) analyzed 20 

years (1988–2008) of stratospheric ozone and temperature 

20

30

40

50

8

−6

4

−4

−2

−2

−2

−2

−

0 0

Altitude or Pressure Altitude [km] 20

30

40

50

0

−6

−6

−6

−4

−4

4

−4

−4

−4

−2

−2

−2

−2

−2

−2

0

Latitude

−60 −30 0 30 60

20

30

40

50

−

6

−6

−4

−6

0

(a)

(b)

(c)

−6

−6

−4

−4

−4

−4

−4

−2

−2

−2

−2

−8

−6

Figure 2-4. Stratospheric ozone trends as functions 

of latitude and altitude or pressure altitude from various data sources. The trends were estimated using 

regression to an EESC curve and converted to % 

per decade using the variation of EESC with time in 

the 1980s. The plots display trend estimates from 

various published data sets: (a) SAGE I+II (adapted 

from Randel and Wu, 2007), (b) the NASA merged 

SBUV(/2) data set, (c) SAGE-corrected SBUV(/2) 

(McLinden et al., 2009). For panel (c) SAGE I+II data 

are converted onto a pressure using a temperature 

trend from Randel et al. (2009) and then vertically 

smoothed to match the SBUV vertical resolution. 

Shading indicates the trends are significant at the 2σ

level. Panel (a) is plotted in altitude, the remaining 

panels in pressure-altitude. Pressure-altitude is defined as z∗ = −16 log10(p/1000), where p is in hPa and 

z∗ is km. Trend contours are every 2% per decade.
2.13

Stratospheric Ozone and Surface UV

profiles measured by differential absorption lidar (DIAL) 

data at the National Institute for Environmental Studies 

in Tsukuba (36°N, 140°E), Japan. Ozone data in the upper stratosphere exhibit a strong negative trend from 1988 

to 1997 and a statistically insignificant trend after 1998. 

Similar results were obtained for SAGE II coincident data 

over the station.

The lack of a significant ozone trend during the re he lack of a significant ozone trend during the recent period (since 1996) in the upper stratosphere has also 

been found in the Arosa Umkehr data (Zanis et al., 2006). 

Figure 2-5. Ozone anomalies over 

the 1979 to early 2010 period from 

different data sets at five NDACC 

stations. Anomalies are averaged 

over the 35–45 km range. Light 

blue: SBUV(/2)-MOD version 8. 

Dark blue: SAGE I and II version 

6.20. Green: HALOE version 19.0. 

Red: Lidar. Magenta: Microwave. 

Yellow: SCIAMACHY IUP–Bremen version 2.0. Violet: GOMOS 

ESA IPF 5.00. Black: Average of 

all available instruments. Gray 

underlay: CCMVal model simulations, 24-month running average 

±2 standard deviations. Observed 

data are smoothed by a five-month 

running mean. Lidar and microwave data are station means; all 

other data are zonal means. The 

thin black lines at the top and bottom show negative 10 hPa zonal 

wind at the equator as a proxy for 

the QBO, and 10.7 cm solar flux as 

a proxy for the 11-year solar cycle, 

respectively. The thin magenta line 

near the bottom shows inverted 

effective stratospheric chlorine as 

a proxy for ozone destruction by 

chlorine (ESC, 4 years mean age, 

2 years spectral width, no bromine; 

see Newman et al., 2006). Updated from Steinbrecht et al. (2009).

1980 1985 1990 1995 2000 2005 2010

5

0

−5

35 to 45 km Ozone Anomaly [%]

5

0

−5

5

0

−5

5

0

−5

5

0

−5

−10

−15

−20

Year
2.14

Chapter 2

In contrast, analyses of the homogenized Umkehr record 

for Belsk yielded a statistically significant upward trend in 

the upper stratosphere in the period 1996–2007 but not a 

decisive trend at other altitudes (Krzyścin and RajewskaWiech, 2009). From FTIR measurements, trends in the 

upper stratospheric layer of the ozone profile retrieval 

(27–42 km) range from an unsignificant 0.8% per decade 

trend at Jungfraujoch (47°N) to a significant positive trend 

of up to 11% per decade at high latitude stations (see Table 2-1). Trend results from relatively short time series 

in the Arctic should be considered with caution due to the 

high variability of ozone in this region, especially during 

wintertime.

Jones et al. (2009) provide a global estimate of 

ozone trends from the average of various satellite ozone 

anomaly records. Using the PWLT statistical model with a 

turning point in 1997, they find that the largest statistically 

significant ozone declines in 1979–1997 are found in the 

midlatitude regions between 35 and 45 km altitude in both 

hemispheres, with trend values of approximately −7%/ 

decade. For the period 1997 to 2008, they derive trends of 

1.4 and 0.8%/decade in the NH and SH respectively, but 

these are not statistically significant (see Table 2-2).

Ozone Changes in the Lower Stratosphere

The lower stratosphere between 20 and 25 km over 

middle latitudes is another region where a statistically significant decline of about 4 to 5%/decade (or 7–8% total decline) occurred between 1979 and the mid-1990s, followed 

by stabilization or a slight (2–3%) ozone increase thereafter.

Angell and Free (2009) analyzed long-term ozone 

profile time series from four Northern Hemisphere Dobson 

Umkehr and 9 ozonesonde stations for trend analysis between 1970 and 2007. The 5-year trends were derived 

from the 11-year running mean of the time series to minimize the impact of the 11-year solar cycle and QBO signals in the data. Both Umkehr and sonde data showed that 

nearly half of the increase in north temperate total-ozone 

trend between 1989 and 2000 was due to an increase in the 

10–19 km layer in the lower stratosphere, with the troposphere contributing only about 5% of the change. Nonsignificant positive ozone trends at the end of the record in 

2000 were found at four Umkehr layers in the middle and 

high stratosphere, as well as between 10 and 32 km altitude 

in sonde data.

Murata et al. (2009) could not detect any trend from 

a 14-year data set of ozone profiles measured with a balloonborne optical ozone sensor beginning in 1994 at 

Sanriku, Japan. This lack of trend was attributed to the 

leveling off of ODSs in the stratosphere. The extension of 

the FTIR trend analysis up to 2009 shows no significant 

trend at the midlatitudes station for the 18–27 km layer 

(Table 2-1). Similarly, the global trend analysis of Jones 

et al. (2009) shows no significant trend for the 20–25 km 

altitude range in the NH and SH midlatitudes for the period 

1997–2008.

Figures 2-6a and 2-6b show the temporal evolution 

of deseasonalized ozone monthly means in three pressure 

ranges (upper, lower, and lowermost stratosphere) based 

on ozonesondes, Umkehr, and SBUV(/2) observations 

over Europe and Lauder in the SH, respectively (adapted 

from Terao and Logan, 2007). The various time series 

show very similar interannual variation, although some 

biases are apparent between the measurements. In the upper and lower stratosphere, ozone levels have stabilized 

after a decrease from the early 1980s to the mid-1990s. In 

the lower stratosphere, the decrease was more pronounced 

over Europe than in the SH. In the lowermost stratosphere, no significant long-term variation is observed at 

Table 2-2. Average ozone trends and uncertainties (95% confidence limits) in %/decade in the lower 

and upper stratosphere in the NH and SH midlatitudes, from various data sources for the period 1996–

2008. The ozonesondes and Umkehr results correspond to the PWLT trends in Figure 2-7. The FTIR results 

are for the Jungfraujoch station only, for the 1995–2009 period, and correspond to respectively the 18–27 km 

and the 27–42 km altitude ranges for the lower and upper stratosphere.

Ozone Trend

30°S–60°S

(%/decade)

Ozone Trend

30°N–60°N

(%/decade)

Data Source 20–25 km 35–45 km 20–25 km 35–45 km

Satellite (from Jones et al., 2009) −1.0 ± 2.0 0.8 ± 2.1 0.2 ± 1.9 1.4 ± 2.3

Umkehr 0.2 ± 2.6 2.0 ± 1.5 3.2 ± 2.1 1.5 ± 1.3

Ozonesondes 1.5 ± 0.6

FTIR (updated from Vigouroux et al., 

2008) 0.4 ± 0.9 0.8 ± 0.9
2.15

Stratospheric Ozone and Surface UV

1.6−6.4 hPa

Smoothed Deseasonalized Monthly Means

−10

−5

0

5

10−63 hPa

−40

−20

0

20

Sondes

SBUV

Umkehr

100−250 hPa

−40

−20

0

20

Year

Ozone Anomaly (DU)

(a) Europe

1980 1985 1990 1995 2000 2005 2010

1.6−6.4 hPa

Smoothed Deseasonalized Monthly Means

−10

−5

0

5

10−63 hPa

−40

−20

0

20

Sondes

SBUV

Umkehr

100−250 hPa

−40

−20

0

20

Year

Ozone Anomaly (DU)

1980 1985 1990 1995 2000 2005 2010

(b) Lauder

Figure 2-6. (a) Monthly ozone anomalies in Dobson units for Europe as measured by ozonesondes (black line), 

SBUV(/2) (red line), and Umkehr (blue 

line) at three pressure layers. The 

monthly anomalies were computed as 

the difference between a given monthly 

mean and the average of monthly 

means for 1979–1987 for each data 

set. The average of the monthly-mean 

anomalies for 1979–1981 was then 

subtracted from each anomaly time series to set the zero level in each panel. 

A three-month running mean was applied to the anomalies. The SBUV(/2) 

data were selected within a grid box of 

45°N–55°N and 10°W–30°E. The 

ozonesonde data are the average of 

measurements at three European stations: Hohenpeissenberg, Payerne, 

and Uccle. The Umkehr data are from 

Arosa, Belsk, and Haute-Provence 

Observatory. The sonde and SBUV(/2) 

analysis is updated from Terao and 

Logan (2007). (b) Same as for (a) but 

for the Southern Hemisphere. The 

sonde data are from Lauder, New 

Zealand. The SBUV(/2) data were selected within a grid box of 40°S–50°S 

and 150°E–170°W. The monthly anomalies were computed using the 

monthly means for 1987–1991.
2.16

Chapter 2

either location over the whole period, but higher shortterm variability was seen during the nineties in Europe.

The vertical profile of ozone trends computed from 

SBUV(/2), Umkehr, and ozonesonde data over Northern 

midlatitudes stations is displayed in Figure 2-7 for both the 

increasing and decreasing periods of EESC (e.g., 1979–

1995 and 1996–2008). The trends were derived using 

EESC as a regression term accounting for the variation of 

mean age of air as a function of altitude (see Waugh and 

Hall (2000) for a discussion of age of air and its spatial 

dependence). In the case of ozonesondes, trends were 

computed as the average of trends derived for nine northern 

midlatitude stations, as in the previous Ozone Assessment 

(Chapter 3). For Umkehr, the trend was derived from the 

average of ozone anomalies at four northern midlatitude 

stations and for SBUV(/2), the 40°N–50°N zonal mean 

data were used. Piecewise linear trends with inflection 

point in January 1996 derived from ozonesonde and 

Umekhr data are also represented in the figure. As shown 

in WMO (2007), ozone trends during the first increasing 

period of EESC display two maxima in the upper and lower 

stratosphere, reaching −5 to −7%/decade and −4 to −5%/

decade respectively (total decline of about 10% and 7% 

respectively), with generally good agreement between the 

various observations, except for Umkehr in the lower 

stratosphere. In both cases, the EESC and PWLT trend 

models give similar results for this period. For the decreasing EESC period, positive ozone trends are derived. In the 

upper and lower stratosphere, EESC and PWLT models 

provide similar trends of about 2%/decade. The PWLT 

trends are significant in the lower stratosphere and barely 

significant in the upper stratosphere. These results indicate 

that while the decrease of ODSs is indeed causing an 

increase of ozone over these midlatitude stations, this in-

−10 −5 0 5 10

1

10

100

400

200

40

20

4

2

0.4

0.2

Pressure (hPa)

Trend (%/decade)

1979−1995

−10 −5 0 5 10

10

14

18

22

26

30

34

38

42

46

50

54

58

62

Altitude (km)

EESC SBUV

EESC Umkehr

PW Umkehr

EESC sondes

PW sondes

−10 −5 0 5 10

1

10

100

400

200

40

20

4

2

0.4

0.2

Pressure (hPa)

Trend (%/decade)

1996−2008

−10 −5 0 5 10

10

14

18

22

26

30

34

38

42

46

50

54

58

62

Altitude (km)

EESC SBUV

EESC Umkehr

PW Umkehr

EESC sondes

PW sondes

(a) (b)

Figure 2-7. Vertical profile of ozone trends over Northern midlatitudes estimated from ozonesondes, Umkehr, 

and SBUV(/2) measurements for the period 1979–2008. The trends were estimated using regression to an 

EESC curve and converted to % per decade using the variation of EESC with time from 1979 to 1995 in panel 

(a) and from 1996 to 2008 in panel (b). Piecewise linear trends with inflection point in January 1996 derived from 

ozonesonde and Umkehr data are also shown. The trend models also include QBO and solar cycle terms. The 

sonde results are an average of trends for Churchill, Goose Bay, Boulder, Wallops Island, Hohenpeissenberg, 

Payerne, Uccle, Sapporo, and Tateno, along with two standard errors of the nine trends. The Umkehr trends 

were derived from averaged ozone anomalies at Belsk, Arosa, OHP, and Boulder. For SBUV(/2), the 40°N–

50°N zonal mean data were used. The altitude scale is from the standard atmosphere. The error bars correspond to 95% confidence interval.
2.17

Stratospheric Ozone and Surface UV

crease is still barely significant, especially in the upper 

stratosphere where trends derived from PWLT and EESC 

models are expected to show the best agreement. In contrast, in the lowermost stratosphere, EESC and PWLT 

trends derived from sondes data differ significantly, with 

large positive trend values in the latter case, suggesting that 

the ozone increase is due to factors other than chlorine decline, for example dynamical processes (see Section 2.4).

Table 2-2 summarizes the average trends found 

from various data sources using the PWLT model in the 

NH and SH midlatitudes in the lower (20–25 km) and 

upper (35–45 km) stratosphere. Most results show positive ozone trends (1–3% increase) since 1996 in the various regions. These trends are significant at some locations 

(e.g., over Northern midlatitudes) but the results from 

global satellite data are still not significant at the 95% 

confidence level (Jones et al., 2009).

Northern Hemisphere midlatitude (35°N–60°N) 

ozone between 12 and 15 km decreased by about 9% between 1979 and 1995, and increased by about 6% between 

1996 and 2009 (Figure 2-7). The increase since the mid1990s is larger than the changes expected from the decline 

in ODS abundances.

2.2 POLAR OZONE

Chapter 4 of WMO (2007) (“Polar Ozone: Past and 

Present”, Newman and Rex et al., 2007) builds upon the 

sequence of polar ozone chapters in the WMO Assessment 

series. The present discussion updates WMO (2007), 

highlighting changes over the past four years. Discussion 

of polar ozone recovery is found in Chapter 3 of this 

Assessment, and interactions of polar chemistry and 

climate are found in Chapter 4.

2.2.1 State of Science in 2006

The discovery of the Antarctic ozone hole by 

Farman et al. (1985) prompted considerable effort to develop the scientific basis necessary to model and predict 

polar ozone loss. As noted in the previous chapter, the 

stratospheric chlorine burden reached its peak in the late 

1990s and has since begun to decrease. During the period 

of increasing chlorine concentrations, the springtime polar ozone values decreased in both hemispheres. Consistently low values in springtime ozone have been observed 

since the mid-1980s in the Southern Hemisphere. A 

unique dynamical situation, the first major sudden stratospheric warming in the Southern Hemisphere, led to the 

anomalously high ozone levels in 2002; this situation is 

discussed in detail in Chapter 4 of WMO (2007). The 

Arctic polar ozone loss is much more variable, depending not just on the stratospheric chlorine level but also on 

whether or not the winter is cold enough and of sufficient 

length for chlorine-catalyzed ozone loss to occur. Because 

chlorofluorocarbons are long-lived, atmospheric chlorine 

loading is declining slowly.

The springtime averages of total ozone poleward of 

63° latitude in the Arctic and Antarctic are shown in Figure 2-8 (an update of Figure 4-7 from WMO, 2007). Interannual variability in polar stratospheric ozone abundance 

and chemistry is driven by variability in temperature and 

transport due to year-to-year differences in dynamics. 

The horizontal gray lines in Figure 2-8 are the averages 

of ozone values obtained between 1970 and 1982, and the 

shading emphasizes the differences between these averages and subsequent years.

WMO (2007) outlined the processes important 

to polar ozone loss. The rate-limiting step of the dominant cycle for polar ozone destruction is the photolysis of 

chlorine peroxide (ClOOCl, also known as the chlorine 

monoxide dimer). Since the previous Assessment, a laboratory study suggesting a much lower photolysis rate of 

ClOOCl than previously recommended (Pope et al., 2007) 

prompted a number of subsequent laboratory studies on 

the subject. The implications of the laboratory studies 

since WMO (2007) for the interpretation of the observations of chlorine monoxide (ClO) and ClOOCl and for the 

assessment of the uncertainty in computation of ozone loss 

rates are discussed below along with other updates to the 

photochemical data in Section 2.2.2.

WMO (2007) concluded that Antarctic ozone loss 

had stabilized over the time period 1995–2005, with higher ozone levels in 2002 and 2004 that were dynamically 

63o-90o Total Ozone Average

1970 1980 1990 2000 2010

Year

200

250

300

350

400

450

500

Total Ozone (DU)

BUV

EPTOMS_V8C

METEOR3_V8

NIMBUS7_V8

OMI_V8F

Merged

SH October

NH March

Figure 2-8. Total ozone average (Dobson units) 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 for the years prior to 

1983 in March for the NH and in October in the SH. 

Updated from Figure 4-7, WMO (2007).
2.18

Chapter 2

driven and not related to reductions in the stratospheric 

halogen load. For the Arctic it was concluded that for the 

coldest Arctic winters, the volume of air cold enough to 

support polar stratospheric clouds (PSCs) had increased 

significantly since the late 1960s. Arctic spring total 

ozone was reported to be lower than in the 1980s and was 

also noted to be highly variable from year to year depending on dynamical conditions. This is discussed further in 

Sections 2.2.4 and 2.2.5.

Transport and mixing both affect high-latitude winter ozone, making it challenging to diagnose the chemical loss rate from observations. WMO (2003) presented 

an overview of various methods that have been used to 

separate these effects, mainly for the Arctic winter. WMO 

(2007) included a comparison of the methods, focusing on 

the 2002/2003 winter. This colder-than-average winter 

was chosen because aircraft and ground field campaigns 

provided the data needed to assess our understanding of 

polar ozone loss, particularly for the large solar zenith 

angle-conditions of early winter. WMO (2007) noted that 

the various methods had been refined since WMO (2003) 

and produced consistent results.

WMO (2007) included evidence that nitric acid trihydrate (NAT) polar stratospheric cloud particles nucleate 

above the ice frost point and are widespread. Improved 

NAT mechanisms in chemistry-transport models (CTMs) 

produce more realistic denitrification, but fail to capture 

observed interannual variability for the northern winters. 

Issues concerning polar stratospheric clouds and their 

representation in models are discussed in Section 2.2.3. 

This section emphasizes new measurements from the 

Cloud-Aerosol Lidar and Infrared Pathfinder Satellite 

Observation (CALIPSO) satellite launched in 2006.

2.2.2 Polar Ozone Chemistry

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 → Cl + ClOO (1b)

ClOO + M → Cl + O2 + M (1c)

2 [Cl + O3 → ClO + O2] (1d)

Net: 2O3 → 3O2

Cycle 2

BrO + ClO + hν → Br + Cl + O2 (2a)

Br + O3 → BrO + O2 (2b)

Cl + O3 → ClO + O2 (2c)

Net: 2O3 → 3O2

Loss of ClOOCl by thermal decomposition

ClOOCl + M → ClO + ClO + M (1e)

or chemical processes that recycle ClO without the formation of O2 do not cause ozone depletion. Production of 

Br + OClO by BrO + ClO also leads to a null cycle. Small 

contributions to polar ozone loss occur due to cycles involving the reactions ClO + O and ClO + HO2.

Since WMO (2007), attention has focused on resolving uncertainties in the photolysis cross sections and 

quantum yields for the ClO dimer, ClOOCl. Pope et al. 

(2007) reported a ClOOCl absorption spectrum with cross 

sections at wavelengths between 300 and 350 nm much 

lower than recommended and than reported in prior studies 

(e.g., Sander et al., 2006 (referred to in this chapter as JPL 

06-2) and references therein), challenging the fundamental 

understanding of polar ozone depletion (i.e., Schiermeier, 

2007; von Hobe, 2007). Photochemical models using 

the cross sections reported in Pope et al. (2007), with all 

other kinetic parameters from JPL 06-2, underestimate the 

ozone loss rate (von Hobe et al., 2007) as well as observed 

abundances of ClO (von Hobe et al., 2007; Santee et al., 

2008; Schofield et al., 2008). A workshop entitled “The 

Role of Halogen Chemistry in Polar Stratospheric Ozone 

Depletion” was convened in summer 2008 to assess the 

fundamental understanding of polar ozone depletion in 

light of the Pope et al. (2007) measurements. The following material incorporates findings from this workshop 

(SPARC, 2009).

The small ClOOCl cross sections reported by Pope 

et al. (2007) have been contradicted by all subsequent laboratory studies, as detailed below. There is now consensus 

in the community that photolysis of ClOOCl occurs much 

faster than implied by Pope et al. (2007). Further, no credible “missing chemical process” has been proposed that 

can be included in models that use the Pope et al. (2007) 

cross section values to adequately account for observed 

levels of [ClO] (or in some cases [ClO] and [ClOOCl]; 

brackets denote concentration of the species) as well as 

ozone loss derived from observations. The fundamental 

understanding that polar ozone depletion is caused primarily by reactions involving ClO + ClO, with significant contribution from reactions involving BrO + ClO, has been 

strengthened since WMO (2007), based on new laboratory 
2.19

Stratospheric Ozone and Surface UV

studies and analyses of field observations. The renewed 

focus on ClOOCl photolysis has improved knowledge of 

the UV absorption spectrum and cross sections such that a 

better quantitative understanding of polar ozone depletion 

and the relation to halogens has been achieved in the past 

few years. Present uncertainties in the ClOOCl absorption 

cross section, as recommended in Sander et al., 2009 (referred to in this chapter as JPL 09-31), are consistent with 

our fundamental understanding that halogens cause depletion of polar ozone. The following discussion supports the 

statements given above.

2.2.2.1 Laboratory Studies of the ClOOCl UV 

Absorption Spectrum

Figure 2-9 shows the ClOOCl ultraviolet-visible 

(UV/vis) absorption spectrum, σClOOCl(λ), reported in the 

laboratory studies of Burkholder et al. (1990) and Pope 

et al. (2007) combined with the spectra recommended 

by Atkinson et al., 2007 (hereafter referenced as IUPAC, 

2007) (based on Huder and DeMore, 1995) and JPL 06-2 

(Sander et al., 2006) (based on Cox and Hayman, 1988; 

DeMore and Tschuikow-Roux, 1990; Permien et al., 1988; 

and Burkholder et al., 1990), which represents the state of 

knowledge in 2007. The laboratory measurements published after 2007 and the recent NASA JPL Panel for Data 

Evaluation recommendation (JPL 09-31, Sander et al., 

2009) are also shown in Figure 2-9. The JPL Panel revised 

the recommendations for a number of kinetic parameters 

for processes important in modeling stratospheric ozone 

depletion. The recommended ClOOCl cross sections in 

JPL 09-31 were unchanged from JPL 06-2, since the recommendation was formulated prior to the appearance of 

several more recent studies, but the recommended uncertainty limits were decreased. The estimated uncertainty 

limits from JPL 06-2 and JPL 09-31 are included in Figure 

2-9, which also illustrates the most critical wavelength region for the calculated photolysis rate constant of ClOOCl 

(J1b) in polar regions. The uncertainty in the calculated 

photolysis rate constant stems primarily from the uncertainty in the absorption cross sections and increases considerably for wavelengths (λ) greater than 350 nm, where 

there is more limited experimental data. At the time of 

WMO (2007), only Burkholder et al. (1990) and DeMore 

and Tschiukow-Roux (1990) had reported measured cross 

section data for λ > 360 nm, while several other studies 

provided extrapolated values. The recent studies of von 

Hobe et al. (2009) and Papanastasiou et al. (2009) confirm 

substantial contributions to J1b from this spectral region. 

The uncertainty in the calculated photolysis rate constant 

is much less using the JPL 09-31 recommendation than 

using the JPL 06-2 recommendation.

Pope et al. (2007) reported a ClOOCl absorption 

spectrum at 195 K that was normalized at 245 nm to the 

absolute cross section value recommended in JPL 06-2. 

The ClOOCl spectrum was obtained using a fitting procedure of measured spectra, which contained significant 

contributions from molecular chlorine (Cl2). For the photolytically active region λ > 300 nm, their inferred cross 

sections are significantly lower than reported in all earlier studies, prompting investigations of the impact of the 

cross sections reported by Pope et al. (2007) on observations of stratospheric ClO and polar ozone loss (von Hobe 

et al., 2007; Santee et al., 2008; Schofield et al., 2008; see 

also Section 2.2.2.2).

Laboratory studies of σClOOCl(λ) (Papanastasiou et 

al., 2009), the relative absorption spectrum (von Hobe et 

al., 2009), and the product of the cross section and quantum yield, σClOOCl(λ)Φ(λ) (Chen et al., 2009; Lien et al., 

2009; Jin et al., 2010; Wilmouth et al., 2009) have been 

published subsequent to Pope et al. (2007). These studies 

used complementary experimental techniques designed to 

reduce the uncertainty in the photochemistry of ClOOCl 

and to address the discrepancy between Pope et al. (2007) 

and earlier studies. The more recent laboratory studies 

either included a method for quantification of Cl2 (von 

Hobe et al., 2009; Papanastasiou et al., 2009; Wilmouth 

et al., 2009) or used a mass-selected detection method 

independent of interference from Cl2 and other impurities (Chen et al., 2009; Lien et al., 2009; Jin et al., 2010). 

Indeed, one of the salient points from the Stratospheric 

Processes and their Role in Climate (SPARC) initiative 

was the need for laboratory studies to address the sensitivity of σClOOCl(λ) to the presence of Cl2 (e.g., Figure 2.2 of 

SPARC, 2009). All of these studies indicate that ClOOCl 

photolyzes much more rapidly than suggested by Pope et 

al. (2007). These studies have led to a reduction in the 

overall uncertainty in σClOOCl(λ) relative to the state of 

knowledge at the time of WMO (2007). However, only 

one of the new gas-phase studies (Papanastasiou et al., 

2009) extended cross section measurements to λ > 352 

nm and, thus, a higher level of uncertainty remains for the 

longer wavelengths.

von Hobe et al. (2009) measured a ClOOCl UV 

absorption spectrum, for λ between 220 and 400 nm, in 

a neon (Ne) matrix at ~10 K. Absolute cross sections 

were not measured, but cross sections were determined by 

scaling the measured spectrum to the peak value of the 

gas-phase cross section reported by JPL 06-2. When interpreting the von Hobe et al. (2009) spectrum it needs to 

be considered that a spectrum measured in a Ne matrix at 

10 K is not directly comparable to a gas-phase spectrum 

at atmospheric temperatures (>190 K). For the photolytically active region (λ> 300 nm), the matrix cross sections 

are significantly greater than the Pope et al. (2007) cross 

sections when both spectra are normalized to the same 

peak cross section. The cross sections reported by von 

Hobe et al. (2009) are somewhat less than the JPL 06-2 
2.20

Chapter 2

recommendation in the range 300–350 nm (Figure 2-9). 

However, significantly greater cross sections would result 

by scaling the von Hobe et al. (2009) absorption spectrum 

to a larger value of the peak cross section, as suggested by 

new observations (see below).

Chen et al. (2009) (308 and 351 nm), Lien et al. 

(2009) (248 and 266 nm), Jin et al. (2010) (330 nm), and 

Wilmouth et al. (2009) (248, 308, and 352 nm) have reported values of σClOOCl(λ)Φ(λ) at the wavelengths given 

in parentheses (Figure 2-9). The first three of these studies measured the loss of ClOOCl following photolysis in 

an effusive molecular beam. Assuming Φ(λ) = 1, the derived cross sections are greater than those given in JPL 

09-31 and agree most closely with those of Burkholder 

et al. (1990) and Papanastasiou et al. (2009). These three 

studies also report a weak temperature dependence to the 

200 250 300 350 400 450

Wavelength (nm)

 JPL 09-31

 Chen et al. (2009)

 Lien et al. (2009)

 Papanastasiou et al. (2009)

 von Hobe et al. (2009)

 Wilmouth et al. (2009)

 Jin et al. (2010) 4

3

2

1

0

J(λ) (1

0-5 s-1 nm-1)

200 250 300 350 400 450

Wavelength (nm)

10-21

10-20

10-19

10-18

10-17

Cross Section (c

m2 molecule-1)

 JPL 06-2

 Burkholder et al. (1990)

 IUPAC (2007)

 Pope et al. (2007)

240 280

0

ClOOCl UV/vis Spectrum

State of Knowledge: 2007

New Studies Since 2007

and Revised JPL Recommendation

1 × 10-17

Figure 2-9. The left column summarizes the ClOOCl photochemical state of knowledge (UV/vis absorption 

spectra, σ(λ), and representative photolysis rates, J(λ), for the polar region) in 2007 and the right column gives 

the advancements since 2007 (as indicated in the legends). The cross sections for Chen et al. (2009), Lien 

et al. (2009), and Jin et al. (2010) at 200 K and Wilmouth et al. (2009) at 240 K were calculated from their reported values of σ(λ)Φ(λ) assuming Φ(λ) = 1. The Pope et al. (2007) and Papanastasiou et al. (2009) spectra 

were recorded at ~200 K. Included for comparison is the von Hobe et al. (2009) spectrum recorded in a Ne 

matrix at ~10 K. The dashed lines represent wavelength regions where extrapolated cross section data rather 

than measured values were reported. The error bars shown were taken from the individual studies. The gray 

shaded regions in the upper panels are the error limits in σ(λ) reported in JPL 06-2 and JPL 09-31. J(λ) values 

were obtained for a solar zenith angle of 86 degrees at 20 km and the gray shaded regions are derived from 

the estimated uncertainties in σ(λ) from JPL 06-2 and JPL 09-31. The hashed region was derived from the 

reported 2σ uncertainty in the Papanastasiou et al. (2009) absorption cross section data. Inset in the top right 

panel shows the absorption spectrum of ClOOCl in the wavelength region 220–280 nm. The choice of a linear 

y-axis allows a better visual assessment of the uncertainties in the peak absorption.
2.21

Stratospheric Ozone and Surface UV

cross sections at 308, 330, and 352 nm. Wilmouth et al. 

(2009) used a discharge flow apparatus combined with Cl 

atom resonance fluorescence detection of photolysis products. Their reported values of σClOOCl(352 nm) at 240 K, 

assuming Φ(352 nm) = 1, are in good agreement with JPL 

09-31. These values are a factor of 12 greater than Pope et 

al. (2007), a factor of 2.3 greater than the IUPAC (2007) 

recommendation, and ~30% lower than the value reported 

by Chen et al. (2009). The cross section measurements at 

several wavelengths from the effusive beam studies (e.g., 

Figure 5 of Lien et al., 2009) are in agreement with the 

spectrum reported by von Hobe et al. (2009).

Papanastasiou et al. (2009) reported σClOOCl(λ) from 

200 to 420 nm at ~200 K. The ClOOCl cross sections obtained at λ > 300 nm are in agreement with values reported 

by Burkholder et al. (1990). Their measurements agree

with the results of Lin and co-workers and Wilmouth et al. 

(2009), assuming Φ(λ) = 1, within the combined estimated 

measurement uncertainties (Figure 2-9). Their σClOOCl(λ > 

300 nm) values are significantly greater than those of Pope 

et al. (2007) and are also somewhat greater than the values 

recommended in JPL 09-31.

The studies of Papanastasiou et al. (2009) and 

Lien et al. (2009) report ClOOCl absorption cross sections near the peak of the spectrum of 7.6+

−

0

0

.

.

8

5 × 10−18 cm2

molecule−1 at 244.25 nm and (8.85 ± 0.42) × 10−18 cm2

molecule−1 at 248.4 nm, respectively. These are greater 

than the JPL 09-31 recommended cross section of 6.4 × 

10−18 cm2 molecule−1 at 244 nm. Wilmouth et al. (2009) 

report a value of (6.6 ± 1.0) × 10−18 cm2 molecule−1 at 

248 nm, in close agreement with JPL 09-31. The spread 

in recent measurements of the peak absorption cross 

section is greater than the estimated uncertainty given 

by JPL 06-2 and many prior evaluations. Furthermore, 

the uncertainty limit of the Lien et al. (2009) peak cross 

section does not overlap with the uncertainty limits reported by Wilmouth et al. (2009) and Papanastasiou et 

al. (2009) (Figure 2-9). Scaling the absorption spectrum for ClOOCl to the peak cross section reported by 

Lien et al. (2009) would result in larger values for the 

ClOOCl cross section, leading to greater photolysis rates 

for ClOOCl as reported by various prior studies. However, even if the Pope et al. (2007) spectrum was scaled 

to the maximum cross section compatible with current 

measurements and the reported error bars, the resulting 

cross sections for λ > 300 nm would still be significantly 

less than reported in all other studies.

Papanastasiou et al. (2009) also reported estimated 

2σ uncertainties in the photolysis rate constant of ClOOCl, 

J1b, based on their cross section data (Figure 2-9). The 

uncertainty is a function of solar zenith angle (SZA) (i.e., 

the wavelength dependence of the ClOOCl spectrum) 

and was estimated to be +40%/−15%, +50%/−20%, and 

+80%/−25% at SZAs of 80°, 86°, and 90°, respectively. 

Better quantification of σClOOCl for λ > 350 nm would reduce this level of uncertainty.

There is a consensus in the community that photolysis of ClOOCl occurs much faster than implied by the 

measured spectrum and the determined cross sections reported by Pope et al. (2007). The weight of new laboratory evidence suggests that values of the ClOOCl absorption 

cross section (σClOOCl) in the atmospherically important region of wavelengths greater than 300 nm reported by Pope 

et al. (2007) are erroneous due to overcorrection of their 

measured spectra for the contribution of a Cl2 impurity. 

The study by von Hobe et al. (2009) showed that ClOOCl 

exhibits an absorption feature similar to the spectral shape 

of the Cl2. Papanastasiou et al. (2009) demonstrated that 

the spectral interference by Cl2 in the Pope et al. (2007) 

experiment led to an underestimate of σClOOCl at λ > 300 

nm. The JPL 09-31 recommended value of Φ (λ > 308 

nm) is 0.9, although only limited experimental studies are 

available (Moore et al., 1999; Plenge et al., 2004). The 

understanding that polar ozone depletion is caused by 

reactions involving halogens has been reaffirmed by the 

numerous laboratory studies conducted since the publication of the paper by Pope et al. (2007).

2.2.2.2 Field Observations of Chlorine

Partitioning

A large number of field studies over the past several 

decades have focused on the quantitative understanding of 

the partitioning of ClO and ClOOCl. The chemistry linking ClO and ClOOCl is thought to be especially simple. 

During daytime when temperatures are low enough that 

loss of ClOOCl occurs mainly by photolysis, the ratio 

[ClO]2/[ClOOCl] essentially equals J1b/k1a (e.g., Stimpfle 

et al., 2004) (brackets denote concentration of the species). During night after hours of darkness, when loss of 

ClOOCl occurs exclusively by thermal decomposition, 

this ratio equals k1e/k1a, which is the equilibrium constant 

(KEQ) between ClO and ClOOCl. Since the rate of ozone 

loss by Cycle 1 is controlled by the parameters J1b and k1a, 

comparisons of measured and modeled daytime values of 

[ClO] and [ClOOCl] provide a quantitative measure of the 

speed of this cycle in the atmosphere. Thermal decomposition of ClOOCl completes a null cycle. Precise knowledge of KEQ and an accurate measurement of nighttime 

[ClO] enable [ClOx] ([ClOx] = [ClO] + 2×[ClOOCl]) to be 

estimated in a manner that is independent of σClOOCl.

Stimpfle et al. (2004) introduced a quantitative 

basis for comparison of modeled [ClO]2/[ClOOCl] to the 

measured value of this quantity at various SZAs during 

daytime, termed β, to quantify how well models represent 

the true value of J1b/k1a. The notion that β represents the 

value of J1b/k1a assumes that the partitioning of ClO and 

ClOOCl is dominated by the self-reaction of ClO and the 
2.22

Chapter 2

photolysis of ClOOCl (i.e., the temperature is low enough 

that thermal dissociation of ClOOCl is much slower than 

photolysis of ClOOCl). The Stimpfle et al. (2004) data 

were obtained at sufficiently low temperature that this 

assumption is valid given known chemistry. This data set 

is also notable for having achieved quantitative closure of 

the chlorine budget (Wilmouth et al., 2006).

Figure 2-10 (left) shows β, as a function of SZA, for 

the laboratory studies and recommendations of σClOOCl(λ) 

that existed in 2007 and provided enough spectral information for calculation of J1b. The black dotted lines indicate the 1σ uncertainty in β based on measured [ClO] 

and [ClOOCl], whereas the error bars indicate the uncertainty in the modeled value of β (see caption). The β ratio 

indicates that the partitioning of ClO and ClOOCl is not 

consistent with J1b/k1a based on the Pope et al. (2007) measurement of σClOOCl(λ). The slight change in k1a in JPL 

09-31 does not affect this or any other conclusion of this 

Figure 2-10. Analysis of β versus SZA, for all daytime measurements obtained during the SAGE III Ozone Loss 

and Validation Experiment (SOLVE), where β = {([ClO]2 / [ClOOCl])MODEL} / {([ClO]2 / [ClOOCl])OBSERVATION}. The 

left panel shows results for J1b using values of σClOOCl(λ) available as of 2007 (same colors and studies as used 

in Figure 2-9). The right-hand side shows results for J1b using values of σClOOCl(λ) available after 2007. The 

black dotted lines depict the ±25% uncertainty (1σ) in β attributable to uncertainties in observations of [ClO] and 

[ClOOCl]. The thick colored error bars denote the standard deviation about the mean for all of the individual 

determinations of β within a particular SZA bin. The thin error bar represents total uncertainty in the modeled 

component of β, found by combining the JPL 06-2 (left panel) or JPL 09-31 (right panel) uncertainty in k1a in a 

root-sum-square fashion with the standard deviations. If the thin error bar falls within the range of the dotted 

lines, this is agreement of model and measurements to within combined 1σ uncertainties. The black dashed line 

on the right hand side depicts results of an illustrative calculation, the scaling of the JPL 09-31 value of σClOOCl(λ), 

which minimizes the distance between the scaled cross section and determinations of the peak cross section 

near 244 nm reported by Lien et al. (2009), Papanastasiou et al. (2009), and Wilmouth et al. (2009). The bottom 

panel compares J1b as a function of SZA, for the altitude and surface albedo of the observations, for the various 

values of σClOOCl(λ) shown in the upper panels. After Stimpfle et al. (2004) and Figure 4-15 of WMO (2007).
2.23

Stratospheric Ozone and Surface UV

section. As noted by Stimpfle et al. (2004), the cross section recommended by IUPAC (2007), based solely on the 

laboratory study of Huder and DeMore (1995), also yields 

a value of β inconsistent with field observations. The JPL 

06-2 cross section is consistent with the field observations 

(overlap of error bars) and the Burkholder et al. (1990) 

value is most consistent (β near unity).

Figure 2-10 (right) shows a similar comparison for 

the laboratory measurements of σClOOCl(λ) published since 

2007. Here, the JPL 09-31 value of k1a is used. As noted 

above, von Hobe et al. (2009) normalized their spectrum 

to the JPL 06-2 peak cross section. Use of this cross section leads to values of β that are outside of the 1σ uncertainty on the measured value of [ClO]2

/[ClOOCl]. Clearly, if the von Hobe et al. (2009) spectrum were normalized 

to a higher peak, as perhaps is warranted by more recent 

measurements of the peak cross section, then β would lie 

closer to unity (β scales in an approximately linear fashion 

with the peak cross section). The JPL 09-31 spectrum and 

cross sections (unchanged since JPL 06-2, except for the 

uncertainty limits) is consistent with the field observations 

(overlap of error bars). The Papanastasiou et al. (2009) 

spectrum and cross sections results in the best agreement 

(value of β close to unity).

Many of the ClOOCl laboratory studies to date report the wavelength dependence of the ClOOCl spectrum 

normalized to an absolute value of the cross section near 

the peak. Typically, the JPL 09-31 peak value of 6.4 × 

10−18 cm2 at 244 nm has been used (e.g., Huder and 

DeMore, 1995; Pope et al., 2007; von Hobe et al., 2009). 

The analysis presented in Figure 2-10 is complicated by 

the fact that three laboratory studies published in 2009 

report peak absolute cross sections greater than the JPL 

09-31 value (see Figure 2-9). To illustrate the importance 

of this scaling, the black dashed line in Figure 2-10 (right) 

shows the result of a calculation where the JPL 09-31 

cross section has been multiplied by a factor of 1.21, 

which scales the peak cross section of JPL 09-31 to the 

mean of the cross section values near 244 nm reported by 

Lien et al. (2009), Papanastasiou et al. (2009), and Wilmouth et al. (2009). The JPL 09-31 cross section, scaled 

in this manner, provides a better representation of field 

data than found using the recommended cross section. 

We highlight this sensitivity to illustrate the importance 

of this laboratory parameter for quantitative understanding of halogen photochemistry in the polar vortex. The 

Pope et al. (2007) absorption spectrum for ClOOCl is entirely inconsistent with field data, for any reasonable 

amount of scaling.

The bottom panels of Figure 2-10 show the photolysis first-order rate constant (J1b) as a function of SZA, for 

various cross section data sets. The shaded region shows 

propagation of the JPL 06-2 uncertainty (left) and the JPL 

09-31 uncertainty (right). This figure reveals consistency 

between laboratory studies of σClOOCl(λ) published after 

2007 and the uncertainty given by JPL 09-31, which was 

not based on these studies. Absorption cross sections derived from the measurements of Lin and co-workers and 

Wilmouth et al. (2009) at specific wavelengths in the photolytically active region are, as shown in Figure 2-9, generally consistent with the values reported by Papanastasiou 

et al. (2009). Therefore, all new laboratory studies of 

σClOOCl(λ) conducted since 2007 lead to a consistent picture 

of good understanding of polar ozone chemistry, in contrast to the state of knowledge that existed upon the 2007 

publication of the Pope et al. results.

Figure 2-11 extends Figure 2-10 by summarizing the high level findings for J1b/k1a from Stimpfle et al. 

(2004) as well as seven other studies, relative to the value 

of J1b/k1a recommended by JPL 06-2. Similar to the conclusions noted above, the seven other studies also suggest 

the value of J1b/k1a is as large as, or larger than, the value 

found using σClOOCl(λ) from JPL 06-2 and JPL 09-31. The 

field observations are most consistent with values of J1b/

k1a found using the cross sections from Burkholder et al. 

(1990) and Papanastasiou et al. (2009) (the central portion of most of the blue bars lies closest to the arrow denoting J1b/k1a from these two laboratory studies). Both of 

these laboratory studies measured absolute values of the 

ClOOCl cross section.

Figure 2-11 shows that field data are not consistent 

with J1b/k1a found using the recommendation for σClOOCl(λ) 

from IUPAC (2007) (based on Huder and DeMore, 1995) 

or von Hobe et al. (2009) (both spectra scaled to JPL 06-2). 

As described above, three 2009 studies report values of the 

peak absolute cross section greater than given in JPL 06-2. 

The value of J1b/k1a found using the spectra reported by 

von Hobe et al. (2009) and Huder and DeMore (1995) will 

exhibit closer agreement with the field data if these spectra 

are scaled to peak cross section values reported in the new 

laboratory measurements, as was illustrated by scaling the 

JPL 09-31 recommended spectrum in Figure 2-10. At 

present the peak cross section and thus scaling factor has 

an uncertainty in the 20 to 30% range. Most of the literature is based on the JPL 06-2 estimate (identical to the JPL 

09-31 value), which is at the low end of this range. If the 

peak cross section is revised by future evaluations, these 

revisions will likely indicate a greater role for ClO in polar 

ozone loss, since dO3/dt is roughly proportional to J1b.

Finally, Figure 2-11 shows that the field data are not 

consistent with J1b/k1a found using σClOOCl(λ) from Pope et 

al. (2007). The range of uncertainty in k1a (red bar) does 

not come close to encompassing the value of J1b/k1a found 

using σClOOCl(λ) from Pope et al. (2007) and the slight revision in k1a in JPL 09-31 is inconsequential. No reasonable scaling of the Pope et al. (2007) spectrum will resolve 

the inconsistency with field data. An analysis of satellite 

observations reported by Santee et al. (2008) supports the 
2.24

Chapter 2

0 0.5 1 1.5 2 2.5 3

 4.8 4.8

Shindell & de Zafra, 1996

Avallone & Toohey, 2001

Solomon et al. 2002

Vogel et al. 2003

Stimpfle et al. 2004

von Hobe et al. 2007

Schofield et al. 2008

JPL 06-2 uncertainty in k1a 

J1b

k1a

J1b

k1a obs JPL 06-2

/

Frieler et al. 2006 Ozone Loss 

JPope et al. 2007

J

J J J

JBurkholder et al. 1990

Papanastasiou et al. 2009

IUPAC JPL 06-2

von Hobe et al. 2009

2007

Figure 2-11. The ratio J1b/k1a inferred from analysis of daytime measurements of [ClO] or 

[ClO] and [ClOOCl] divided by the value of J1b/k1a from JPL 06-2, from various studies (blue 

bars; length of each bar represents range of uncertainty, generally 1σ). Horizontal black arrows denote lower limits for the ratio determined by two of the studies. The red bar shows the 

1σ uncertainty in k1a from JPL 06-2, evaluated at 190 K. The gray bar shows the value of 

J1b/k1a needed to match observed ozone loss in the Arctic and Antarctic vortices. The value 

of J1b/k1a using absorption cross sections of ClOOCl from various laboratory studies and data 

evaluations, relative to the JPL 06-2 value, is indicated by downward pointing arrows. The 

JPL 06-2 value of k1a was used throughout the analysis because most of the cited papers 

relied on k1a from JPL 06-2. The slight change in the high-pressure limit of this rate constant 

recommended by JPL 09-31 has a barely discernable effect on the appearance of this plot. 

Schofield et al. (2008) analyzed airborne in situ observations of [ClO] during the “self Match 

flight” of the European Polar Stratospheric Cloud and Lee Wave Experiment (EUPLEX) in the 

Arctic; von Hobe et al. (2007) examined airborne in situ observations of [ClO] and [ClOOCl] in 

the Arctic obtained during the SOLVE, EUPLEX, and Envisat Validation campaigns; Stimpfle 

et al. (2004) analyzed airborne in situ observations of [ClO] and [ClOOCl] in the Arctic obtained 

during the SOLVE campaign; Vogel et al. (2003) examined balloonborne in situ observations 

of [ClO] obtained in the Arctic during the Third European Stratospheric Experiment on Ozone 

(THESEO) 2000; Solomon et al. (2002) and Shindell and de Zafra (1996) analyzed groundbased observations of [ClO] over Antarctica; and Avallone and Toohey (2001) examined airborne in situ observations of [ClO] obtained in the Arctic during Airborne Arctic Stratospheric 

Expedition (AASE) I and II. Figure first published as Figure 3.1 of SPARC (2009).
2.25

Stratospheric Ozone and Surface UV

conclusion that the Pope et al. (2007) spectrum cannot be 

reconciled with atmospheric measurements of [ClO].

Section 4.2.1.3 of WMO (2007) examined our understanding of calculated and observed polar ozone loss 

rates. A major advance is recognition that calculated polar 

ozone loss rates using values for σClOOCl(λ) from Pope et 

al. (2007) are not consistent with ozone loss rates derived 

from observations (von Hobe et al., 2007; Santee et al., 

2008; Schofield et al., 2008; Kawa et al., 2009; SPARC, 

2009). A detailed summary is given in SPARC (2009). 

The gray bar in Figure 2-11 summarizes the understanding 

articulated in these studies: ozone loss rates derived from 

observations are consistent with values of J1b/k1a ranging 

from slightly less than that obtained using the JPL 06-2 

recommended cross sections to about a factor of 2 larger 

than that found using JPL 06-2, with best agreement found 

for the Burkholder et al. (1990) and Papanastasiou et al. 

(2009) spectrum and cross sections.

2.2.2.3 Other Issues Related to Polar Ozone

Chemistry

Kawa et al. (2009) used Monte Carlo model calculations to evaluate the impact and significance of the estimated uncertainties in the kinetic parameters given in JPL 

06-2 on polar stratospheric ozone loss. The simulations 

indicate that the most of the uncertainty in the calculated 

ozone loss and the rate of ozone loss is due to the uncertainty in the ClOOCl photolysis reaction. Uncertainties 

in the BrO + ClO reaction rate coefficient and its product 

branching ratio were also found to be important. Canty et 

al. (2005) note an inconsistency between theory and observation of nighttime chlorine dioxide (OClO) that could 

be resolved by higher yields of the branches of BrO + ClO 

that lead to ozone loss. Uncertainties in the rate coefficients for the reactions ClO + OH → HCl + O2 and Br + 

H2CO → HBr + HCO were found to be significant, but 

to have a smaller overall impact on the calculated polar 

ozone loss. The uncertainties in these processes were reevaluated by the NASA JPL Data Panel and reductions in 

the uncertainties were made in JPL 09-31.

Theoretical calculations by Matus et al. (2008) 

found the ClClO2 isomer was more stable than ClOOCl by 

3.1 kcal mol−1 at 298 K. However, as described in SPARC 

(2009), observational evidence suggests ClClO2 is not 

present in appreciable quantities during times of chlorine 

activation in the Arctic vortex. SPARC (2009) concluded 

that if unknown chemistry plays a role for the polar ozone 

loss mechanism, it can only be a minor modification of 

known mechanisms.

Analyses of field observations by von Hobe et al. 

(2007), Wetzel et al. (2010), and Santee et al. (2010) are 

somewhat more consistent with the revised JPL 09-31 

recommendation for KEQ at temperatures below ~210 K 

than with the JPL 06-2 recommendation. Observations 

obtained during nighttime conditions when photolysis 

is negligible all point to higher abundances of ClO and 

lower abundances of ClOOCl than found in models using 

JPL 06-2 kinetics. However, uncertainty in KEQ, which is 

believed to drive nighttime chemistry, has no bearing on 

the rate of polar ozone depletion (e.g., Kawa et al. (2009) 

and references therein). Nevertheless, there is interest in 

reducing the uncertainty in KEQ because accurate knowledge of this quantity will improve estimates of [ClOx] 

from nighttime observations of [ClO]. Such improvement is needed, for example, because representation of 

polar ozone chemistry in chemistry and transport models 

(CTMs) and chemistry-climate models (CCMs) is better 

evaluated by comparing modeled and measured [ClOx] 

rather than comparing modeled and measured [ClO]. If 

[ClO] is used, one must factor in SZA and temperature, 

which greatly complicates the comparison (Chapter 6 of 

SPARC CCMVal, 2010).

2.2.3 Polar Stratospheric Cloud Processes

As discussed in previous WMO reports, heterogeneous reactions on the surfaces of stratospheric particles at 

cold temperatures convert chlorine reservoir species that 

do not react with ozone, such as hydrogen chloride (HCl) 

and chlorine nitrate (ClONO2), to chlorine radical species 

that lead to catalytic ozone destruction (Solomon et al., 

1986). Liquid-phase binary sulfuric acid/water (H2SO4/

H2O) droplets, commonly known as background stratospheric aerosols, are ubiquitous throughout the stratosphere. Under cold conditions, these background aerosols 

take up nitric acid (HNO3) and H2O (Carslaw et al., 1994; 

Tabazadeh et al., 1994) and evolve into ternary HNO3/

H2SO4/H2O droplets, commonly referred to as supercooled 

ternary solution (STS) polar stratospheric clouds (PSCs). 

PSC particles may also take the form of H2O ice and solid 

hydrates of nitric acid, likely nitric acid trihydrate, or NAT 

(Voigt et al., 2000). Particle ensembles in the polar winter stratosphere are primarily mixtures of liquid (binary 

or ternary) droplets and solid particles (NAT and H2O 

ice) in varying sizes and number densities (e.g., Toon et 

al., 2000; Biele et al., 2001; Drdla et al., 2003). Chlorine 

activation rates on stratospheric particles are dependent 

on the uptake coefficient of the particle and the particulate surface area density (SAD) (Lowe and MacKenzie, 

2008). Both the uptake coefficients and the available surface area of the liquid particles are generally much higher 

than that of NAT PSC particles, making liquid particles 

much more efficient in chlorine activation (Portmann et 

al., 1996; Lowe and MacKenzie, 2008). Liquid particles 

also increase ozone loss by extending both the height 

range (Hofmann and Oltmans, 1993) and the season over 

which heterogeneous chemistry can occur (Portmann et 
2.26

Chapter 2

al., 1996). Solid particles play an important indirect role 

in ozone depletion by their influence on the abundance of 

gas-phase nitrogen. Formation and sedimentation of large 

NAT particles (Waibel et al., 1999; Fahey et al., 2001) 

can irreversibly redistribute HNO3 (denitrify) in the polar 

stratosphere, allowing the ozone depletion process to continue for a longer period by delaying the reformation of the 

chlorine reservoir ClONO2.

2.2.3.1 New Observational Data Sets

Since the previous Assessment, the PSC observational database has been greatly expanded by measurements from the Cloud-Aerosol Lidar and Infrared Pathfinder Satellite Observation (CALIPSO) satellite that was 

launched in 2006. Measurements from the polarizationsensitive lidar on CALIPSO (Pitts et al., 2007; Noel et al., 

2008) provide comprehensive daily information on the occurrence of PSCs in both the Arctic and Antarctic over the 

entire polar region, including the polar night that cannot be 

sampled by solar occultation instruments. The general climatology of PSCs included in previous Ozone Assessments (e.g., WMO, 1995; WMO, 1999) was established 

from long-term solar occultation data records (e.g., Poole 

and Pitts, 1994; Fromm et al., 2003), which were the only 

available data source at that time. Pitts et al. (2007) compared the PSC frequency observed by CALIPSO in 2006 

over the entire Antarctic region (50°S–82°S) with the PSC 

frequency derived by subsampling the CALIPSO database 

only at the time-varying latitudes observed by the SAM II 

(Stratospheric Aerosol Measurement II) solar occultation 

sensor, which ranged from 65°S in June to 80°S in September. As shown in Figure 2-12, the temporal distribution of PSCs derived from solar occultation data is not 

representative of the polar region as a whole. For example, in early June the solar occultation sensor samples only 

near the edge of the vortex (~65°S) and underestimates the 

PSC frequency of the vortex as a whole. In September the 

0 .1 0

10 0.

0.10

0.10

0.10

10 0.

10 0.

0 .3 0

0 .3 0

0 .3 0

30 0.

0 .5 0

0.50

Altitude, km

30

25

20

15

10

J t un Jul Aug Sep Oct 0.00

0.01

0.10

0.20

0.30

0.50

0.60

0.70

PSC Frequenc

y

0.40

CALIPSO PSC Frequency Subsampled at SAM II Occultation Latitudes

Antarctic 2006

Altitude, km

30

25

20

15

10

J t un Jul Aug Sep Oct 0.00

0.01

0.10

0.20

0.30

0.50

0.60

0.70

PSC Frequenc

y

0.40

CALIPSO Vortex-Wide PSC Frequency (50 S-82 S)

Antarctic 2006

o o

Figure 2-12. PSC frequency for the 2006 Antarctic season (adapted from Pitts et al., 2007). Top panel: PSC 

frequency as observed by CALIPSO over the entire polar region (50°S–82°S). Bottom panel: PSC frequency 

that is deduced by subsampling the CALIPSO data set at measurement latitudes seen by a solar occultation 

instrument (e.g., SAM II).
2.27

Stratospheric Ozone and Surface UV

occultation sensor samples only near the core of the vortex 

(80°S) and hence overestimates the PSC frequency of the 

vortex as a whole. CALIPSO provides a more accurate 

representation of PSC occurrence on a day-by-day basis 

over the entire polar region than the historical solar 

occultation-based observations. But on a season-long basis, both CALIPSO and the historical solar occultation 

data show that PSCs occur much more frequently and over 

a longer time period in the Antarctic than in the Arctic.

Optical modeling results (Pitts et al., 2009) suggest that from an ensemble point of view, CALIPSO PSC 

detection limits are as good as or better than those of historic solar occultation sensors. The ensemble detection 

limits are also comparable to those of ground-based and 

airborne lidars, except that CALIPSO cannot resolve very 

tenuous PSCs with the optical characteristics of a mixture of binary liquid aerosols and very low number densities (<3–5 × 10−4 cm-3) of NAT particles. For reference, 

NAT particle surface areas in these PSCs are <0.1 μm2

cm−3, which is about 10% of the surface area of the liquid 

binary aerosols and <1% of the surface area of a fully 

developed STS PSC.

2.2.3.2 PSC Composition

A more complete picture of PSC occurrence and 

composition has emerged since the previous Assessment. 

Pitts et al. (2009) examined the seasonal evolution of PSC 

composition utilizing CALIPSO lidar observations and 

found that the vast majority of PSC observations over 

the Antarctic from mid-June until mid-September consisted of liquid/NAT mixtures, while liquid STS clouds 

were predominant in the Antarctic in late May through 

early June and again in late September and October. Ice 

PSCs were much more episodic in nature and accounted 

for only about 10% of all CALIPSO PSC observations in 

the Antarctic. Pitts et al. (2009) found that mixtures containing NAT particles in higher number densities/volumes 

(similar to the so-called type 1a enhanced PSCs) are much 

more common in the Antarctic than in the Arctic, while 

the relative frequency of liquid STS clouds is higher in 

the Arctic than in the Antarctic. These results are generally consistent with the climatologies from ground-based 

lidars (Adriani et al., 2004; Maturilli et al., 2005; Blum et 

al., 2005; Massoli et al., 2006).

The spaceborne Michelson Interferometer for Passive Atmospheric Sounding (MIPAS) on Envisat, a limbsounding infrared spectrometer with full coverage of the 

Arctic and Antarctic regions, also provides information 

on PSC composition. Höpfner et al. (2006a) analyzed 

MIPAS measurements of mid-IR emissions by PSCs during the 2003 Antarctic winter and found evidence of NAT, 

STS, and water ice clouds. Information on PSC composition is derived from CALIPSO and MIPAS data using 

fundamentally different approaches, yet the results are 

consistent (Höpfner et al., 2009).

2.2.3.3 PSC Forcing Mechanisms

Several studies have examined formation mechanisms for PSCs. In the NH, Felton et al. (2007) showed 

that wave-induced temperature perturbations account for 

about 11% of PSC detections in the Arctic during the 

SOLVE-THESEO campaign. In the Antarctic, recent 

studies provide evidence that mountain wave PSCs may 

be a significant source of NAT particles that can be transported throughout the Antarctic polar vortex. Höpfner 

et al. (2006b) attribute the large-scale outbreak of NAT 

particles observed by MIPAS during June 2003 to heterogeneous nucleation on ice in the cooling-phases of largeamplitude mountain waves over the Antarctic Peninsula 

and Ellsworth Mountains. Eckermann et al. (2009) corroborate this hypothesis, showing that a small region of 

mountain wave activity over the Antarctic Peninsula on 

10–14 June likely served as the source of this circumpolar NAT outbreak. Eckermann et al. (2009) conclude 

that this is observational evidence of the “mother cloud” 

theory, which posits that vortex-wide NAT formation and 

growth are driven by ice formed in mountain wave activity. Noel et al. (2009) used CALIPSO data to examine 

wave-induced PSCs with near-unity optical depths, concluding that this relatively rare class of PSC can lead to 

widespread NAT and ice particles downstream from the 

mountain wave. Based on a combination of Polar Ozone 

and Aerosol Measurement (POAM) III aerosol extinction measurements and Challenging Minisatellite Payload 

GPS Radio Occultation (CHAMP GPS/RO) temperature 

measurements, McDonald et al. (2009) found that gravity wave-induced temperature perturbations may explain 

enhanced PSC incidence over the Antarctic Peninsula in 

June, while they contribute to only about 15% of the PSC 

observations later in the winter at higher latitudes. Innis 

and Klekociuk (2006), using lidar observations over Davis, 

Antarctica, also found that gravity wave perturbations influence PSC formation about 15% of the time. Wang et 

al. (2008) propose a different Antarctic PSC formation 

mechanism based on their analysis of observations by the 

NASA A-train satellites that showed two thirds of PSCs 

over west Antarctica and one half of PSCs over east Antarctica can be related to deep tropospheric cloud systems.

PSCs both influence and are influenced by climate 

change. Randel et al. (2009) report large trends in temperature in the lower stratosphere in spring. David et al. 

(2010) show that trends in the mean temperature and also 

trends in extreme temperatures are important for prediction of PSC occurrence. The long-term trend in globalmean lower stratospheric temperatures reported by Randel 

et al. (2009) is strongly driven by changes in stratospheric 
2.28

Chapter 2

ozone. Thus, both stratospheric cooling and the increase 

in PSC occurrence may reverse substantially as ozone will 

recover in the future. This conclusion is supported by 

model predictions by Hitchcock et al. (2009).

2.2.3.4 Use of Proxies to Represent PSC 

Processes

PSC occurrence is often represented in modeling and diagnostic studies of polar ozone loss by simple 

thermodynamic proxies. A proxy used frequently in the 

past is that PSCs are present if the ambient temperature 

is below TNAT, the theoretical threshold temperature for 

NAT existence. CALIPSO provides the first observational estimates of vortex-wide PSC areal coverage that can 

be used to directly assess the validity of simple thermodynamic proxies for PSC occurrence. Pitts et al. (2007) 

showed that the use of TNAT alone is a poor quantitative 

proxy for PSC occurrence and significantly overestimates 

the PSC areal extent; Figure 2-13 illustrates this point for 

the 2007 Antarctic season. The observation that TNAT

is not an accurate proxy for PSC occurrence is not surprising since analyses of CALIPSO and MIPAS data, as 

well as earlier studies, suggest that PSCs are primarily 

mixtures of liquid droplets and solid particles. However, 

as discussed below in Section 2.2.4, the empirical relationship between ozone loss and the volume of vortex air 

below TNAT indicates that TNAT is a useful gauge of low 

stratospheric temperatures that trigger chlorine activation 

and ozone loss, even though the activation is primarily 

occurring on liquid particles.

CALIPSO PSC Area Antarctic 2007

10

15

20

25

30

Altitude, km

May Jun Jul Aug Sep

2.00

0 .0

2

0 .0 2

0 .0 2

0 .0 6

0 .0 6

0 .0

6

6.00

0 .0 01

May Jun Jul Aug Sep Oct

0

2

6

10

14

18

22

26

30

PSC Area (× 106 km2)

T<Tnat Area Antarctic 2007

10

15

20

25

30

Altitude, km

May Jun Jul Aug Sep

.00 2

2.00

00 2.

2.00 6.00

.00 6

0 .0 6

10.00

.00 10

0 .0 01

.00 14

4.00 1

0 .0 41

00 8. 1

0 .0 22 0 .0 22

0 .0 62

May Jun Jul Aug Sep Oct

0

2

6

10

14

18

22

26

30

TNAT Area (× 106 km2)

Figure 2-13. Daily time series of CALIPSO PSC area (km2) versus area with T < TNAT for the 2007 Antarctic season (adapted from Pitts et al., 2007, and Pitts et al., 2009). Note that very tenuous PSCs containing a mixture of 

binary liquid aerosols and very low number densities (<3–5 × 10−4 cm−3) of NAT particles are below the CALIPSO 

ensemble detection threshold. NAT particle surface areas in these PSCs are < 0.1 μm2 cm−3, or about 10% of 

the surface area of the liquid binary aerosols and <1% of the surface area of a fully developed STS PSC.
2.29

Stratospheric Ozone and Surface UV

2.2.4 Arctic Polar Temperatures and 

Ozone

The annual cycle and variability for the minimum 

temperature poleward of 50° latitude in the Northern 

Hemisphere for 1979–2004 was shown in WMO (2007) 

(Newman and Rex et al., 2007, their Figure 4-1, top panel). The temperatures at 50 hPa for recent years fall within 

this range of variability. Temperatures in the Arctic winter stratosphere exhibit a high degree of natural variability. 

The winters 2006/2007 and 2007/2008 were among the 

ten coldest winters from 1965 to present. The 2008/2009 

winter was very cold in midwinter but PSC conditions 

were terminated by a strong warming in late January.

Polar ozone during recent Arctic winters remains 

low compared with values observed during the 1980s and 

continues to strongly vary interannually (Figure 2-14, top 

panel). The figure shows the minimum total ozone over 

the polar cap for March (calculated as the minimum of 

the daily average column ozone poleward of 63° equivalent latitude). This is an indicator of polar ozone loss that 

shows a reasonable correlation with observed chemical 

ozone depletion (Müller et al., 2008). In the years since 

WMO (2007), the minimum spring ozone values over the 

Arctic polar cap remained in a range comparable with values prevailing since the early 1990s (Figure 2-14, top panel). This indicates that substantial chemical loss continues 

to occur in cold Arctic winters.

Arctic winter and spring ozone loss has varied 

between 2007 and 2010, but remained in a range comparable to the values that have prevailed since the early 

1990s. Chemical ozone destruction on the order of 100 

DU (about 80% of the values derived for the record cold 

winters of 1999/2000 and 2004/2005) is deduced for 

both Arctic winters 2006/2007 and 2007/2008 (derived 

from ozonesonde measurements following the approach 

described by Rex et al., 2006). A strong reduction in 

column ozone during these winters is in accordance with 

strong chemical destruction at 475 K in mid-March reported for 2006/2007 based on Odin data (Rösevall et 

al., 2007) and 2007/2008 based on Microwave Limb 

Sounder (MLS) data (Kuttippurath et al., 2009). Reliable ozone loss estimates are not possible for the Arctic 

winter 2008/2009 because a strong midwinter warming 

in late January led to extensive mixing of air from low 

latitudes with the polar vortex air.

Tegtmeier et al. (2008) showed that the variability of (1) chemical loss of ozone in the Arctic and (2) the 

amount of ozone transported into the Arctic during winter 

each contribute about half of the observed variability of 

high latitude total ozone during spring (Figure 2-15). The 

supply of ozone due to transport has so far usually been 

larger than the chemical loss (with possibly very small 

net loss only in 1995, 1996, and 2000) such that the total ozone column has not declined in the course of Arctic 

winters and the extreme ozone anomalies associated with 

the Antarctic ozone hole are not observed in any longterm Arctic record (Solomon et al., 2007). Tegtmeier et 

al. (2008) also showed that these terms are correlated and 

variability in both is driven by variability in wave driving from the troposphere, making both terms sensitive to 

potential changes in wave driving due to climate change.

Minimum of daily mean polar ozone (Φe > 63o)

320

340

360

380

400

420

Ozone column (DU)

March, Arctic

1980 1990 2000 2010

160

180

200

220

240

260

280

Ozone column (DU)

October, Antarctic

Figure 2-14. Time series of 

minimum total ozone (Dobson units) over the polar 

cap, for March in the Arctic 

(top panel) and October in 

the Antarctic (bottom panel), calculated as the minimum of daily average column ozone poleward of 63° 

equivalent latitude. Winters 

in which the vortex broke up 

before March (1987, 1999, 

2001, and 2006) are not 

shown for the Arctic time 

series. Figure adapted from 

Müller et al. (2008), updated 

using the NIWA combined 

total column ozone database (version 2.7).
2.30

Chapter 2

Several diagnostics use volume of vortex air below TNAT, the theoretical threshold temperature for NAT 

existence for typical values of stratospheric water and nitric acid mixing ratios (Rex et al., 2004; Rex et al., 2006; 

Tilmes et al., 2004). This volume has commonly been 

referred to as VPSC. Because VPSC calculated in this way 

is not a good proxy for the existence of NAT particles (see 

Section 2.2.3.4 above), other proxies for the onset of polar 

heterogeneous chlorine activation have been considered 

that involve also the surface area density of binary sulfate aerosol particles, e.g., Tilmes et al. (2008a) consider 

the potential for activation of chlorine (PACl). However, 

observations show that TNAT may still be a useful proxy 

because the onset of chlorine activation does not necessarily occur on NAT particles (Toohey et al., 1993; SPARC, 

2009). Therefore, it is meaningful to consider a relation 

between ozone loss and such calculated values of VPSC. 

The ozone loss observed in Arctic winters since the previous Assessment falls along the relation between chemical 

ozone loss and VPSC that was discussed in the previous 

Assessment (WMO 2007, Figure 4-13).

Rex et al. (2006) noted a tendency toward higher 

extreme values in the overall volume of air at temperatures 

below the PSC threshold (VPSC) over the past four decades, 

i.e., a cooling of the “cold” Arctic winters. An update of 

this relation is shown in Figure 2-16. Values of VPSC since 

WMO (2007) were all less than that computed for NH 

winter 2004/05, but on a statistical basis established over 

the past four decades, a new maximum occurs about only 

once in five-year intervals. Therefore, it is not possible to 

draw a conclusion about the continuation of the changes 

in the severity of extreme values based on absence of a 

new record VPSC in the past four years. It is notable that 

three of the past four winters had VPSC values in the upper 

25% of the long-term distribution even though a midwinter major warming occurred in all of the winters since the 

previous Assessment.

2.2.5 Antarctic Polar Temperatures 

and Ozone

Winters remain very cold in the Antarctic, and the 

seasonal cycles for the minimum temperatures poleward 

of 50°S for recent years fall within the range of variability 

for 1979–2005 that is shown in the bottom panel of Figure 

4-1 of WMO (2007). Exceptional dynamical activity and 

associated warming, like that in Antarctic winters 2002 

and 2004, has not occurred since the previous Assessment, 

but in 2006 the springtime increase in the minimum temperature took place slightly later than usual.

The minimum of total ozone over the polar cap 

for October (calculated as the minimum of the daily average column ozone poleward of 63° equivalent latitude) 

has remained approximately at the low levels observed 

in the mid-1990s until 2009 (Figure 2-14 bottom panel). 

Chemical ozone loss in the Antarctic since the previous 

Assessment remained stable at the level of the mid-1990s. 

Ozonesonde observations at South Pole (Hofmann et al., 

2009a) consistently indicate that, with the exception of the 

years 2002 and 2004, more than 90% of the ozone is removed each year since the mid-1990s, at about 18 km (~70 

hPa). Maximum ozone losses at this altitude reach 99% 

(Solomon et al., 2005; Solomon et al., 2007). The lowest 

ozone partial column (2 DU) in the altitude range 14–21 

km observed at South Pole (in the time period 1986–2007) 

range of October observations

no data

1992 1994 1996 1998 2000 2002 2004 2006 2008

Year

400

300

200

100

0

-100

500

Total Ozone (DU)

vortex average

late March total ozone

chemical loss

dynamical supply

Chemical and dynamical contribution to the 

Interannual variability of the late winter Arctic ozone column

Figure 2-15. Interannual variability of 

the observed late winter (March) total ozone column (Dobson units) averaged over the polar vortex (black), 

the dynamical supply of ozone to the 

polar vortex (blue) and the chemical 

loss of ozone over the winter inside 

the vortex (red). The hashed range 

around 290 DU illustrates the very 

limited range of early winter (October) variability. The dashed black 

line close to the late winter observations is the sum of early winter ozone, 

dynamical supply and chemical loss 

(updated from Tegtmeier et al., 2008; 

based on ozonesonde data combined 

with meteorological data and transport modeling).
2.31

Stratospheric Ozone and Surface UV

occurred in 2006 (Hofmann et al., 2009a). The chemical ozone loss in the Antarctic vortex core for the altitude 

range 350–550 K deduced from satellite data for the 

period 1992–2005 is ~130 DU, more than found for even 

the coldest Arctic winters (Tilmes et al., 2006).

In summary, ozone loss during Antarctic winters 

remained stable with little year-to-year variability since 

the previous Assessment. This is expected since moderate 

changes in EESC are not expected to have a detectable 

effect on measures of Antarctic ozone loss (WMO 2007; 

Newman et al., 2006; see also Section 2.4.5 for further discussion). Since about 1997 the ODS amounts have been 

nearly constant and the depth and magnitude of the ozone 

hole have been controlled by variations in temperature and 

dynamics. The October mean column ozone within the 

vortex has been about 40% below 1980 values for the past 

fifteen years.

2.2.6 The Onset of Antarctic Ozone 

Depletion

The average total ozone poleward of 63° latitude 

in March (NH) and October (SH) is compared to an average of the observations between 1970 and 1982 to provide an indication of the amount of ozone lost each year 

due to polar processes (Figure 2-8). This choice of reference level is somewhat arbitrary. The 1970–1982 average includes observations from 1970–1972 and 1979–

1982 in the SH and from 1971–1972 and 1979–1982 in 

the NH. These observations and data from ground-based 

instruments at several locations suggest that ozone decreased between the early 1970s and the early 1980s in 

both hemispheres, but computing a trend from the satellite observations is not meaningful because of the data 

gaps. The ground-based data records have smaller gaps 

but exhibit year-to-year variations and are not representative of the entire polar region. A near-continuous, 

near-global record of satellite observations of total ozone 

has been available only since the Nimbus 7 Total Ozone 

Mapping Spectrometer (TOMS) instrument began taking 

data late in 1978.

2.3 SURFACE ULTRAVIOLET RADIATION

2.3.1 State of Science in 2006

Chapter 7 of the 2006 Assessment (Bais and Lubin 

et al., 2007) explored factors that affect surface UV irradiance, examined the methods of assessing this irradiance 

(ground-based measurements, satellite retrievals, and radiative transfer modeling), and reviewed measurement 

series, reconstructions, and numerical models of future 

scenarios to identify past and potential future changes in 

surface UV irradiance. Much of our understanding of UV 

radiative transfer remains valid today, while updates since 

2006 follow later in the section.

The influence of stratospheric ozone on the transmission of short wavelength UV radiation through the 

atmosphere is well understood. Since the ozone-related 

change in UV radiation is wavelength dependent (becoming negligible at wavelengths greater than 340 nm), the 

way in which UV radiation is defined or measured (e.g., 

1970 1980 1990 2000 2010

0

20

40

60

Year

VPSC [ 106 km3

]

Figure 2-16. Evolution of VPSC (km3) 

for the Arctic over the past four decades obtained from European 

Centre for Medium-Range Weather 

Forecasts (ECMWF) and Free University of Berlin (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 ERA-15 

data extended by operational analyses. The gray shading represents 

the uncertainty of VPSC assuming a 

1-K uncertainty of the long-term stability of radiosonde temperatures. 

Updated from Figure 4-3 of WMO 

(2007) and Rex et al. (2004).
2.32

Chapter 2

spectral, UV-B (280–315 nm) waveband, UV-B + UV-A 

(315–400 nm), or biologically weighted UV) determines 

its sensitivity to ozone. Broadband erythemal radiometers 

are the most widespread instruments for measuring surface 

UV irradiance, but they cannot identify the contribution 

of individual wavelengths to the observed change in erythemal irradiance and so cannot separate ozone-induced 

changes from other influences on UV radiation.

For UV effects for which there is a known action 

spectrum, e.g., erythemally weighted UV, the sensitivity 

to ozone changes is often expressed as a radiation amplification factor (RAF) (see Box 2-1), which differs from the 

monochromatic RAF.

Clouds, aerosol, surface albedo, and ozone all affect surface UV radiation. For many locations, the observed changes in UV radiation are influenced by changes 

in all of these. The importance of each varies with latitude, climate, and the amount of atmospheric pollution 

at each site. Cloud transmission is the greatest and most 

changeable atmospheric determinant of UV variability at 

Box 2-1. Radiation Amplification Factor for Erythemal Irradiance

The relationship between change in total column ozone (O3) and change in biologically effective radiation, e.g., 

erythemal irradiance (E), can be quantified using the radiation amplification factor (RAF) (Booth and Madronich, 1994). 

For example, RAF = 1.5 means that a 1% decrease in ozone will lead to a 1.5% increase in the biologically effective 

radiation. For small fractional changes in ozone (of the order of 10%) the RAF for erythema can be defined as

 

where ΔE and ΔO3 are the respective changes of erythemal irradiance and column ozone.

As the relationship between ozone and erythemal irradiance is nonlinear, the RAF is not a constant in all conditions but depends on factors that change the shape of the solar spectrum, primarily solar zenith angle and ozone column, 

as illustrated in Figure 1. While for solar zenith angles (SZAs) from 0° to 50° and typical midlatitude ozone the RAF for 

erythema is 1.1 ± 0.1, for larger SZAs and large total ozone values the RAF gradually decreases.

0.2

0.4

0.6

0.8

1

1.2

1.4

Solar Zenith Angle

Radiation Amplification Factor

RAF between 100 and 600 DU

RAF at 100 DU

RAF at 200 DU

RAF at 300 DU

RAF at 400 DU

RAF at 500 DU

RAF at 600 DU

0 10 20 30 40 50 60 70 80 90

Box 2-1, Figure 1. Radiation amplification factor for erythemal irradiance as a function of SZA, for several total 

column ozone amounts. RAF values were calculated in SZA-steps of 10° and steps of 10 DU in ozone column. 

The first value in each SZA-bin refers to 100 DU; the last value refers to 600 DU (from WMO, 2008).

RAF = E

E

O3

O3
2.33

Stratospheric Ozone and Surface UV

the surface. Although UV radiation is reduced slightly 

less than visible radiation by clouds, heavily overcast 

conditions can still reduce surface UV irradiance by 90%. 

Broken-cloud conditions can produce short-term enhancements in UV irradiance up to 25% over clear-sky conditions. Clouds vary on many timescales and thus hinder 

attempts to identify small ozone-induced changes in UV 

irradiance from ground-based measurements at locations 

where ozone depletion has been modest.

Pollutants in the lower troposphere, notably ozone, 

nitrogen dioxide, and sulfur dioxide, are strong attenuators 

of UV radiation. Ozone and sulfur dioxide (SO2) attenuate at UV-B wavelengths, while nitrogen dioxide (NO2) 

attenuates most strongly in the range from 330 to 400 nm, 

with a smaller influence in the UV-B waveband. The UV 

response to changes in these gaseous pollutants, and UVabsorbing aerosols such as organic carbon from motor 

vehicles or biomass burning, may be larger than the UV 

response to changes in stratospheric ozone. In some locations increased pollution and aerosols may have masked 

an increase in UV irradiance due to ozone decrease in the 

1980s and 1990s, and present cleaner air policies may hide 

the effect of ozone recovery on UV irradiance.

High surface albedo (reflectivity) greatly enhances surface UV irradiance where there is extensive snow 

and/or ice cover. Enhancements of erythemal irradiance 

of ~20% have been reported in regions with snow cover; 

even greater enhancements are seen in Antarctica due to 

very clean, high-albedo snow over a large area.

In the previous Assessment, ground-based UV measurement records that began in the 1990s covered a period 

with no significant ozone trends except at high latitudes, 

and reducing the uncertainties in quantifying surface UV 

irradiance remained a challenging task. Ozone-related 

changes in UV irradiance were apparent in unpolluted regions of the SH where previous increases in surface UV 

irradiance had ceased or reversed along with observed increases in ozone. At polluted midlatitudes in the NH, UV 

surface irradiance was generally still increasing by small 

amounts (a few percent per decade) at the time of the previous Assessment, but at unpolluted sites any increase had 

ceased. Short events with low column ozone and high UV 

irradiance had also been recorded at a variety of locations.

Satellite-derived erythemal irradiance estimates 

from 1979, based on ozone, aerosol, and cloud measurements, were being evaluated against ground-based data at 

the time of WMO (2007). Satellite data have the advantage of providing global coverage, but the retrieved products must be validated. TOMS (Total Ozone Mapping 

Spectrometer) data, and since July 2004 data from Ozone 

Monitoring Instrument (OMI) on the Aura spacecraft, had 

been analyzed. Satellite-estimated UV irradiance compares well with ground-based measurements in clean conditions, but at polluted sites the satellite retrievals failed to 

adequately account for tropospheric pollution, leading to 

overestimates of about 10% and up to 40%.

Both ground-based measured and satellite-estimated 

UV radiation data have increased in availability since 

WMO (2007). In addition to improved and reanalyzed 

ground-based data, there has been substantial progress in 

development of both satellite-derived UV products and 

ground-based reconstructed UV time series, and the climatologies and trend estimates derived from them. The following sections detail the advances in our knowledge and 

understanding since WMO (2007).

2.3.2 Update on Factors Affecting 

UV Radiation

The main factors affecting UV radiation are well established and recent studies were mostly focused on quantification of the effect of different factors on UV irradiance. 

In particular, better estimates of the effects of air pollution 

on UV irradiance, including their magnitude and wavelength dependence, were the subject of many recent studies.

2.3.2.1 Ozone Effects

Where ozone changes are large, e.g., in the Antarctic, they produce a clear signature in the surface UV 

radiation. Measurements from Summit Station, Greenland (72°N, established in 2004), show that for comparable solar zenith angles, the spring erythemal irradiance 

at the South Pole is 50–130% larger than at Summit Station, despite the similar location of the two sites on vast 

high-altitude ice sheets (Bernhard et al., 2008). Bernhard 

et al. (2008) also show that these differences are almost 

entirely caused by the difference in spring total ozone at 

South Pole and Summit.

Due to the ozone decline, the average erythemal irradiance measured at the Pole between 1991 and 2006 was 

up to 85% larger than the estimate for the years 1963–1980 

(Bernhard et al., 2010). At McMurdo, the average erythemal irradiance for October and November is estimated to 

be about 30–60% higher than historically.

The ozone-related influence on UV radiation can 

be identified from the spectral dependence of the changes, 

because ozone absorption increases strongly toward the 

shorter wavelengths in the UV-B region. Seckmeyer et al. 

(2008a) demonstrated that the ratio between UV values at 

Garmisch, Germany, and Lauder, New Zealand, as a function of wavelength clearly shows features related to the 

ozone absorption spectrum. Monthly erythemal UV radiation is up to 50% lower in the European summer compared 

to sites with comparable latitudes in New Zealand. These 

large differences are caused mainly by differences in total ozone (~15%), cloudiness, aerosol loading, and Sun–
2.34

Chapter 2

Earth separation. At similar SZAs, the ratio between UV 

irradiance values at Garmisch and Lauder as a function 

of wavelength clearly shows features related to the ozone 

absorption spectrum.

2.3.2.2 Other Influences on UV

Since ozone is not the only influence on surface 

UV, and may not be the only parameter to change in the 

future, or to have done so in recent years, we update and 

further quantify our knowledge of these other influences.

Clouds can cause large and rapid changes (increas can cause large and rapid changes (increases as well as decreases) in surface UV radiation and in total radiation (i.e., radiation integrated over the entire spectrum from about 300 nm to 3000 nm). On average, clouds 

have an attenuating effect of 15–32% in the UV waveband 

(Seckmeyer et al., 2008a; den Outer et al., 2005), but for 

any given situation the cloud effect will depend on the 

cloud type, depth, and distribution across the sky. Thus, 

cloud transmission of UV irradiance is difficult to quantify in sufficient detail to provide an exact determination 

of its effect at a given time and place (Thiel et al., 2008). 

This latter statement becomes important when considering 

satellite UV retrievals, detailed in Section 2.3.3.3, where 

cloud effects are averaged over a large field of view.

Another localized influence on surface UV radiation is air pollution, i.e., scattering and absorption by aerosols (black and organic carbon, hydrocarbons, dust, and 

smoke) as well as absorption by tropospheric O3, NO2, and 

other gases. Such pollution can reduce UV radiation by up 

to 15% at polluted sites, with reductions greater than 25% 

in highly polluted cities, e.g., occasionally in Los Angeles 

and frequently in Beijing. McKenzie et al. (2008) show 

that tropospheric extinctions in Tokyo account for much 

of the 40% reduction in UV-B radiation when compared 

to the clean atmosphere of New Zealand, and that UV-A 

radiation could occasionally be affected by high levels of 

tropospheric NO2.

Black carbon is the main absorbing component 

present in atmospheric aerosols. In addition, soil dust 

absorbs radiation in the UV and visible wavebands, and 

some organic materials absorb at UV wavelengths. Recently there have been several studies including measurements of enhanced UV absorption due to organic carbon. 

The sources and formation of these compounds and their 

temporal variability, however, are currently not well 

understood (Andreae and Gelencsér, 2006; Bergstrom et 

al., 2007).

The wavelength dependence of aerosol absorption 

optical depth is often expressed as an Absorption Angstrom Exponent (AAE). The theoretical AAE value for 

black carbon, assuming spectrally constant refractive index, is one. The higher values (spectrally steeper absorption) are usually attributed to dust and absorbing organic 

components. Recent studies that include UV wavelengths 

show large values of AAE for organic species (Bergstrom 

et al., 2010; Barnard et al., 2008; Martins et al., 2009; M. 

Yang et al., 2009).

Corr et al. (2009) analyzed measurements at UV 

wavelengths from Mexico City and found that there is no 

significant single scattering albedo (SSA, i.e., the ratio of 

scattering optical depth to the total optical depth (scattering + extinction) of the atmosphere) variability between 

332 and 368 nm. Barnard et al. (2008) analyzed data from 

the same Megacity Initiative: Local and Global Research 

Observations (MILAGRO) campaigns and also concluded there was little SSA wavelength dependence between 

about 300 and 400 nm for aerosols in the Mexico City 

area. However, Barnard et al. (2008) found a steep increase in SSA between ~400 and 500 nm and attributed 

the enhanced absorption at UV wavelengths to organic 

matter. Marley et al. (2009) also concluded that enhanced 

UV absorption was due to organic aerosols.

Surface reflectivity in the UV is usually less than 

0.1, unless the surface is snow covered. The albedo for 

snow-covered surfaces varies with snow type and age, and 

can approach unity for fresh, pure snow (Wuttke et al., 

2006a; Meinander et al., 2008). The UV irradiance over a 

high-albedo surface is greater than that above a snow-free 

surface, and this effect can be further enhanced by cloudiness and the multiple reflections between snow and cloud 

surfaces (Wuttke et al., 2006a). Changes in the extent and 

seasonality of snow cover due to climate change could 

thus change surface UV.

Surface UV radiation is a response to the combination of all influences on UV radiative transfer. Thus, while 

latitudinal and seasonal variations for UV radiation are 

firstly determined by the solar zenith angle, instantaneous 

UV irradiance values at high-latitude sites may sometimes 

significantly exceed those at lower latitudes due to the 

effect of clouds, aerosols, or ozone, particularly at sites 

affected by the ozone hole (Diaz et al., 2006; Vernet et 

al., 2009). Under similar solar zenith angles, there can 

still be large differences in UV irradiance between urban 

and rural locations, or between sites in the NH and generally cleaner sites in the SH. McKenzie et al. (2006), using 

measurements from the United States Department of Agriculture UV network, show that peak UV Index values at 

45°S in New Zealand are approximately 40% greater than 

those at 45°N in North America. Subsequently, it has been 

shown that cloud effects are also less severe at SH sites in 

the South Pacific region compared with similar latitudes in 

Europe (Seckmeyer et al., 2008a).

Seasonal differences in surface UV irradiance are 

highly dependent on latitude, moderated by local cloud 

and aerosol climatologies. They also depend on the UV 

wavelength/waveband, or biological weighting, of the irradiance being considered or measured. For erythemally 
2.35

Stratospheric Ozone and Surface UV

weighted UV, the ratio between summer and winter mean 

noon erythemal irradiance is in the range 17–24 around 

55°N (24 in Moscow at 57°N (Chubarova, 2008), and 17 

at 53°N, representative of, e.g., Liverpool, U.K., or Berlin, 

Germany (Seckmeyer et al., 2008b)), while it is 5 at 35°N. 

These seasonal contrasts tend to be greater, latitude for 

latitude, in the SH (McKenzie et al., 2009).

2.3.3 Ground-Based and Satellite UV Data

The majority of sites making direct measurements 

of UV radiation became active in the 1990s. With rare 

exceptions, ground-based data prior to this are reconstructions. To estimate the long-term evolution of UV radiation at the surface they use proxy data. While the approach is not new, recent studies used it for the first time 

to reconstruct UV over a large area (Europe) and study 

regional effects of ozone and clouds on UV irradiance. 

Satellite retrievals of UV irradiance have the advantage of 

a nearly global coverage and are available since 1979, the 

beginning of the almost continuous record of total column 

ozone and surface reflectivity from satellite instruments. 

While total ozone data merged from several satellite instruments have long been available (see Section 2.1.3.1), 

it is only recently that satellite reflectivity measurements 

have been merged together into a uniform data set (Herman et al., 2009), making it possible to study long-term 

trends in satellite-derived erythemal irradiance over a long 

period.

2.3.3.1 Ground-Based UV Measurements

The main classes of instruments presently in use for 

measuring solar UV irradiance reaching the ground have 

been described in previous Assessments (WMO 2003; 

WMO, 2007). Broadband instruments provide a measurement across a specified waveband, most commonly 

designed to mimic the erythemal action spectrum (Hülsen 

et al., 2008; Webb et al., 2006). Multispectral filter instruments measure several narrow wavelength bands (usually less than 10 bands from 1 nm to 10 nm wide) at the 

same time or sequentially (e.g., Petkov et al., 2006), thus 

providing some spectral information. Spectroradiometer 

systems measure the spectrally resolved solar irradiance, 

ideally at a resolution of 1 nm or better. No instrument is 

perfect, with uncertainties in the data coming from a variety of instrument-dependent sources (see, e.g., Johnsen et 

al., 2008; Cordero et al., 2008a; Cordero et al., 2008b). It 

follows that instruments must be carefully characterized 

and regularly calibrated (Webb et al., 1998; Wuttke et al., 

2006b). Since standards of spectral irradiance have uncertainties in the UV-B of ~3%, Cordero et al. (2008a) 

estimate uncertainties between 4 and 6% in the absolute 

calibration of double monochromator-based spectroradiometers. Additional uncertainties (e.g., imperfect cosine 

response) result in an overall uncertainty for solar measurements of 7–9%. This accords well with the work of 

Gröbner et al. (2006), who show results of an intercomparison of 25 European spectroradiometers relative to a 

transportable reference spectroradiometer. Almost half of 

the instruments had absolute agreement with the reference 

spectroradiometer to within ±4% for UV-B and UV-A 

wavelengths. Broadband instruments, usually calibrated 

by a spectroradiometer, have uncertainties in the range 

7–16% (Gröbner et al., 2007), with imperfections in the 

cosine response being a major source of error.

While absolute calibration is important for comparing measurements at different locations, or between different techniques (e.g., ground versus satellite), the most 

important requirement for trend detection is good longterm stability in instrument sensitivity. Stability can be 

determined by repeated checks against standard lamps in 

the laboratory or by various techniques applied to the data 

obtained in the field, i.e., the Langley method (Slusser et 

al., 2000). Some instruments are known to be very stable 

against standard lamps and can produce radiance stability 

near 1% (Cede et al., 2006).

Despite the need for geographical comparison between traceable quality-controlled UV measurements, 

ground-based monitoring stations provide sparse coverage. Most UV measuring stations are located in the NH 

in Europe and North America, with gradually increasing 

numbers in Asia. There are a few stations in South America, New Zealand, Australia, and Africa. A well calibrated 

polar network of seven stations has been maintained by the 

U.S. National Science Foundation, while the international 

Network for the Detection of Atmospheric Composition 

Change (NDACC) also maintains a long-term database of 

quality-assured UV data from a small number of observation sites. The World Ozone and Ultraviolet Data Center 

(WOUDC), the repository for Global Atmospheric Watch 

(GAW) and GAW-related station data, holds both spectral 

and broadband data.

Given the limited time series of UV measurements at most ground-based sites, there have been recent 

attempts to extend some of the data sets back in time using 

other data sets available for the sites, as detailed in the 

next section.

2.3.3.2 Ground-Based UV Reconstruction

Consistent records of reliable UV data covering 

more than 15–20 years are generally not available. Recent 

studies show the possibility of calculating the surface UV 

irradiance using variables directly affecting UV radiation, 

e.g., total ozone and cloud characteristics from standard 

weather station observations. The reconstructed data sets, 

which can extend backward to the beginning of the UV 
2.36

Chapter 2

proxy observations, are valuable for examining the UV 

variability in periods without UV measurements.

Radiative transfer calculations and statistical models are additional tools used to reconstruct the UV irradiance reaching the Earth’s surface for data-poor periods. 

These calculations provide hypothetical irradiance that 

would be measured if input parameters (ozone, aerosols, 

and ground albedo) were known for selected measuring 

sites. However, uncertainties in the model input parameters yield uncertainties in the model outputs of about 6% 

for clean sites, and up to 20% for sites with very large 

aerosol load (Cordero et al., 2007; Badosa et al., 2007).

The most common hybrid algorithm calculates allsky UV irradiance by multiplying the radiative transfer 

model output for clear-sky conditions by an empirical 

cloud modification factor (CMF) to account for cloud 

attenuation effects. CMF represents a ratio between 

measured shortwave radiation and its potential clearsky value. The UV-relevant CMF is usually derived via 

model calculations from an empirical formula based on 

total irradiance, which is measured routinely on many 

meteorological stations (Staiger et al., 2008; den Outer 

et al., 2005; den Outer et al., 2010). Comparison of 16 

European UV reconstruction models (COST 726, 2010) 

shows that the models using total irradiance for the CMF 

calculation prevail over models based on other parameterizations of the cloud effects, since total irradiance represents all the influences on radiation transmission through 

the atmosphere.

Ultraviolet irradiance data have been reconstructed 

by merging five different UV reconstruction models and 

ground-based measurements for eight European sites from 

Finland to Greece (den Outer et al., 2010) for the period 1963–2004. Reconstruction techniques are based on 

neural networks (Junk et al., 2007; Janouch and Metelka, 

2007), and radiative transfer modeling (Lindfors et al., 

2007) or combine radiative transfer modeling with empirical relationships (den Outer et al., 2005; Kazantzidis 

et al., 2006). Besides total column ozone, all models use 

pyranometer data (total radiation) to determine the cloud 

impact on the UV irradiance. The five different models 

have been validated against ground-based measurements 

on a range of timescales using daily to yearly sums of 

erythemally weighted irradiance. Good correspondence 

between reconstructed and measured data was found for 

all-sky conditions at sites with concurrent observations 

of total radiation and total ozone used as input of the reconstruction model. The study by den Outer et al. (2010) 

compares the output of the models with erythemal doses 

measured at the eight European stations. Standard deviations in the ratios of modeled to measured daily sums vary 

between 10 and 30% depending on site and model, and for 

yearly totals are less than 5%.

While no single model is able to fully represent all 

sites (each performs best at its site of origin), the combined (best estimate) approach based on all models has 

resulted in significant reduction in uncertainties in the 

reconstructed data: the range of deviations compared to 

measurements is at least halved for the best estimate when 

compared to the individual models.

Fixed aerosol properties are usually assumed in 

the UV reconstruction model. Simulations over Europe 

have used gridded climatological values of aerosol optical 

depth in the UV range that are calculated on a monthly 

basis by combining AERONET (AErosol RObotic NETwork) and MODIS (Moderate Resolution Imaging Spectroradiometer) aerosol data (COST 726, 2010). Variable 

aerosol properties appear as input to the reconstruction 

model at a few sites (e.g., Moscow; Chubarova, 2008) and 

are important where there has been a significant change in 

aerosols (e.g., in Thessaloniki due to reduction of urban 

pollution). Kazadzis et al. (2007) found the decrease of 

aerosol optical depth to be 2.9% ± 0.9% per year at 320 

nm between 1997 and 2006. den Outer et al. (2010) discuss the dramatic effect of this on the positive trend in the 

erythemal UV daily sums, up to 1.4% ± 0.1% per year, 

whereas without the aerosol change the trend would be 

around 0.3% ± 0.1% per year. Several authors demonstrate that local pollution (aerosols and trace gases) over 

urban regions significantly affects surface UV irradiance 

levels (Chubarova, 2008; McKenzie et al., 2008; Panicker 

et al., 2009).

A further important input to reconstructions in 

some climates is elevated ground-albedo during the snowcovered season, which can be incorporated if snow data 

are available (Rieder et al., 2008; Pribullová and Chmelík, 

2008; COST 726, 2010).

While reconstructions provide a historical view of 

UV irradiance for as long as the required ancillary data 

are available, recent reconstruction data intended for trend 

analysis exist only for a small number of sites, concentrated in Europe and Canada (Fioletov et al., 2001).

2.3.3.3 UV Estimates from Satellite

Observations

Global estimates of long-term UV records from 

satellite ozone and reflectivity data were previously discussed using observations from Nimbus 7 Total Ozone 

Mapping Spectrometer (N7-TOMS), and have now been 

updated and extended up to 2008 by Herman (2010). The 

updated time series, 1979–2008, uses ozone data obtained 

from the NASA merged satellite data set of monthly and 

5° zonal average band values (Stolarski and Frith, 2006). 

The UV estimates are obtained from measurements of 

backscattered UV by using radiative transfer calculations 
2.37

Stratospheric Ozone and Surface UV

that include ozone, surface reflectivity, and aerosol estimates embedded in a Rayleigh scattering atmosphere adjusted for the local terrain height within a satellite field 

of view. Another TOMS-based global climatology was 

published recently (Lee-Taylor et al., 2009). This study 

also discusses changes in UV irradiance in the 1990s relative to the 1980s and separates contributions from ozone 

and cloud changes.

Recent validation studies of satellite estimates have 

focused on OMI data using comparisons to ground-based 

measurements (Tanskanen et al., 2006; Tanskanen et al., 

2007; Ialongo et al., 2008; Buchard et al., 2008; Weihs 

et al., 2008; Kazadzis et al., 2009a). All find that the 

OMI UV estimates are biased high, particularly at more 

polluted sites. In snow-free conditions, they are biased 

high by 10% compared with ground-based measurements; 

the high bias is up to 40% in some highly polluted environments. This bias is similar to that for TOMS-based 

estimates: analysis of Belsk (Poland) data shows that the 

ratio between satellite-derived and measured erythemally 

weighted doses calculated for the snowless period (April–

October) is stable over the period 1979–2008 (Krzyścin et 

al., 2011). The ratio for the earliest part of the satellite UV 

data (from TOMS measurements onboard of the Nimbus 

7 satellite in the period 1979–1993) is 1.23 ± 0.06, only 

slightly larger than that obtained from the measurements 

by the TOMS onboard the Earth Probe satellite (1.19 ± 

0.07 for the period 1996–2003) and from the measurements by OMI onboard the Aura satellite (1.19 ± 0.05 for 

the period 2004–2008).

One of the sources of differences is that satellite measurements represent a much larger region (OMI 

minimum pixel at nadir: 13 × 24 km2) than ground-based 

measurements. Ground-based measurement of erythemal dose at various sites within one OMI satellite pixel 

showed deviations of ±5% in cloud-free conditions, or 

20% if urban areas were included (Weihs et al., 2008; 

Kazadzis et al., 2009b). For partly cloudy conditions and 

overcast conditions the discrepancy of instantaneous values between the stations can exceed 200%. If 3-hourly 

averages are considered, the agreement is better than 20% 

within a distance of 10 km (Weihs et al., 2008). This spatial discrepancy can explain much of the random variation between ground-based and satellite data but does not 

generally explain a strong systematic bias. The largest 

relative differences between the satellite-derived and the 

measured irradiance are observed in urban areas, where 

UV-absorbing aerosols play an important role. The OMI 

overestimate of ground-based UV measurements may be 

partly explained by the lack of sensitivity of satellite instruments to the boundary layer (Tanskanen et al., 2007; 

Ialongo et al., 2008; Weihs et al., 2008; Kazadzis et al., 

2009a; Buchard et al., 2008). The bias between OMI and 

ground-based measurements increases with increasing 

aerosol absorption optical thickness (Arola et al., 2005; 

Kazadzis et al., 2009b; Ialongo et al., 2010). The largest positive OMI biases were found at 305 nm (the lowest common wavelength routinely available), reaching 

32% and 27% for cloudy and cloudless measurements 

(Kazadzis et al., 2009b), indicative of ozone variability in 

addition to aerosol variability.

Additionally, over snow-covered surfaces the OMIderived daily erythemal dose is generally lower than the 

ground-based measurement because the OMI surface UV 

algorithm uses climatological surface albedo that may 

then be lower than the actual effective surface albedo. 

Part of the problem is that a portion of the observed reflectivity may be incorrectly interpreted as cloud cover, which 

reduces the estimated irradiance. All-conditions data and 

snow-free data have been compared separately to evaluate 

the effect of albedo. For example, a recent comparison 

by Buchard et al. (2008) demonstrated that OMI overestimates erythemal daily doses by 14% for days without 

snow on the surface and only by 8% if days with snow are 

included in the comparison.

UV irradiance is also derived by using column 

ozone in conjunction with cloud fraction data from geostationary satellites in combination with polar-orbiting satellites (Verdebout, 2004a and 2004b; Wuttke et al., 2003; 

Gadhavi et al., 2008; Janjai et al., 2010). The difference of 

about 10% between satellite and ground-based measurements is due to the limited information from the boundary 

layer (Arola et al., 2009; Kazadzis et al., 2009a). Schallhart et al. (2008) developed a method to generate near real 

time UV-index maps from Meteosat Second Generation 

(MSG) for Austria. The method is similar to that developed by Verdebout (2000) but with an additional input 

from ground-based UV measurements.

2.3.4 Long-Term Changes in UV

2.3.4.1 Ground-Based Observations

Over much of the Earth’s surface, long-term increases in UV radiation due to decreases in ozone derived 

from records of ground-based instruments are small and 

difficult to separate from the sometimes larger effects of 

changes in clouds, aerosols, and pollutants. Spectral instruments at clean sites in regions with significant ozone 

change are in the best position to identify ozone-related 

changes in UV radiation, either from ozone depletion or 

recovery. Elsewhere, with small ozone changes, changes 

in climate (clouds/albedo) and human activities (pollution, 

gaseous and aerosols) may have greater importance for 

UV radiation than the ozone.

At clean-air sites such as Lauder, New Zealand, the 

increase in summertime UV irradiance reported previously 
2.38

Chapter 2

has not continued, and in recent years, the peak summertime UV amounts have been lower than in the late 1990s. 

Ozone is no longer decreasing at this site, with lowest 

values occurring in the summer of 1998/1999. However, 

year-to-year differences are comparable with the measurement uncertainty.

One of the longest series of UV measurements has 

been taken at Belsk (Poland). The erythemally weighted daily doses have been homogenized for the period 

1976–2008 to account for instrument differences and time 

drift. Figure 2-17 shows that the mean erythemal irradiance level at Belsk in the warm subperiods of the year 

(April–October) in the 2000s has been ~10% larger than 

the overall mean level for the whole period of observations 1976–2008 (Krzyścin et al., 2011). The increase for 

clear-sky days is inferred from a simple model taking into 

account only ozone anomalies multiplied by the radiation 

amplification factor (RAF) of 1.1 (see Box 2-1). However, for data measured under all atmospheric conditions, 

the erythemal irradiance continues to increase when ozone 

reduction ceases. Borkowski (2008) shows an increase of 

2.3% per decade in annual dose during the period 1976–

2006 using Belsk’s monthly data.

Similarly, but for a shorter time period, increasing 

trends in erythemal irradiance have been seen at Sapporo, 

Tsukuba, and Naha since the early 1990s. At these sites 

total ozone was at its lowest around the early 1990s and 

since then there has been little increase in ozone. The 

increasing trends in erythemal irradiance since the early 

1990s cannot be attributed only to changing ozone (JMA, 

2009). Changes in aerosol and weather conditions at the 

Japanese sites are suggested as reasons for the observed 

increases in UV radiation.

Analysis of surface 280–320 nm UV irradiance 

measured by Multi-Filter Rotating Shadowband Radiometer (UV-MFRSR) from the U.S. Department of Agricul- U.S. Department of Agriculture (USDA) network shows annual irradiance changes 

from −5% per decade to +2% per decade at eight stations 

with approximately 10-year records (Hicke et al., 2008). 

Interannual variability of surface UV-B radiation was 2 to 

5% of the mean. Trends at each site were calculated for 

individual months, but during most months trends were 

not statistically different from zero.

2.3.4.2 Reconstructed UV Data

Using the erythemally weighted yearly doses reconstructed by various models (see Section 2.3.3.2), den 

Outer et al. (2010) find statistically significant UV trends 

for eight European sites. The trends vary between 3–6% 

per decade between 1980 and 2006. Upward trends were 

observed from 1980 to the mid-1990s for most sites, with 

levels in the 1980s being lower than the long-term average. Thereafter the rate of change altered at some sites and 

certainly cannot be extrapolated into the future. Ozone 

change accounts for ~1–2% of the increase per decade and 

clouds account for about 2–3% per decade. Observed increases in surface UV radiation are partly attributed to the 

decreased pollution since the late 1980s (e.g., Wang et al., 

2009). The European continent has “brightened,” possibly 

due to air pollution abatement policies (e.g., Kazadzis et 

al., 2007; Ruckstuhl et al., 2008).

Lindfors et al. (2007) reconstructed erythemally 

weighted irradiance back to the early 1980s at four Northern European stations based on measured total ozone and 

total irradiance. The reconstructed time series show an 

increase in erythemal irradiance (3–4% per decade) for the 

stations having a clear increase in total radiation. Feister 

et al. (2008) found a similar long-term pattern for Central 

European stations, with a decrease between the 1950s and 

1980 followed by an increase, resulting in UV irradiances 

~5–10% higher than the overall (1950–2004) mean level 

at the end of the time series. Sites in Finland show similar 

long-term behavior since the early 1980s, but an increase 

is found between 1951 and the end of the 1960s, opposite 

to the Central European sites.

Junk et al. (2007) and Feister et al. (2008) recon (2008) reconstructed series of erythemal irradiance (see Feister et al., 

2008, Figure 12) since 1893 based on sunshine duration 

1974 1978 1982 1986 1990 1994 1998 2002 2006 2010

-15

-10

-5

0

5

10

15

20

-15

-10

-5

0

5

10

15

20

Observation: clear-sky

Observation: all-sky

Model: clear-sky

Belsk: 1976–2008

April – October

Deviation (%)

Year

Figure 2-17. Deviations (%) of the monthly mean erythemal dose relative to the long-term (1976–2008) 

monthly means averaged over the warm subperiod 

of the year (April–October). The homogenized results of measurements by various broadband instruments at Belsk, Poland, for the period 1976–2008 

are shown for all-sky conditions (blue dots) and 

clear-sky conditions (magenta dots). Curves illustrate the smoothed data. The dashed magenta line 

shows the modeled ozone signal in the UV series 

(corresponding total ozone deviations are multiplied 

by RAF = 1.1, see Box 2-1). Adapted from Krzyścin 

et al. (2011).
2.39

Stratospheric Ozone and Surface UV

measurements at Potsdam. Additional measurements of 

global and diffuse irradiation were included since 1937, 

and variability of total ozone was included since 1964. 

The analysis indicates a few percent decrease in erythemal dose between 1893 and about 1910, a slight increase 

in the 1950s, followed by a gradual decrease of a few 

percent until 1980 and a subsequent increase after 1980 

of a few percent to the highest levels in the time series. 

Chubarova (2008) discusses the long-term UV changes 

over Moscow using reconstructed time series since 1968. 

The overall trend for 1968–2006 is not statistically significant because of a significant decline (−11% per decade) in cloud transmission at the beginning of time series 

(1968–1980). Since 1980, the growth of ~6% per decade 

in the yearly sums of erythemally weighted doses has 

been caused by a decrease in total ozone (2.5% per decade), cloud (2.1% per decade), and aerosol effects (1.1% 

per decade). In Austria (Vienna and Sonnblick), changes 

in erythemally weighted doses relative to a reference decade (1976–1985) showed that changes in total ozone had 

a larger influence on erythemally weighted doses than 

changes in cloudiness; here ozone accounted for about 

66% of the annually averaged change. However, in recent years the relative influence of changes in cloudiness 

on UV-doses became larger (reaching 50%) over Vienna 

(Rieder et al., 2008). Curylo et al. (2007) used reconstructed erythemally weighted data for Poland since 1964 

for four stations, finding positive trends in the yearly data 

(with a maximum trend of about 4% per decade in Warsaw). The trends disappeared over some stations in summer due to increasing cloud attenuation.

Kvalevåg et al. (2009), addressing an extended period, found an extensive reduction of erythemally weighted irradiance of up to 15–20% over most land areas since 

1750. An increase is suggested only in polar regions, most 

strongly in the SH and associated with the appearance of 

the Antarctic ozone hole. The estimate is based on a UV 

reconstruction taking into account preindustrial (1750) and 

present (2000) atmospheric conditions including changes 

in the stratospheric and tropospheric ozone (within the 

last 2–3 decades) and centurial changes in SO2, NO2, the 

direct and indirect effects of aerosol changes, and albedo 

changes. The long-term changes in pollutants could mask 

an increase of surface UV radiation due to stratospheric 

ozone decline in recent decades. It seems that the increase 

of UV extinction by carbonaceous aerosols is mostly responsible for such compensation.

Figure 2-18 presents a summary of estimated 

changes in UV radiation based on the reconstructions at 

eight European locations (den Outer et al., 2010). The 10-

year running means of the yearly sums of erythemal irradiance are shown separately for all-sky and clear-sky conditions. To make the results of these eight different sites 

comparable, each reconstruction was first normalized with 

respect to the average level in 1983–2004, which is the 

overlap period of available data. The uncertainty bands 

shown in Figure 2-18 stem from the uncertainties in the 

reconstructed yearly erythemal sums. These uncertainties, 

assigned to each year separately, were derived from the 

long-term stability and underlying agreement of the models, and the agreement with actual UV measurements. The 

uncertainties to higher and lower values than the estimate 

are independent, and generally differ in magnitude. The 

high (low) limit of the final uncertainty band is based on 

the high (low) limits for the individual yearly sums. Summations and averaging in these processes were always carried out with statistical weights assigned to the individual 

data points. The final running means, as presented in Figure 2-18, are normalized again with respect to the period 

1970–1980. Note the divergence between the trends for 

clear-sky and all-sky data since the mid-1990s, similar to 

the observations in Figure 2-17.

In conclusion, reconstructed trends, like those observed from direct UV measurement, are a superposition 

of the ozone, cloud, and aerosol effects on UV radiation. 

The magnitude of the trend depends very much on the period selected for the trend analyses, and on local characteristics of the cloud and aerosol changes as well as ozone. 

The statistically significant positive trends in erythemally 

weighted irradiance identified at many stations in the last 

half-century are due to reduced attenuation of radiation by 

1960 1970 1980 1990 2000

-4

-2

0

2

4

6

Deviation (%)

Year

Figure 2-18. The 10-year running means of the yearly sums of erythemal irradiance at eight European 

sites for all-sky conditions (blue diamonds) and clear 

sky-only conditions (red dots). Prior to averaging, 

each reconstructed time series is normalized with respect to the 1970–1980 average. The light blue and 

red areas depict the uncertainty envelope for all-sky 

and clear sky-only conditions, respectively. Adapted 

from den Outer et al. (2010).
2.40

Chapter 2

clouds since the beginning of the 1980s, and the general 

tendency of declining ozone until the mid-1990s.

2.3.4.3 Satellite Estimates of Irradiance

Changes

Herman (2010) analyzed satellite data (total ozone 

from multiple satellites) to estimate zonal average percent 

changes in UV irradiance from 290 to 400 nm and percent changes in biologically weighted irradiance reaching 

the Earth’s surface (using the concept of RAFs and Beer’s 

Law) for the period 1979 to 2008.

As shown in Figure 2-19 (top), clear-sky biologically weighted UV irradiance increased significantly except in the tropics. The increase was caused by decreases 

in ozone amount from 1979 to 1998. Since 1998, ozone 

amounts and UV irradiance levels have been approximately constant. Also shown for comparison are the percent 

changes in irradiance at six wavelengths (305, 308, 310, 

315, 320, and 325 nm), which have different responses to 

ozone and contribute in different ways to biological effects. The annual estimates of zonal average UV irradiance changes are restricted to latitudes between 53°S and 

53°N to avoid SZA > 80° during winter solstice, where 

spherical geometry effects become significant.

When all atmospheric conditions are considered, 

changes in cloud and aerosol transmission have a significant effect on UV irradiance in both hemispheres. In the 

SH the annual average UV increase is partially offset by a 

decrease in cloud and aerosol transmission (hemispherical 

dimming), while in the NH the effect was minimal (Figure 

2-19, bottom).

For clear skies, the largest increases were at the 

higher latitudes and in the SH (about 8.5% for erythemally 

weighted irradiance at 50°S compared to 4% at 50°N). At 

30°S the increase was about 5%, which is comparable to 

the increase at 30–40°N. As mentioned earlier, the weakness in this method is that the satellite sensors do not adequately probe the lowermost regions of the atmosphere, 

so the method is insensitive to changes in UV-absorbing 

pollution in the boundary layer of the atmosphere.

The average all-sky changes in erythemally weighted irradiance are more similar at higher latitudes in both 

hemispheres: 3.1% at 50°N, near the Canadian border, and 

5.2% at 50°S latitude near the southern tip of South America. In the SH, changes in cloud transmission partially 

offset the clear-sky increase. The zonal average irradiance 

increases vary locally and regionally because of different 

amounts of local cloud and aerosol cover.

Figure 2-20 shows the monthly percent change in 

erythemally weighted irradiance, caused by ozone, from 

1979 to 2008. Erythemally weighted irradiance has increased dramatically at higher latitudes in the SH and 

moderately elsewhere. In the SH spring (October and 

November) the increases are similar to the NH spring 

(April and May) changes and are about 5% to 7% at 40° 

to 45° latitude. However the SH summer changes in 

December to February are much larger than in the NH 

(June to August) at latitudes >40°. The major population 

centers in middle and southern South America, Southern 

Africa, Australia, and New Zealand experience significant 

increases in erythemally weighted irradiance compared to 

30 years ago. The increases have occurred during most 

of the spring and summer when the solar UV irradiance 

 clear sky

 all sky

 clouds

Irradiance Change (%)

Latitude

Latitude

28

26

24

22

20

18

16

14

12

10

8

6

4

2

0

-2

Erythemal Irradiance Change (%)

-70 -60 -50 -40 -30 -20 -10 0 3 10 20 0 40 50 60 70

-60 -40 -20 0 20 40 60

12

10

8

6

4

2

0

-2

-4

305

ERY

325

320

315

310

308

Figure 2-19. (top) Percent changes in erythemal irradiance from total change in ozone amount for 30 

years (1979 to 2008). Also shown for comparison 

are six monochromatic irradiance changes. (bottom) 

The change in erythemal irradiance (solid black line) 

caused by changes in reflectivity and cloud transmission (solid gray line) and the changes in ozone 

amount represented by the clear-sky changes in erythemal irradiance (dashed gray line). Error bars are 

estimated from the linear least squares fitting procedure. Adapted from Herman (2010).
2.41

Stratospheric Ozone and Surface UV

exposure is at a maximum (more clear days as well as 

seasonally declining ozone going into the summer).

2.3.4.4 Consistency of UV Estimates from

Observations, Reconstructions, and

Satellite Data

Since most of the ground-based UV measurement 

records are relatively short, they are not suitable for validation of long-term UV trends derived from satellite data. 

Long-term records of reconstructed data have to be used 

instead. Figure 2-21 shows summertime daily erythemally weighted doses estimated from ground-based UV 

measurements, ground-based reconstructions, and satellite estimates for the Northern Hemisphere (May–August) 

and the Southern Hemisphere (November–February). The 

comparison was limited to the summertime to avoid conditions with snow on the ground, when satellite estimates are 

unreliable. In general, satellite estimates are able to reproduce year-to-year variability, although there is a substantial bias between them and ground-based measurements 

and reconstructions (discussed in Section 2.3.3.3). Linear 

trend estimates for the period 1980–2003 from groundbased reconstructions and TOMS satellite estimates show a 

substantial difference between individual sites that reflects 

the importance of local conditions, and is well captured by 

both data sets. The average of four linear trend values for 

seven European sites shown in Figure 2-21 is 4.8% ± 1.7% 

per decade for the reconstructed data and 5.0% ± 1.4% per 

decade for satellite estimates (the uncertainties represent 

the 95% confidence level of the average of trend values at 

individual sites). While satellite estimated irradiance data 

has a positive bias compared to ground-based data, this 

does not influence trend detection if the bias is consistent 

over time. The agreement between independent methods 

of measurement and trend analysis in Figure 2-21, and discussed in Section 2.3.3.3, implies that at least for the European sites explored, this is the case.

It should be noted that summertime trends over 

seven European sites with reconstructed data from Figure 

2-21 are noticably larger than those over Toronto, Lauder, 

and two Japanese stations and substantially larger than 

trends in zonal mean UV radiation over northern midlatitudes shown in Figure 2-19 (about 4% over the 1979–2008 

period). This is likely due to long-term changes in the 

cloud cover over Europe. The mean trend over the same 

seven stations (Sodankyla is excluded to avoid snow influence on satellite estimates) is 4.3% ± 0.9% per decade for 

all atmospheric conditions, but the UV trend due to ozone 

is only 1.5% ± 0.4% per decade or about 4% over the entire period (see den Outer et al., 2010, their Table 6, for 

the trend values at individual sites), which is in line with 

satellite estimates from Figure 2-20.

2.4 INTERPRETATION OF OBSERVED 

OZONE CHANGES

The focus of this section is to interpret the observed ozone changes reported earlier in the chapter. This 

is accomplished primarily through comparisons of observations with simulations produced by chemistry-climate 

models (CCMs). These models couple the general circulation of the atmosphere with a photochemical mechanism using the latest available evaluation of chemical 

reaction rates and photolysis cross sections (e.g., JPL 

06-2). Eighteen of these models were extensively compared with each other and with data in the second 

Chemistry-Climate Model Validation (CCMVal) exercise 

documented in SPARC CCMVal (2010). The models all 

produced simulations of the past, the future, and a 

present-day time slice. One of the models included a 

coupled ocean and some had interactive tropospheric 

chemistry. Most, but not all, simulated the 11-year solar 

cycle. Most also simulated the response to large enhancements of aerosol surface area following the eruptions of 

the El Chichón and Mt. Pinatubo volcanoes.

2.4.1 State of Science in 2006

WMO (2003) concluded that chemistry-transport 

models (CTMs) including observed changes in halocarbons, other source gases, and aerosols captured the longterm behavior of ozone in the midlatitudes. This conclusion, repeated in WMO (2007) following examination of 

Change in Erythemal Irradiance 1979–2008

Latitude

Month 12

11

10

9

8

7

6

5

4

3

2

1

-40 -20 0 20 40

Figure 2-20. Percent change for erythemally weighted irradiance caused by changes in ozone amount 

for the period 1979 to 2008 as a function of latitude 

and month. From Herman (2010).
2.42

Chapter 2

four additional years of measurements and analyses, was 

based on the faithful representation of general features of 

the observed ozone changes as functions of latitude, altitude, and season obtained from a suite of models (both 

CTMs and CCMs).

WMO (2007) pointed out unresolved issues relating to deficiencies of then-current models. These include 

(1) the failure of models to reproduce the difference between observed southern and northern middle latitude 

ozone trends; (2) the consistent simulation of a Southern 

Hemisphere effect of the Mt. Pinatubo aerosols on ozone 

contrary to observed ozone behavior; and (3) worse comparisons of upper atmospheric ozone trend derived from 

observations with trends produced by models with temperature feedback than from models without this important process. The models used in WMO (2007) did not 

include a contribution from short-lived bromine compounds, perhaps leading to an underestimate of the decline 

in ozone following injection of volcanic aerosols into the 

stratosphere. Finally, although reproduction of the vertical structure of the solar cycle variation of ozone could be 

an important critical test of models, observational uncertainties made this test of limited value.

WMO (2007) noted increased evidence that changes 

in atmospheric dynamics had a significant influence on 

ozone over the northern midlatitudes on decadal timescales. WMO (2007) concluded that changes in tropospheric and stratospheric dynamics were partially responN E

E

Daily mean UV (J 

m-2) Daily mean UV (J 

m-2) Daily mean UV (J 

m-2) Daily mean UV (J 

m-2)

Daily mean UV (J 

m-2) Daily mean UV (J 

m-2) Daily mean UV (J 

m-2) Daily mean UV (J 

m-2)

Daily mean UV (J 

m-2) Daily mean UV (J 

m-2) Daily mean UV (J 

m-2) Daily mean UV (J 

m-2)

Figure 2-21. Summertime (May–August for NH and November–February for SH) mean daily erythemal UV doses 

(joules per square meter) from ground-based UV measurements (red), ground-based reconstructions (green), 

and satellite estimates (blue). The dashed lines show the linear trends for 1980–2004 calculated from reconstructions (green) and satellite estimates (blue). The derived trend values in percent per decade are also shown.
2.43

Stratospheric Ozone and Surface UV

sible for both the ozone decline up to the mid-1990s and 

for the increase after that time.

2.4.2 Updates to Kinetic and 

Photochemical Data

There have been many revisions to JPL 02-25 

(Sander et al., 2003) in chemical rate coefficients, photochemical parameters, and estimated uncertainties reported 

in JPL 06-2 (Sander et al., 2006) and JPL 09-31 (Sander 

et al., 2009). These do not result in major changes in our 

understanding of the response of stratospheric ozone to 

release of ozone-depleting substances. Recent developments in kinetic parameters of importance to polar ozone 

chemistry are discussed in Section 2.2.2.

2.4.2.1 Update from JPL 2002 to JPL 2006

An important difference between JPL 02-25 and JPL 

06-2 is revision of the recommended rates for reactions of 

excited-state oxygen atoms O(1D) with molecular nitrogen 

(N2), nitrous oxide (N2O), H2O, and methane (CH4) based 

on several laboratory studies. Dunlea and Ravishankara 

(2004) show that the revised rate coefficient for O(1D) + 

H2O reduces the production rate of hydrogen radicals in the 

stratosphere by 10–15% compared to the 2002 rate coefficient recommendation.

JPL 06-2 introduced the reaction of O + BrONO2

for the first time in the NASA evaluation using data from 

Soller et al. (2001). Sinnhuber et al. (2002) and Salawitch 

et al. (2005) showed that including this reaction in chemical models increases the daytime bromine monoxide/total 

inorganic bromine (BrO/Bry) ratio by about 20%. The presence of a greater fraction of inorganic bromine in radical 

form increases the sensitivity of ozone to halogens as well 

as the contribution of bromine to EESC (Section 1.4.4). 

The CCMVal-2 comparison (Section 2.4.3.5) shows that 

many CCM groups did not add this reaction when adopting JPL 06-2 kinetics (Chapter 6 of SPARC CCMVal, 

2010). Including this reaction also impacts the amount of 

Bry inferred from measured BrO, one method for estimating the impact of very short-lived bromocarbons on the 

stratospheric bromine budget. The photochemical models 

used in this Assessment to estimate Bry from BrO have all 

included this reaction (Chapter 1, Table 1-14).

The most notable impact of kinetics changes between JPL 02-25 and JPL 06-2 is likely to be the increase 

in the computed BrO/Bry ratio. Other reaction rates constants updated in JPL 06-2 include reactions of chlorine 

atoms (Cl) with hydrochlorofluorocarbons (HCFCs) and 

hydrofluorocarbons (HFCs); the hydroperoxyl radical 

(HO2) self-reaction; OH reactions with oxygenated organics; the yield of pernitrous acid (HOONO) from OH + NO2

+ M and HO2 + NO2 + M; and the photochemistry for a 

number of organic carbonyl compounds. WMO (2007) 

used the kinetic evaluations from JPL 02-25 with a few 

updates. To our knowledge there are no published estimates of the impact on global ozone of the change from 

JPL 02-25 kinetics to JPL 06-2 kinetics.

2.4.2.2 Updates since JPL 2006

O(1D) Reactions 

Additional studies (Carl, 2005; Takahashi et al., 

2005; Dillon et al., 2008; Vranckx et al., 2008) have led 

to a further revision in the recommended O(1D) + N2O 

and CH4 reaction rate coefficients in JPL 09-31. Recent 

simulations using the NASA Goddard two-dimensional 

(2-D) model described by Fleming et al. (2007) show a 1 

part per billion by volume (ppbv) increase in NOy (~8%) 

due to the increase in the recommended rate coefficient for 

O(1D) + N2O → 2NO. Atmospheric measurements do not 

differentiate between these results. The same 2-D model 

also produces a 10–12% deeper minimum in the Antarctic 

springtime ozone with the JPL 09-31 rates versus JPL 06-

2, with about half of the difference due to the change in the 

O(1D) + N2O reaction rate coefficient.

HO2 + NO

Laboratory studies by Butkovskaya et al. (2007, 

2009) examined the reaction product yields of the HO2 + 

NO reaction, reporting that the yield of the HNO3 product 

increased at lower temperature (223–323 K) and higher 

pressure, 72–660 Torr (96–880 hPa), and in the presence 

of water vapor. Cariolle et al. (2008) tested the impact of 

the updated reaction rate coefficient on the HNO3 formation 

from the HO2 + NO reaction using 2-D and 3-D models and 

found a significant impact on nitrogen oxides (NOx) and odd 

hydrogen (HOx) in the free troposphere and O3 abundance 

in the troposphere at low latitudes. They found only small 

changes in minor species distributions in the stratosphere.

ClO + HO2

JPL 09-31 revised the recommendation for the rate 

constant for the reaction ClO + HO2 based on the new laboratory study by Hickson et al. (2007). The revised rate 

constant has been increased by 25% at room temperature 

and by 37% at 200 K. Kovalenko et al. (2007) showed that 

atmospheric measurements of hypochlorous acid (HOCl) 

obtained by several balloonborne instruments are consistent 

with a value for the ClO + HO2 rate constant considerably 

larger than values recommended by JPL 06-2 or reported by 

Hickson et al. (2007). Kovalenko et al. (2007) also showed 
2.44

Chapter 2

good agreement between measured and modeled HOCl 

using the older rate constant measurement of Stimpfle et 

al. (1979), which was weighted more heavily in recommendations prior to JPL 02-25. Comparison of measured 

and modeled HOCl serves as a proxy for the rate of ozone 

loss by the ClO + HO2 cycle, which is often the dominant 

halogen loss process in the midlatitude lower stratosphere. 

Kovalenko et al. (2007) suggest that models underestimate 

ozone loss by this process because models using this kinetic 

input commonly underestimate observed HOCl.

2.4.3 The Distribution and Variability of 

Stratospheric Ozone and Their 

Representation in Models

The detection and attribution of trends in ozone depend on both long-term changes in ozone and on the variability of ozone. In Section 2.1, it was noted that ozone 

concentrations over the last decade or so remained more 

or less constant. A central problem is attributing these 

changes to causes such as dynamical variability, long-term 

climate change, decreasing ODSs, or the changing phase 

of the solar cycle.

WMO (2007) relied on 2-D models, 3-D CTMs, 

and CCMs, a mix of models that included and excluded 

radiative feedback. The current Assessment makes nearexclusive use of 3-D CCMs because they represent the 

state of the art of our understanding of the physical and 

chemical processes controlling atmospheric composition. 

These models have been developed significantly over the 

last four years. Recently their radiative, dynamical, chemical, and transport processes have been subjected to extensive evaluation using a wide range of diagnostics derived 

from observations (SPARC CCMVal, 2010). Understanding the mean ozone distribution and ozone variability requires not only analyses of observations but also evaluations of the processes that control the mean distribution 

and variability found in models.

The following sections summarize some of the key 

conclusions drawn from extensive evaluations of the 18 

CCMs presented in the SPARC CCMVal report (2010). 

Table 3-1 of Chapter 3 gives a description of the CCMs.

2.4.3.1 Annual Cycle and Natural Variability

Stratospheric ozone is known to vary in response 

to natural factors, such as the QBO, El Niño-Southern 

Oscillation (ENSO), variations in transport associated 

with the Brewer-Dobson circulation, and dynamical variability associated with the annular modes. Ozone observations have demonstrated variations on many spatial and 

temporal scales. Diagnostics for each of the different 

sources of natural variability in stratospheric ozone have 

been applied to the chemistry-climate models participating in this Assessment. The relative importance of the different sources of natural variability in stratospheric ozone 

has been assessed using multiple linear regression. When 

possible, the connection between the sources of natural 

variability and ozone has been addressed by analyzing the 

processes that determine it. Understanding and quantifying the underlying natural ozone variations are necessary 

to quantify the impact of anthropogenic perturbations of 

the ozone layer and to make reliable predictions of future 

ozone abundances.

The annual cycle is a forced variation of the atmosphere that is reflected in ozone concentrations. Figure 

2-22 (panels a and b) shows the annual cycle in ozone concentration at 1 hPa for 40°S and 40°N derived from MLS 

data compared to the annual cycles calculated by CCMs. 

Figure 2-22 (panels c and d) shows the annual cycle also 

at 72°S and 72°N calculated in the lower stratosphere at 46 

hPa. The vertical and latitudinal distribution of the annual 

cycle in stratospheric zonal monthly-mean ozone is well 

represented in stratosphere- and mesosphere-resolving 

models with a few outstanding issues. In the lower stratosphere, models tend to have a late occurrence of the polar 

ozone depletion (Antarctic ozone hole, Figure 2-22c). In 

the upper stratosphere (Figure 2-22a) the zonal mean MLS 

data at 40°S show an interesting peak in May and June for 

three of the four years that is not reproduced by models. 

SBUV data show a similar peak.

Model simulations reproduce many of the key features of the observed interannual variability of column 

ozone. All models show the expected minimum in polar 

variability in the summer season. However, in the NH dynamically active period, most of the models underestimate 

the interannual polar ozone variability. In the SH some 

models overestimate while others underestimate interannual variability in ozone.

Most models reproduce the connections between 

the dynamical processes responsible for the interannual 

polar ozone variations and the ozone response. Models 

with poor performance in interannual polar variability 

also tend to perform poorly in the diagnosed dynamicsozone connections. There are various techniques that use 

observations to cause a model to produce a QBO. These 

improve simulated ozone variability. However, there are 

biases in the amplitude of the QBO ozone signal from 

these simulations that are comparable to biases for models 

with internally generated QBO signals.

2.4.3.2 Solar Cycle

Here we update the extensive discussion of the relationship between ozone and solar radiation associated 

with the 11-year solar cycle and the 27-day solar rotation 

period given in WMO (2007). Variations of total ozone 
2.45

Stratospheric Ozone and Surface UV

column in phase with the solar cycle of 2–3% were reported, along with the stratospheric ozone profile response as 

a function of latitude. The detailed mechanism to explain 

the ozone response observed in the lower stratosphere remained uncertain. The following updates the discussion 

of the vertical structure of the tropical solar response given 

in WMO (2007) using additional observations and model 

results (see also Gray et al., 2010).

Recent analysis of the CCMVal models shows that 

although the 11-year solar cycle in column ozone is reproduced by the models, the amplitude of the response varies 

among models. Differences in radiation, photolysis, and 

transport all contribute to the spread in response. The latitudinal dependence of the solar response in column ozone 

derived from observations compares better with that derived from present models than earlier studies (Austin et 

al., 2008). The large spread at mid to high latitudes is due 

to large interannual variability and limits the discussion of 

extratropical signals.

Recent observational studies using ground-based 

and satellite data sets agree with previous results, reporting ozone variations in phase with the 11-year and 

27-day solar cycle in the upper stratosphere (e.g., Fioletov, 2009; Remsberg, 2008) and the lower stratosphere 

(e.g., Tourpali, et al., 2007; Sitnov, 2009). Similarly, the 

vertical structure of the solar signal in SAGE I and II 

data (Randel and Wu, 2007) shows a maximum response 

in the tropical upper stratosphere, a smaller (statistically insignificant) response in the middle stratosphere 

(~30–35 km), and a secondary maximum in the tropical 

lower stratosphere. This structure of the stratospheric 

ozone response in the tropics is supported by analysis of 

ground-based Umkehr measurements at the Mauna Loa 

station (Tourpali et al., 2007).

O3 1 hPa 40°S

J F M A M J J A S O N D

-1.0

-0.5

0.0

0.5

1.0

Ozone Seasonal Cycle (ppmv)

Month

(a)

J A S O N D J F M A M J

-1.0

-0.5

0.0

0.5

1.0

O3 1 hPa 40°N

Ozone Seasonal Cycle (ppmv)

Month

(b)

O3 46 hPa 72°S

J F M A M J J A S O N D

-2

-1

0

1

2

Ozone Seasonal Cycle (ppmv)

Month

(c)

J A S O N D J F M A M J

-2

-1

0

1

2

O3 46 hPa 72°N

Ozone Seasonal Cycle (ppmv)

Month

(d)

Figure 2-22. (a) and (b): Ozone seasonal cycle (in parts per million by volume) from MLS data at 1 hPa compared to models at 40°S and 40°N. (c) and (d): Same but for 46 hPa and 72°S and 72°N. Heavy lines are MLS 

data for each of four years (2005–2008); light lines are from models. Adapted from Figure 8.2 in Chapter 8 of 

SPARC CCMVal (2010).
2.46

Chapter 2

The QBO complicates the statistical detection of 

the solar signal. Smith and Matthes (2008) found that the 

presence of a QBO in their simulation could contaminate 

the detection of the solar signal when 2–4 solar cycles were 

included in the analysis. The near coincidence of the El 

Chichón and Mt. Pinatubo volcanic eruptions in 1982 and 

1991 with two solar maxima also complicates the quantification of solar signals derived from data. Thus, there is 

still uncertainty in quantifying the solar cycle by statistical analysis from observations, particularly in the lower 

stratosphere. Chapter 8 of SPARC CCMVal (2010) finds 

that the direct solar response in temperature and ozone in 

the upper stratosphere is well represented, but the vertical 

structure in the tropics below 10 hPa varies considerably 

among the models and between models and observations. 

Chapter 8 also notes large uncertainties in the middle to 

lower stratosphere that are the result of short observational 

records, possible aliasing of signals from QBO and volcanoes, as well as possible nonlinear interactions of the solar 

cycle signal with QBO, ENSO, and volcanic signals.

Recent model analyses show an improved representation of the vertical distribution of the solar signal in ozone 

in the tropics as compared to WMO (2007). However, 

reasons for the better agreement in the 2-D and 3-D models are still under discussion and conflicting explanations 

for the tropical solar ozone response in the lower stratosphere are presented. Austin et al. (2008) noted that models participating in the first CCMVal exercise reproduced 

the observed vertical structure of the stratospheric tropical 

solar ozone signal, although the magnitude of the response 

was smaller than derived from observations. Climate forcings, including observed sea surface temperatures (SSTs) 

as well as time-dependent solar cycle forcings, were considered essential to simulating this response. Neither the 

QBO nor the upper atmospheric effects of energetic particles were found to be necessary to explain the observed 

solar cycle response of ozone or temperature. Large discrepancies among the models and a small (insignificant) 

solar signal found in the tropical lower stratosphere in 

models that did not include solar forcing suggest possible 

influences from other processes (SST variations and aliasing between the signals) or from random variability in the 

records (Austin et al., 2008).

Schmidt and Brasseur (2006) used the Hamburg Model of the Neutral and Ionized Atmosphere 

(HAMMONIA) to simulate the difference between solar maximum and solar minimum. They produced a solar signal in the tropical lower stratosphere using climatological sea surface temperatures and a repeating solar 

cycle. This solar signal is independent of the presence of 

a self-consistently produced QBO (Schmidt et al., 2010). 

Tsutsui et al. (2009) found a small, not significant (0.6%) 

solar cycle response of ozone in the tropical lower stratosphere in a simulation using the Whole-Atmosphere Community Climate Model (WACCM) forced with climatological SSTs and a time-varying solar cycle. However, 

McCormack et al. (2007) highlight the importance of the 

presence of a QBO for the solar signal in tropical ozone 

in their 2-D model simulations. Their internally generated QBO is itself modified by the solar cycle. Additional 

complications arise from aliasing between ENSO and the 

solar cycle in the tropical lower stratosphere (Marsh and 

Garcia, 2007). Recent statistical analysis applied to observations and model simulations includes a proxy term 

that accounts for ENSO variations (Randel et al., 2009; 

Chapter 8 of SPARC CCMVal, 2010).

In summary, the upper stratospheric ozone response 

(2–3% between solar minimum and solar maximum) is reproduced by the CCMVal models and is a direct radiative 

effect of heating and photochemistry. The lower stratospheric solar cycle in tropical ozone appears to be caused 

indirectly through a dynamical response to solar ultraviolet variations. The origin of such a dynamical response to 

the solar cycle is not fully understood.

2.4.3.3 Volcanic and Aerosol Effects

Sulfate volcanic aerosols affect stratospheric circulation and temperature, provide surfaces for chemical heterogeneous reactions, and lead to ozone depletion 

(Hofmann and Solomon, 1989). The aerosols produced 

by Mt. Pinatubo, which erupted in June 1991, decayed by 

1996. No major volcanic eruptions have occurred between 

1996 and present, but efforts to quantify fully the effects 

of volcanic aerosols on stratospheric ozone continue, in 

part due to the importance of quantifying potential effects of continuous injection of sulfur into the stratosphere 

that has been discussed as a “geoengineering” approach 

to counteract global warming from increased greenhouse 

gases (e.g., Crutzen, 2006; Rasch et al., 2008). Geoengineering is discussed further in Section 3.2.6 of Chapter 3.

1-D, 2-D, and 3-D models have been used to determine the chemical and dynamical effects of volcanic aerosols (e.g., Zhao et al., 1995; Tie et al., 1994; Stolarski et al., 

2006). The polar vortex strengthens in NH after a volcanic eruption because of the changes in the equator-to-pole 

temperature gradient in the lower stratosphere (Stenchikov 

et al., 2002). A stronger polar vortex increases the probability of formation of polar stratospheric clouds, thereby 

enhancing the rate of heterogeneous chemical destruction 

of stratospheric ozone (Tabazadeh et al., 2002). The effects on stratospheric ozone depletion in the wake of a volcanic eruption are largest when chlorine and bromine levels are largest (Tie and Brasseur, 1995). The surface area 

density (SAD) of sulfate aerosols in models, the crucial 

input parameter to simulate this aspect of ozone depletion, 
2.47

Stratospheric Ozone and Surface UV

is usually prescribed based on observations. Comparisons 

of simulations with observations suggest that there are no 

significant missing processes or greatly inaccurate reaction rates (e.g., Fahey et al., 1993; Dessler et al., 1997).

Analysis of CCMVal-2 models (SPARC CCMVal, 

2010) shows that the simulated post-eruption changes in total column ozone are well correlated with changes in lower 

stratospheric chlorine monoxide. Although most models 

use the same aerosol SAD data set to drive the anomalous 

post-eruption chemistry, the sensitivity to aerosols and the 

background SAD values differ among models, leading to 

different amounts of chlorine activation and associated 

ozone loss. Mäder et al. (2007) used statistical analysis of 

total ozone data from 158 ground stations to show that the 

SAD had the largest impact on the decrease of total ozone 

in the Northern Hemisphere. This is consistent with the 

conclusions in WMO (2007) that the Mt. Pinatubo aerosols 

appeared to have a stronger effect on ozone in the Northern 

Hemisphere than in the Southern Hemisphere.

Telford et al. (2009), using a nudged CCM (UK 

Chemistry and Aerosols (UKCA)) simulation, concluded 

that the depletion of stratospheric ozone was produced 

mainly as a result of photochemical changes due to heterogeneous reactions, but there was some evidence for 

dynamically induced ozone reduction, especially in the 

NH midlatitudes. Fleming et al. (2007) found that the 

interannual dynamical variability in two-dimensional 

model simulations acts to reinforce the ozone perturbation in the NH midlatitudes, while in the SH the observationally derived circulation mitigates or even cancels 

the aerosol-induced chemical ozone reactions. Robock et 

al. (2007), using National Centers for Environmental Prediction (NCEP) and European Centre for Medium-Range 

Weather Forecasts (ECMWF) analyses and Goddard Institute for Space Studies (GISS) Model E, and Feng et al. 

(2007), using a 3-D CTM, did not find a clear signal of 

chemical ozone loss in the SH midlatitudes after the eruption of Mt. Pinatubo. In contrast Brunner et al. (2006), 

using Candidoz Assimilated Three-dimensional Ozone 

(CATO), demonstrated a significant effect of volcanic 

aerosols on ozone in SH mid (south of about 45°) and 

high latitudes. In summary, observations and model simulations found a significant effect of Mt. Pinatubo volcanic aerosols on stratospheric ozone in the NH compared 

to the SH. The absence of a stronger observation of ozone 

loss in the SH is not yet understood.

Polar ozone is also sensitive to volcanic enhancement of stratospheric aerosols. There have been no new 

findings on this subject since WMO (2007), but results 

are summarized here due to the potential importance of an 

increase in stratospheric aerosols through geoengineering 

as discussed above and in Section 3.2.6. Studies of the 

Arctic winter 1991/1992 consistently find enhanced polar 

ozone loss due to the presence of Mt. Pinatubo aerosol 

(Rex et al., 2004; Tilmes et al., 2008b) although the quantitative estimates of the volcanically induced signal differ. Further, Portmann et al. (1996) showed that Antarctic 

ozone depletion in the 1980s and early 1990s was influenced by enhanced sulfate aerosol SAD, and speculated 

that the ozone hole might have been detected later than 

the mid-1980s if there had been no increase in aerosol 

loading from El Chichón and several earlier minor eruptions. Portmann et al. (1996) also suggested that future 

Arctic ozone depletion could be severe in unusually cold 

winters with large volcanic aerosol SAD present. Several 

studies attempt to discern the effects of volcanic aerosol 

perturbations on PSC characteristics (Deshler et al., 1994; 

David et al., 1998; Fromm et al., 2003). These studies 

are inconclusive, largely because it is much more difficult 

to distinguish PSCs – especially STS – when the background (non-PSC) sulfate aerosol level is elevated and 

varies temporally.

As noted above, there have been no recent major 

volcanic eruptions, but there is evidence that the background stratospheric aerosol layer is changing. Hofmann 

et al. (2009b) analyzed ground-based measurements to 

show that stratospheric aerosol amount exhibits an increasing trend since 2000, attributing the background stratospheric aerosol since 2002 to an increase in the SO2 emissions caused by an increase in global coal consumption, 

mainly in China. Hofmann et al. (2009b) estimate the contributions to the increase in the stratospheric aerosol from 

major volcanic activity and from increased tropical upwelling due to change in the Brewer-Dobson circulation, 

concluding that these are not significant. However, analysis of observations from the Cloud-Aerosol Lidar with 

Orthogonal Polarization (CALIOP) suggests that small 

volcanic eruptions may play a role. CALIOP data show 

significant changes in aerosol concentration in the tropical 

stratosphere between 2006 and 2008 (Vernier et al., 2009). 

CALIOP data showed slow ascent of the Soufriere plume 

in 2006 from 20 to 25 km, consistent with the BrewerDobson circulation and remnants of aerosols from minor 

volcanic eruptions such as Manam in 2005. Quantitative 

attribution of the sources of the observed change in background aerosols has not been accomplished.

2.4.3.4 Evaluation of Simulated Transport

The evaluation of transport emphasizes the model 

ability to reproduce observations that are directly related 

to large-scale physical processes affecting ozone distribution, such as tropical ascent and Antarctic vortex isolation. The reader is referred to Chapter 5 of the SPARC 

CCMVal (2010) report for detailed evaluation of individual models.

The distributions of long-lived trace gases in the 

stratosphere are controlled by transport processes, mainly 
2.48

Chapter 2

by the balance between the diabatic circulation and quasihorizontal mixing (e.g., Holton, 1986). Transport affects 

the ozone distribution directly and indirectly. In the lower 

stratosphere ozone is long lived, and large-scale transport 

from source regions is important to its spatial and temporal distribution. Transport affects ozone indirectly by 

determining the mixture of chemicals that affect ozone 

loss at a given location. In Chapter 5 of SPARC CCMVal 

(2010), diagnostics for circulation and mixing were developed from observations and applied to CCMs. These 

evaluations identified key processes essential for realistic 

transport. Chapter 5 concluded that for the credible prediction of future stratospheric composition, the following 

transport requirements are essential: (1) realistic tropical 

ascent in the lower stratosphere; (2) realistic mixing between the tropics and extratropics in the lower and middle 

stratosphere; (3) generation of an isolated lower stratospheric Antarctic vortex in spring; (4) local conservation 

of chemical family mixing ratios (e.g., Cly); and (5) good 

agreement on all mean-age diagnostics. A summary of 

the performance of participating CCMs on these criteria 

is given below.

An adequate tracer advection scheme must conserve 

chemical families such as total inorganic chlorine (Cly) for 

credible predictions of future ozone levels. Conservation 

of Cly means that models should not produce higher levels 

of Cly in the upper stratosphere than released in the form 

of organic chlorine at the surface. Three CCMs failed 

to conserve total chlorine, undermining their credibility 

for assessment of ozone trends due to chlorine. A fourth 

model had excess stratospheric chlorine due to insufficient 

tropospheric removal of hydrogen chloride (HCl).

For more than a decade, model transport has been 

evaluated using comparisons of simulated and observationally derived mean age at 50 hPa in the midlatitudes 

and tropics (e.g., Hall et al., 1999; Eyring et al., 2006). 

These comparisons revealed several issues with model 

transport. However, at 50 hPa, age distributions computed by many models participating in CCMVal-2 compare 

well with those derived from observations even though 

other diagnostics reveal persistent transport issues. The 

average mean age diagnostic (AMA) was developed in 

the SPARC CCMVal (2010) report to broaden the comparisons with observationally derived ages in order to assess a model’s overall transport fidelity. AMA is based 

on mean age comparisons at seven locations in the lower 

and middle stratosphere between 60°S and 60°N. The 

mean age is a sensitive function of both the circulation 

and mixing, and the distribution of mean ages throughout 

the stratosphere reflects the balance between them, which 

varies as a function of height. While compensating errors 

in ascent and quasi-horizontal mixing can result in good 

mean age in some altitudes, it is unlikely that a model 

will have perfectly balanced compensating errors in both 

the lower and middle stratosphere, making the AMA diagnostic a more stringent test of transport than mean age 

at 50 hPa alone.

Although AMA is a strict diagnostic, fortuitous 

agreement with mean ages derived from observations is 

possible because excessive recirculation increases mean 

age while a fast circulation decreases it. Together they 

can produce compensating transport errors. It is therefore 

important to independently assess both the circulation 

and mixing in models. Significant problems with tropical stratospheric transport have been identified in half of 

the 18 participating CCMs. Tropical ascent and tropicalmidlatitude (T-M) mixing across the subtropics are crucial 

to distributing ODSs in the stratosphere, and deficiencies 

in the transport of ODSs affect modeled abundances of 

the products of ODS destruction (Cly) everywhere. Ten 

of the 18 CCMs were found to have realistic ascent in 

the tropical lower stratosphere. Of the remaining eight, 

two models showed a very slow circulation, while six had 

faster-than-observed ascent rates. T-M mixing, essentially a measure of the degree of isolation of the tropics, was 

evaluated using as many as four diagnostics in the lower 

and middle stratosphere. Nine of the ten CCMs with realistic ascent rates also showed reasonable mixing across 

the subtropics. Of the other nine models, the two models 

with slow circulations had too little T-M mixing while 

the other seven showed too much mixing between the 

tropics and midlatitudes. Strong T-M mixing increases 

the simulated mean age by allowing older midlatitude air 

to re-enter the tropics and recirculate through the stratosphere, increasing mean age everywhere. The wide range 

of independent observations from which the diagnostics 

are derived strengthens the usefulness and credibility of 

these evaluations.

Other essential transport processes for simulation of the ozone response to changing chlorine levels 

are strong descent in the southern winter vortex and the 

presence of a barrier to mixing between the vortex and 

midlatitudes in early spring in the Antarctic. Chapter 5 of 

SPARC CCMVal (2010) showed that vortex isolation is 

uncorrelated with the AMA, tropical ascent, and T-M mixing and therefore must be evaluated in addition to the other 

circulation and mixing diagnostics. Of the nine models 

showing both good tropical ascent and T-M mixing, seven 

also showed realistic, isolated descent at the end of austral winter. Lower stratospheric vortex isolation in spring, 

which is crucial to maintaining the high levels of Cly necessary for a credible ozone hole, was found to be sufficient 

in only eight of 18 CCMs. Of the eight CCMs producing 

an isolated vortex, only four also showed realistic tropical 

ascent, T-M mixing, average mean age, and polar descent. 

Those models are CMAM, GEOSCCM, UMSLIMCAT, 

and WACCM. (See Table 3-1 of Chapter 3 for descriptions of the CCMs.)
2.49

Stratospheric Ozone and Surface UV

2.4.3.5 Evaluation of the Chemical Mechanism

and Its Implementation

Chapter 6 of CCMVal-2 (SPARC CCMVal, 2010) 

presents the first major attempt at quantifying the behavior of stratospheric chemistry modules within various 

global 3-D CCMs. Eight of the 18 models participating 

in CCMVal-2 took part in a photolysis intercomparison 

(PhotoComp). The chemical mechanism within CCMs 

was evaluated by comparison of radical species output by 

each CCM to calculations conducted using a benchmark 

photochemical steady state (PSS) model, constrained by 

the abundance of long-lived species and sulfate surface 

area density (SAD) from each CCM. This comparison 

was conducted for 14 of the 18 groups. Quantitative metrics were developed for both PhotoComp and PSS, based 

on equation (4) of Waugh and Eyring (2008). Careful attention was devoted to proper definition of uncertainties. 

The majority of models showed good agreement with the 

PhotoComp benchmark.

Models showed a range of agreement with the evaluation of fast chemistry, with some models showing very 

good to excellent representation of both radicals and their 

precursors. The radicals in some models show significant 

discrepancies with the PSS benchmark; in some cases 

the precursors disagree with a suite of atmospheric observations included in the evaluation (Chapter 6, SPARC 

CCMVal, 2010). The most realistic representations of fast 

chemistry were found to be provided by the WACCM, 

EMAC, GEOSCCM, CMAM, and UMSLIMCAT models 

(Figure 6.11 of SPARC CCMVal, 2010). Results for these 

models are highlighted in subsequent figures.

The PSS evaluation provides an important new tool 

for quantitative evaluation of the chemical mechanism 

within CCMs based on analysis of archived model output. A particular CCM model has been observed to exhibit a much larger sensitivity of ozone to anthropogenic 

halogens than any of the other CCMs; the PSS evaluation shows that this sensitivity is the result of omission 

of the HCl production from ClO + OH in their chemical 

mechanism (Chapter 6 of SPARC CCMVal, 2010). Previously, chemical mechanisms have been evaluated either 

by exchange of code or by various groups performing prescribed calculations, both of which are arduous tasks.

The comparisons carried out for CCMVal-2 also 

showed that models used widely varying distributions 

of sulfate SAD, even for simulations designed to simulate past behavior (Chapter 6 of SPARC CCMVal, 2010). 

These differences may be due to the specification of the 

SAD climatology as a function of geometric altitude, a coordinate not commonly used in CCMs.

One notable difference among the various CCMs 

was the abundance of Cly at the tropopause. Some models have near-zero Cly (<<50 parts per trillion (ppt)) while 

other models have much larger values (>>50 ppt). Models with high levels of Cly at the tropopause tend to have 

excess Cly throughout the lowermost stratosphere. These 

differences in Cly are likely due to various representations 

of HCl uptake and removal in the troposphere (Chapter 6 

of SPARC CCMVal, 2010). This is an important detail 

because models with the higher values of Cly in the lowermost stratosphere display greater sensitivity to future 

changes in stratospheric H2O and temperature than models 

with near-zero Cly.

2.4.3.6 Evaluation of Simulations of the Upper

Troposphere/Lower Stratosphere

Chemistry-climate models are important for disentangling and quantifying the effects of halogens on ozone 

from those of climate change. In the upper troposphere/

lower stratosphere (UTLS, approximately defined here as 

the region between 5 and 22 km), dynamical variability and 

its expected response to climate change are large. The ability of CCMs to represent the basic characteristics of the 

dynamical and chemical structure of the UTLS has been 

tested in Chapter 7 of SPARC CCMVal (2010; also see 

Gettelman et al., 2009; Gettelman et al., 2010). Various 

characteristics of the tropical and extratropical chemical 

and dynamical structures have been evaluated using both 

quantitative and qualitative diagnostics derived from robust 

relationships found in aircraft and satellite observations, or 

reanalyses data sets. The tested characteristics include the 

representation of the tropical tropopause layer, the tropopause inversion layer, the extratropical tropopause transition layer, and the seasonal cycles in tropical cold point 

temperatures, lowermost stratospheric mass, and tropical 

and extratropical water vapor and ozone.

Most CCMs represent the basic dynamical and 

chemical characteristics of both the tropics and the extratropics well. Models mostly reproduce the annual cycle 

of tropical cold point temperatures in the tropics, with the 

right amplitude and timing. However, they show some 

significant biases in the annual mean cold point temperature. Similar findings are found for the representation of 

the seasonal cycle in ozone. However, amplitude, timing, 

and mean value of tropical water vapor does not follow 

the tropical cold point temperature as expected, which indicates errors in transport and/or microphysics in the models.

In the extratropics, the dynamical structure including zonal mean winds, the seasonal cycle in the lowermost 

stratospheric mass, and the seasonality and latitudinal behavior of the tropopause inversion layer are well captured. 

This implies that both the climatological distribution of 

radiatively active species and the radiative transfer in the 

models are realistic. The models fail to reproduce the observed sharpness in the tropopause inversion layer, most 

likely due to limited spatial resolution. The seasonal cycles 
2.50

Chapter 2

in ozone and water vapor are generally well reproduced 

but indicate a too large (small) amplitude at 100 (200) 

hPa. The models also resolve the strong tracer gradients 

across the extratropical tropopause, indicating that simulation of the Brewer-Dobson and the tropospheric baroclinic 

circulations is at least qualitatively realistic. Models that 

perform particularly well in the extratropics are CMAM, 

E39CA, GEOSCCM, and WACCM (Chapter 7 of SPARC 

CCMVal, 2010). Models that perform less well have issues that include insufficient vertical and horizontal resolution or use of a diffusive numerical transport scheme.

Outstanding deficiencies in the models that affect the ozone distribution and introduce uncertainties in 

UTLS evaluations are the lack of comprehensive representations of tropospheric chemistry involving nonmethane 

carbons, of wet and dry deposition, and of lightning NOx. 

CCMs include varying lightning NOx parameterizations, 

however, the most common one dates back to Price and 

Rind (1992), which has been shown in recent chemistry 

transport model studies to be inadequate in generating 

observed NOx levels in lightning-affected regions in the 

upper troposphere (Hudman et al., 2007; Sauvage et al., 

2007; Parrington et al., 2008).

2.4.4 Recovery Detection and Attribution

Section 2.1 discussed the data obtained over the 

past decades including a time-series analysis directed toward determining the current stage of ozone recovery as 

defined in WMO (2007). Here we discuss the problem of 

attribution of changes in ozone to changes in chlorine and 

bromine in the stratosphere.

The past evolution of ozone has been influenced 

by changes in ODSs and also by changes in greenhouse 

gas (GHG) concentrations. GHGs influence stratospheric 

ozone in two ways. In the upper stratosphere, cooling due 

to increased GHG concentrations leads to an ozone increase 

by slowing the temperature-dependent catalytic loss cycles. 

In the lower stratosphere where ozone is long-lived, radiatively driven dynamical changes to the Brewer-Dobson 

circulation that result from increased GHG concentrations 

affect the ozone distribution through transport. Attribution 

of past trends is complicated by the need to untangle the 

effects of the increases and leveling off of ODSs from the 

effects of continued increases in GHGs. The implications 

of these processes on the future development of the ozone 

layer are discussed in detail in Chapter 3 of this Assessment.

2.4.4.1 Dynamical Contributions to

Apparent Trend

In addition to the slow trend in ozone expected due 

to changing GHGs, dynamical variability in short time 

series of measurements or model simulations can give rise 

to an apparent trend. WMO (2007) summarized a number 

of studies using chemistry and transport models (CTMs) 

driven by analyzed winds and studies using empirical relations between planetary wave activity and circulation to 

address the question of dynamical contribution to Northern Hemisphere midlatitude ozone trend. WMO (2007) 

concluded that about 30% of the observed negative trend 

from 1979 to the mid-1990s could be attributed to changes 

in the lower stratospheric circulation based on both empirical and model studies.

Numerous model and observational studies have 

demonstrated the importance of dynamical contributions 

to ozone trends. In an extension of previous studies, Feng 

et al. (2007) used a CTM to show how spurious trends 

could result from the use of analyzed winds. They concluded that their simulation with nearly constant ODSs 

could reproduce the increase in NH midlatitude ozone 

from the 1990s. Brunner et al. (2006) empirically analyzed ozone data to show that the Brewer-Dobson circulation was a dominant contributor to interannual variability 

at both high and low latitudes. They further showed that 

this variability accounted for some of the ozone increase 

seen in the Northern Hemisphere since the mid-1990s. 

Fleming et al. (2007) used a 2-D model with winds derived from planetary wave forcing derived from ERA-40 

and also from NCEP. They found a good fit to the overall ozone column variability from 1979 through 2005 

that appears to have a dynamical contribution to both the 

downward trend between 1979 and the mid-1990s and the 

upward trend since then. Harris et al. (2008) examined 

both chemical and dynamical influences on decadal ozone 

trends, primarily in the Northern Hemisphere. They found 

that recent increases in ozone were linked to dynamical 

processes in the lower stratosphere.

Overall, studies agree that dynamical variability contributed to both the downward trend in the NH 

through the mid-1990s and to the upward trend in total 

ozone since the mid-1990s. In addition, Fleming et al. 

(2007) found little impact of dynamical variability on the 

fit to the EESC curve for the overall time period from 

1979 through 2005.

2.4.4.2 Greenhouse Gas Effects on Ozone

Trends

In addition to the possible trends due to interannual variability, the change in composition due to the 

increasing GHGs may cause ozone trends through changes 

in chemistry and dynamics, as noted previously. CCMs 

consistently predict a strengthening of the stratospheric 

Brewer-Dobson circulation (BDC) due to climate change 

(Butchart and Scaife, 2001; Butchart et al., 2006; McLandress and Shepherd, 2009) that leads to changes in 

the spatial distribution of ozone (Shepherd, 2008; Li et 
2.51

Stratospheric Ozone and Surface UV

al., 2009; Hegglin and Shepherd, 2009). Shepherd (2008) 

showed that CMAM results indicated an eventual recovery to values above 1980 (or 1960) (i.e., super-recovery) 

in the mid and high latitudes and under-recovery in the 

tropics. Both Shepherd (2008) and Li et al. (2009) found 

nearly latitude-independent increases in upper stratospheric ozone due to temperature decreases in CMAM and 

GEOSCCM, respectively. Both studies also found lower 

stratospheric ozone decreases in the tropics and increases 

elsewhere due to the simulated change in the BDC. In 

CMAM the effects from the upper and lower stratosphere 

largely cancel in the 60°S–60°N mean total ozone. According to CMAM, total ozone in the tropics will remain 

below pre-1980 values, whereas in the midlatitudes, and in 

particular in the NH between 35°N and 60°N, total ozone 

will increase to values about 5% above pre-1980 levels 

in the second half of the 21st century. The GEOSCCM 

calculates lower stratospheric ozone changes that approximately cancel when averaged globally (Li et al., 2009), 

so that the resulting change in global mean total ozone is 

dominated by the upper stratospheric ozone increase.

The studies by Shepherd (2008) and Li et al. (2009) 

showed the combined effect of GHGs and ODSs on stratospheric ozone, but did not quantify past and future changes to changes in ODSs or GHGs. Shepherd and Jonsson 

(2008) argued that carbon dioxide (CO2)-induced cooling 

had masked about 20% of the ODS-induced ozone depletion. Waugh et al. (2009), using calculations from the 

GEOSCCM to separate the effect of ODSs and GHGs, 

found that simulated stratospheric ozone changes during 

the past decades were clearly dominated by ODSs, except 

for the tropical lower stratosphere. Here the calculated 

ozone change during past decades due to the increased 

strength of the BDC (enhanced tropical upwelling) 

exceeded the calculated effect of ODSs.

CCMs suggest that GHG-induced changes in temperature and in the BDC may partly be responsible for past 

ozone changes. Therefore, it is important to assess the realism of the simulated circulation changes. Changes in the 

strength of the Brewer-Dobson circulation have not been 

unambiguously detected from observations. Engel et al. 

(2009) have analyzed balloonborne measurements of CO2

and sulfur hexafluoride (SF6) over the past 30 years to derive possible trends in stratospheric age of air. They found 

a small but insignificant increase of the age of air in the 

midlatitude middle stratosphere, in contrast to the small decrease in age of air due to the increased BDC calculated 

with CCMs. The results from this analysis can neither confirm nor deny the model results because of the large error 

bars and shortness of the record (see also Waugh, 2009). 

Because the simulated change in the BDC is most apparent 

in the tropical lower stratosphere where the ozone vertical 

gradient is steep, CCMs predict a change in ozone there that 

may be detectable in a long, stable time series of observations. There are significant issues with ozone trends in the 

tropical lower stratosphere due to insufficient length of the 

observational record as discussed below in Section 2.4.5.3.

2.4.4.3 Polar Loss and Dilution to Midlatitudes

WMO (2007) discussed in detail the importance of 

export of ozone-depleted air from the polar vortices to midlatitude trends, attributing about 30% (50%) of the northern (southern) midlatitude trends in total ozone for 1979 

to 1995 to this process. Results from multiannual simulations are generally consistent with detailed process studies 

for a few specific years in both hemispheres. Andersen and 

Knudsen (2006) conclude that longitudinal differences in 

the export of ozone-depleted air accounts for the longitudinal dependence of winter and springtime ozone trends at 

northern midlatitudes. The simulation of polar loss and its 

comparison to data are discussed in Section 2.4.5.4 below.

2.4.5 Simulation of Ozone Changes 

for the Last Three Decades

Previous sections discussed specific perturbations 

of stratospheric ozone in current chemistry-climate models. The CCMVal-2 report evaluates these models with 

a particular emphasis on their representation of processes 

by comparison to measurements of the current or recent 

past atmosphere. In this section we consider the ability of 

these models to reproduce the observed ozone trends over 

the last three decades by comparison to the data shown in 

Section 2.1.

2.4.5.1 Total Ozone Trends

Figure 2-23 shows the total ozone trend for 1979–

1995 as a function of latitude computed by multiple 

linear regression (MLR) as described in Section 2.1.2. 

The sensitivity of total ozone to EESC is calculated 

by applying the MLR to the entire data period (1978–

2009), and the trend is shown for the time period during which EESC increased linearly (1979–1995). The 

models are also fit with EESC from 1978 through the 

end of the REF-B1 simulation (usually 2004 to 2007). 

Note that the EESC fit obtained in this manner is more 

sensitive to the shape of the EESC function than to the 

length of record (see Section 2.1.2). The models generally reproduce the latitude dependence of the total ozone 

trend derived from the merged ozone data, but differ in 

some significant details. Most models produce a negative trend in the tropics while the data show no trend. 

Red lines are from three models (CMAM, GEOSCCM, 

WACCM) that rated highest in CCMVal-2 evaluations 

of transport, photochemistry, and UTLS discussed in 
2.52

Chapter 2

Sections 2.4.3.4–2.4.3.6 above. Blue lines are from two 

models (UMSLIMCAT and EMAC) that rated highest in 

one or two of the three CCMVal-2 tests.

In the southern midlatitudes, with one exception, 

the model trends vary with latitude in a manner similar to 

the trend derived from observations, but with significant 

spread in the magnitude of model trends. A similar spread 

is found for northern midlatitude trends. The model results 

show less spread when output beginning in 1960 is used to 

determine the fit to EESC (not shown), particularly in the 

north-polar region where interannual variability is large. 

In the tropics and midlatitudes, the models highlighted as 

having the best representations of transport, chemistry, or 

the UTLS generally are more similar to each other than 

the collection of all model simulations. The same is not 

true for polar regions. In the tropics most models show a 

negative trend in total column ozone that does not agree 

with the trend deduced from either satellite or groundbased data sets. One exception is the EMAC model that 

has a region of significant positive trend over a portion of 

the tropics due to an ozone increase in the tropical troposphere; however, the EMAC positive trend disagrees with 

the trend derived from measurements.

This raises the question of whether the discrepancy 

in the tropics could be the result of increased tropospheric 

ozone that is of the same magnitude as the decrease in 

stratospheric ozone. Most of the models in this comparison do not include tropospheric chemistry or tropospheric 

physical processes, such as wet and dry deposition and the 

production of NOx (and O3) by lightning. In the tropics, 

about 10% of the ozone column is in the troposphere. To 

offset the 1% per decade decrease found in model simulations would require a ~10% per decade increase in tropospheric ozone from the late 1970s through about the 

mid-1990s. Studies have searched for evidence of a tropospheric trend. Ziemke et al. (2005), using a 25-year record 

of tropospheric column ozone (TCO) derived from TOMS 

and SBUV(/2) data (1979–2003), showed a zero trend 

(± ~1DU per decade) in the tropics in the Pacific. Beig 

and Singh (2007) used TOMS data along with the Model 

for Ozone and Related Chemical Tracers (MOZART) 

model to derive a trend versus longitude that is negative 

but insignificant over most of the tropics, but positive and 

significant over the Indonesian region (~7–9%/decade). 

This trend appears to be the result of large fires in that 

region leading to ozone increases in the last few years that 

dominate the change in the record. Clain et al. (2009) derive marginally positive trends from ozonesonde records 

at Reunion Island (20.8°S, 55.5°E) and Irene, South Africa 

(25.9°S, 28.2°E). In summary, observational studies to 

date find no conclusive evidence for a tropospheric O3

trend large enough to explain the difference between simulated and observed trends in total column ozone, thus the 

reasons for model disagreement with the observed tropical 

trend remain a mystery.

Figure 2-24a-d shows the annual average total 

ozone time series from the merged ozone data set and from 

each of the CCMs for four regions. The quasi-global average from 60°S to 60°N is shown in Figure 2-24a. In general the models reproduce the overall behavior of the data. 

The four models showing the best transport representation 

are highlighted in red and show a smaller difference than 

the ensemble of all models. Some of the models include 

forcing due to solar cycle, volcanoes, and QBO while others do not. Although there is a spread in the magnitude of 

the ozone change from the beginning of the record to the 

last few years, the models agree with the data that the total 

ozone has been relatively constant over the last 7–8 years 

at about 2–4% below 1980 values. The midlatitude time 

series shown in Figure 2-24b and 2-24d show similar behavior with more variability in the data and among models. Figure 2-24c shows results for the tropics where an 

apparent solar cycle is prominent in the data and most of 

the models run below the data, consistent with the trends 

shown in Figure 2-23.

Figure 2-23. Latitude dependence of the annual total ozone trend (%/decade) from 13 of the 18 models 

used in CCMVal-2 and the merged ozone data set. 

Trends derived from measurements with 2σ uncertainties are indicated by vertical bars. Red lines are 

from three models (CMAM, GEOSCCM, WACCM) 

that rated highest in all three CCMVal-2 tests. Blue 

lines are from two models (UMSLIMCAT and EMAC) 

that rated highest in one or two of the three CCMVal-2 

tests. Model trends for 1978-1996 are calculated by 

fitting to EESC using output from 1978 through the 

end of the REF-B1 simulation. Model descriptions 

are given in Table 3-1 of Chapter 3.

Annual Trend Fit to EESC

-90 -60 -30 0 30 60 90

Latitude

-15

-10

-5

0

Total Ozone Trend (%/decade)
2.53

Stratospheric Ozone and Surface UV

2.4.5.2 Midlatitude Profile Trends

In previous Assessments it was noted that 2-D models without temperature feedback showed better agreement with observed ozone trends in the upper stratosphere 

than models that did include the feedback. The CCMs 

used here have improved somewhat in this regard. Figure 

2-25 and Figure 2-26 show trends deduced from measurements (SAGE, SAGE adjusted SBUV(/2), SBUV(/2), and 

ozonesondes) for the northern and southern midlatitudes, 

respectively. The trends calculated from SAGE for the 

upper stratosphere have been converted to pressure coordinates and have been adjusted to account for the temperature trend and the accompanying change of altitude at a 

given pressure level and the conversion of number density 

mixing ratio (see Section 2.1.4; Rosenfield et al., 2005). 

The model results presented in both figures are all from 

simulations including temperature feedback.

The northern midlatitude simulated trends show a 

spread in the maximum upper stratospheric depletion that 

is centered near the trend deduced from data. The red 

region in Figure 2-25a shows the range of the 3 models 

with high ratings in each of CCMVal-2 evaluations of the 

transport, chemistry, and UTLS. The red plus blue regions 

show the range of those 3 models plus the 3 that were with 

high ratings in at least one evaluation. The range of deduced ozone trends for these 6 models is narrower than for 

the entire suite of models (gray shading) and centered near 

the trend derived from satellite observations. The derived 

trends from Umkehr measurements are somewhat larger 

60oS–60oN

1980 1985 1990 1995 2000 2005 2010

Year

-15

-10

-5

0

Ozone Change (%)

(a)

35oN–60oN

1980 1985 1990 1995 2000 2005 2010

Year

-15

-10

-5

0

Ozone Change (%)

(b)

15oS–15oN

1980 1985 1990 1995 2000 2005 2010

Year

-15

-10

-5

0

Ozone Change (%)

(c)

35oS–60oS

1980 1985 1990 1995 2000 2005 2010

Year

-15

-10

-5

0

Ozone Change (%)

(d)

Figure 2-24. (a) Quasi-global time series of annual mean total ozone change (%) from 1980 value calculated 

from merged ozone data set in heavy black line and from 13 models in lighter lines. Changes in both models 

and observations are computed relative to the average for 1979–1981. Red lines are results from the four 

models that rated highest in the CCMVal-2 transport evaluation (CMAM, GEOSCCM, UMSLIMCAT, WACCM). 

The light gray line with large depletion appears to be an outlier to the rest of the models. (b) Same for northern 

midlatitudes. (c) Same for tropics. (d) Same for southern midlatitudes.
2.54

Chapter 2

than the models and the trends derived from satellite measurements. The spread among models is larger in the lower stratosphere but is still significantly reduced by use of 

the diagnostics to filter results. Here the models that perform best on the diagnostics tended to have smaller trends 

than those deduced from ozonesonde measurements.

Figure 2-25b shows trends calculated from model 

simulations for the period after 2000 when the beginning 

of recovery is expected. Although the simulated upward 

trends and trends from measurements are broadly consistent, the observed increase cannot presently be attributed 

to ODS decrease because of observational uncertainty, 

natural ozone variability, and stratospheric cooling. Note 

that both stratospheric cooling and ODS decrease lead to a 

projected upper stratospheric ozone increase.

In the southern midlatitude (Figure 2-26) upper 

stratosphere, trends from the models that perform best on 

the diagnostics are somewhat smaller than those deduced 

from measurements. In the southern midlatitude lower 

stratosphere, the range of ozone trends deduced from models is somewhat more narrow for simulations that perform 

best on the CCMVal-2 diagnostics. In both the northern 

35oS–60oS

-10 -5 0 5 10

Ozone Trend (%/decade)

100.0

10.0

1.0

0.1

Pressure (hPa)

-10 -5 0 5 10

100.0

10.0

1.0

0.1

EESC Fit

Figure 2-25. Altitude profile of calculated ozone trend 

(%/decade) at northern midlatitudes from 17 models. Trends 

derived from measurements are 

as defined in Figures 2-4 and 

2-7. Panel (a): Trends for the 

period 1979–1996 are obtained 

by fitting to EESC. Trends derived from measurements with 

2σ uncertainties are indicated by 

horizontal lines. The solid line 

is the trend derived from SAGE 

I + SAGE II data. The dashed 

line is trend derived from SAGEcorrected SBUV(/2) data. The 

dash-dot line is the trend derived from SBUV(/2) data. The 

dash-dot-dot line is the trend derived from sonde data. The long 

dashed line is the trend derived 

from Umkehr data. Red shading 

is the range of 3 models with high 

ratings in all 3 evaluations. Blue shading adds the range of 3 models that were highest in at least 1 evaluation. 

Panel (b): Altitude profile of trends calculated from model simulations for the time period 2000–2009, when the 

beginning of recovery is expected. Shaded regions are the same as in (a). Although the simulated upward 

trends and trends from measurements are broadly consistent, the observed increase cannot presently be attributed to ODS decrease because of observational uncertainty, natural ozone variability, and stratospheric cooling. 

Note that both stratospheric cooling and ODS decrease lead to a projected upper stratospheric ozone increase.

35oN–60oN

-10 -5 0 5 10

Ozone Trend (%/decade)

100.0

10.0

1.0

0.1

Pressure (hPa)

-10 -5 0 5 10

100.0

10.0

1.0

0.1

-10 -5 0 5 10

100.0

10.0

1.0

0.1

-10 -5 0 5 10

100.0

10.0

1.0

0.1

1979–1996 2000–2009

(a) (b)

Ozone Trend (%/decade)

Figure 2-26. Same as 2-25a for the southern midlatitudes.
2.55

Stratospheric Ozone and Surface UV

and southern midlatitudes, the transport diagnostic selects 

simulations with smaller ozone depletion. Although we 

do not have enough sonde data to compare to southern 

midlatitude trends in the lower stratosphere, it appears that 

the simulated trends are smaller than trends derived from 

measurements at the northern midlatitudes.

2.4.5.3 Tropical Profile Trends

In the tropical lower stratosphere the models predict 

a decrease in column ozone due to increased upwelling 

(e.g., Butchart et al., 2006), in contrast to trends deduced 

from total column ozone measurements. The trends in the 

tropics computed from the simulations by 17 models are 

shown as a function of pressure in Figure 2-27. Almost 

all of the simulations show a similar pattern, with a decrease in the upper stratosphere centered at about 2–3 hPa 

and a smaller, narrower decrease in the lower stratosphere 

between the tropopause and about 60 hPa. The increase 

in tropical upwelling bringing up air with lower ozone 

amounts into the lower stratosphere and mixing of ozonedepleted air into the lower tropical stratosphere may both 

contribute to this trend.

There is evidence for this lower stratospheric negative trend in data as shown by the SAGE I and SAGE I+II 

results in Figure 2-27. These trends are an update of Randel and Wu (2007). The results for SAGE I+II rely on the 

relative calibration of SAGE I to SAGE II, and are consistent with the results from SAGE II only and with the simulated trends. Brunner et al. (2006) used the CATO data 

set that uses measurements of total ozone and the analyses 

from ECMWF to infer global distributions of ozone. They 

found a narrow region of negative ozone trends in the tropical lower stratosphere and noted that these were consistent with those found for 1984–2000 in SAGE II data as 

reported in WMO (2003). As discussed in Section 2.4.5.1, 

the near-zero trend in tropical total ozone measurements is 

inconsistent with the negative trend found in the integrated 

SAGE I + SAGE II stratospheric profiles, and the tropospheric ozone column does not increase enough to resolve 

this discrepancy. The disagreement between trends obtained from vertically integrated SAGE data and trends 

obtained from the merged TOMS/SBUV(/2) column 

ozone data set was discussed in WMO (2007).

2.4.5.4 Polar Trends

Yang et al. (2008) report that Antarctic ozone is in 

the first stage of recovery. Ozone within the springtime 

Antarctic vortex is affected by both chemical and dynamical processes (Newman et al., 2006). Yang et al. (2008) 

used correlations between monthly means of total ozone 

column and temperatures to construct ozone anomaly time 

series, which reflect variations in ozone due to chemical 

forcing (Figure 2-28). The ozone anomaly time series 

reveals a statistically significant leveling off of ozone column since 1997, relative to the previous rate of decline. 

Yang et al. (2008) estimate ozone loss saturation by comparing the frequency distribution of measured ozone with 

the distribution expected from a reconstruction of ozone 

that hypothetically allows ozone abundances to drop 

below zero. This analysis indicates that the recent leveling off of the total ozone anomaly time series is due to 

changes in Antarctic halogen loading rather than loss saturation (bottom panel, Figure 2-28). They also identified 

the minimum of the second derivative of Antarctic EESC 

as a useful quantitative means to specify the break point 

for a piecewise linear fit to the data (further discussed in 

Chapter 1 of this Assessment).

As noted in the previous Assessment (WMO, 

2007), trends in Arctic ozone are much more difficult to 

assess due to strong year-to-year variability of meteorological conditions and the dependence of inferred trends 

in temperature or ozone on start and end date. Furthermore, the trend in Arctic temperature is not sufficient to 

assess the impact of long-term changes on ozone (see 

15oS–15oN

-10 -5 0 5 10

Ozone Trend (%/decade)

100.0

10.0

1.0

0.1

Pressure (hPa)

-10 -5 0 5 10

100.0

10.0

1.0

0.1

Figure 2-27. Ozone trends (%/decade) computed 

from CCMVal-2 simulations by 17 models as a function of altitude for the tropics. Black line: trend derived from SAGE II data alone. Purple line: trend 

derived from SAGE I + SAGE II data. The range of 

model simulations is as in Figure 2-25a.
2.56

Chapter 2

Section 4.1.1.1 of WMO, 2007). Since the previous Assessment, temperature conditions in the Arctic vortex 

have not been extremely cold for any particular ozone 

loss season. Chemical loss of Arctic ozone for recent 

years falls along the compact, near-linear relation with 

VPSC (volume of air in the vortex exposed to temperature 

below the threshold for formation of PSCs) identified by 

Rex et al. (2004).

Much progress has been achieved in the evaluation 

of the representation of polar ozone loss in CCMs since 

the previous Assessment. The CCMVal-2 effort has provided important new insight into the chemical, dynamical, 

and radiative properties of CCMs (Chapter 6 of SPARC 

CCMVal, 2010). The chemical modules contained in 

CCMs are similar to the schemes used in off-line 3-D 

CTMs and other models known to produce good simulations of polar ozone loss under specified meteorological 

conditions (e.g., Chipperfield et al., 2005; Frieler et al., 

2006; von Hobe et al., 2007). However, use of a realistic and complete description of the relevant chemical (and 

microphysical) processes related to polar ozone depletion 

is only the first step toward obtaining an accurate representation of springtime polar ozone loss in coupled CCMs. 

The host general circulation model (GCM) must also simulate the correct descent over the pole, the correct timing 

and isolation of the polar vortex, and realistic temperatures 

that allow PSCs to form in the model.

Most CCMs include simplified schemes for PSC 

microphysics. Chapter 6 of SPARC CCMVal (2010) analyzed how well observed distributions of HNO3, H2O, and 

HCl in the Antarctic were simulated by the CCMs. The 

models generally reproduced observed H2O but overestimated HNO3. The CCMs generally use simplified equilibrium denitrification schemes and this discrepancy for 

HNO3 will be affected by deficiencies in transport. Although conversion of HCl to active chlorine is essentially 

complete in most CCMs near 500–600 K, as expected 

from observations, the CCMs overestimate HCl at lower 

altitudes. This indicates incomplete chlorine activation in 

some CCMs, which would contribute to an underestimate 

of the simulated column ozone loss.

Figure 2-29 shows the time dependence of chemical loss of column ozone for the Arctic (top panel) and 

Antarctic (bottom panel) for 15 CCMs. Ozone loss derived from observations made by the Halogen Occultation Experiment (HALOE) instrument on the Upper Atmosphere Research Satellite (UARS) is also shown. The 

chemical loss of column ozone is calculated using analysis 

of the relationship between ozone and N2O using archived 

records from each CCM. Chapter 6 of SPARC CCMVal 

(2010) shows that this method to quantify chemical ozone 

loss is reasonable because the simulated ozone versus 

N2O relationship is well organized along isopleths, as observed. Also, Figure 2-29 shows that prior to ~1980 there 

Figure 2-28. Time series of Dobson/

Brewer measurements of total column ozone (Dobson units) in the 

core of the Antarctic vortex during October. Top panel: observed 

ozone column. Middle panel: resulting time series after removing 

effects of monthly mean temperature variations from the time series. 

Bottom panel: resulting time series 

after removing effects of loss saturation and monthly mean temperature 

variations from the time series. The 

blue line indicates the downward 

trend in column ozone for the respective time series, calculated for 

1979 to1996 and forecasted linearly 

afterward. Linear trends and 95% 

confidence intervals, DU/decade, 

are given. The red line in the bottom 

panel shows a fit of Antarctic EESC to the time series, considering data for the entire record. Years of “record 

cold” and of a sudden, “early warming” are indicated. Temperatures were warmer than usual in 1988, corresponding to the anomalously high levels of ozone observed that year. The vortex remained intact following the 

rise in temperature, hence 1988 is not classically considered to have experienced a sudden, “early warming.” 

From Yang et al. (2008), updated to include data for 2008 and 2009.

150

200

250

1979-1996 trend = -68.7 +/- 20.7 DU/decade October core ozone

Early

Warming Record

Cold

150

200

250

Ozone (DU)

1979-1996 trend = -49.6 +/- 9.9 DU/decade Temperature adjusted

80 82 84 86 88 90 92 94 96 98 00 02 04 06 08

Year

150

200

250

1979-1996 trend = -57.3 +/- 7.6 DU/decade Temperature adjusted

Saturation corrected

Antarctic EESC fit to data, 1979-2009
2.57

Stratospheric Ozone and Surface UV

is chemical loss of column ozone during the ozone loss 

season. Note that methyl chloride (CH3Cl) supplies a natural background level of ~0.6 ppb Cly and anthropogenic 

ODSs were also increasing during this time period. Cly

from both of these sources leads to chemical loss of ozone 

in the early part of the displayed time series, should meteorological conditions result in chlorine activation. Note 

that the loss due to CH3Cl is natural and helps determine 

background levels of polar ozone prior to the buildup of 

Cly due to anthropogenic release of ODSs.

Figure 2-29 shows that modeled chemical loss of 

column ozone compares better with loss derived from observations in the Antarctic than in the Arctic. There are 

long-standing discrepancies between model behavior and 

observations related to the polar lower stratosphere, particularly for the Arctic (e.g., Eyring et al., 2006). CCMs 

tend to exhibit a cold bias in the Antarctic lower stratosphere and a concomitant westerly circulation that is too 

strong. The polar vortex then persists later into spring 

than observed. Analysis of the CCM tracer transport 

shows that the different CCMs vary widely in their ability to produce an isolated lower stratospheric Antarctic 

vortex, with about half of them producing sufficient isolation for the altitude region and time period critical for 

ozone loss. In the Arctic, where observed temperatures 

are nearer the threshold for PSC formation, the CCMs are 

on average warm-biased and there is a large amount of 

model-to-model variation. The CCMs underestimate the 

observed frequency of cold Arctic winters. Because the 

CCMs are generally warm biased in the Arctic they tend 

to underestimate observed chemical loss of column ozone 

in the NH. In the SH CCMs are generally cold biased, 

and simulated column ozone loss generally scatter about 

the range of column loss derived from observations. In 

the SH ozone loss tends to saturate in the cold, isolated 

vortex. Another factor that may contribute to the spread in 

model results in Figure 2-29 is that although the standard 

CCMVal-2 runs only included bromine from long-lived 

Arctic: Vortex Core 350−550 K

1950 1960 1970 1980 1990 2000 2010

Years

0

20

40

60

80

100

120

140

Ozone Loss (DU)

 Observations 

 CAM3.5 

 AMTRAC3 

 WACCM 

 UMUKCA−METO 

 MRI 

 SOCOL 

 NIWA−SOCOL 

 CCSRNIES 

 EMAC 

 ULAQ 

 CMAM 

 LMDZrepro 

 UMSLIMCAT 

 GEOSCCM 

 CNRM−ACM 

Antarctica: Vortex Core 350−550 K

1950 1960 1970 1980 1990 2000 2010

Years

0

50

100

150

200

Ozone Loss (DU)

 Observations 

 CAM3.5 

 AMTRAC3 

 WACCM 

 UMUKCA−METO 

 MRI 

 SOCOL 

 NIWA−SOCOL 

 CCSRNIES 

 EMAC 

 ULAQ 

 CMAM 

 LMDZrepro 

 UMSLIMCAT 

 GEOSCCM 

 CNRM−ACM 

Figure 2-29. Chemical ozone 

depletion (Dobson units) in 

the polar vortex from January 

through April (top panel) and 

July through October (bottom 

panel) between 350 and 550 

K. Results from observations 

(black triangles) were derived 

from HALOE on UARS for the 

polar vortex core. Model results are shown in different colors and calculated for equivalent latitudes poleward of 80°. 

Model descriptions are given in 

Table 3-1 of Chapter 3. (Reproduction of Figure 6.36, Chapter 

6 of SPARC CCMVal, 2010).
2.58

Chapter 2

source gases, some models included ~6 pptv extra bromine 

to be more consistent with the actual stratospheric loading. 

This increases the loss computed for these models.

2.4.5.5 Trends and Recovery in the Upper

Troposphere and Lower Stratosphere

The upper troposphere and lower stratosphere 

(UTLS) are key regions where ozone changes can affect 

radiative forcing, and hence climate. At middle and high 

latitudes, up to 30% of total ozone resides in the lower 

stratosphere, i.e., the region between the tropopause and 

100 hPa (Logan, 1999). WMO (2007) noted that apparent 

trends in this region were mostly the result of interannual 

variability and could not be ascribed to ODSs. It also concluded that changes in transport were a major contributor 

to the increase in midlatitude total ozone since the mid1990s, particularly in the lowermost stratosphere (tropopause to 380K). Yang et al. (2006) showed that abundance of partial column ozone above 18 km altitude has 

risen slightly since 1997 in response to the decline in the 

atmospheric abundance of ODSs, whereas partial column 

ozone between the tropopause and 18 km rose much faster 

since 1997 than could be attributed to the decline in ODSs. 

The increases in total column were shown to differ significantly from the recovery in ozone expected solely from 

declining levels of ODS.

Vyushin et al. (2007), Mäder et al. (2007), and 

Harris et al. (2008) highlight the importance of accounting for dynamical variability and also volcanic aerosol 

in statistical models explaining observed changes in total ozone. Brunner et al. (2006) draw similar conclusions 

from analysis of vertically resolved ozone trends for the 

period 1979–2004 obtained by multiple linear regression 

analysis of the CATO ozone data set (see also WMO, 

2007). Following the approach of Newchurch et al. 

(2003), Brunner et al. (2006) fitted their regression model 

to the reference period 1979–1995 during which EESC 

was increasing, but analyzed the residuals for the period 

1996–2004 using different sets of explanatory variables. 

Their results show that a combination of EESC, QBO, and 

solar forcing is not sufficient to explain the recent ozone 

trend, as significantly positive residuals remained in the 

extratropical lower stratosphere for the period 1996–2004. 

However, most of the variations since 1996 could be explained when also accounting for volcanic aerosol and the 

Eliassen Palm (EP)-flux as a proxy for the wave forcing of 

the stratospheric circulation. This result demonstrates the 

importance of interannual variations in dynamics in producing the recent variations in ozone in the lower stratosphere. Brunner et al. (2006) also replaced EESC by a 

linear trend, assuming a continuation of the negative trend 

between 1979 and 1995. This resulted in an increase of 

the residuals that is much stronger than that resulting from 

omission of aerosols or EP-flux, demonstrating the importance of the Montreal Protocol in addition to natural variability for explaining the recent observed ozone changes.

The main driver for the natural variations in UTLS 

ozone is climate variability, as has been extensively discussed in the WMO (2007). This includes variability in 

the stratospheric mean meridional circulation (e.g., Fusco 

and Salby, 1999) as well as variability in the tropospheric circulation through synoptic scale wave activity (e.g., 

Weiss et al., 2001). Recent studies explore the impact 

of climate variability on vertically resolved ozone in the 

UTLS. Thouret et al. (2006) used ozone measurements 

from the MOZAIC program to show that ozone in the 

upper troposphere is positively correlated with the North 

Atlantic Oscillation (NAO) index (the dominant mode of 

variability in surface pressure in the North Atlantic) while 

ozone in the lower stratosphere is anticorrelated with the 

extremes of the Northern Annular Mode (NAM) index (extension of the NAO to the full hemisphere) calculated at 

150 hPa. Hess and Lamarque (2007), using the MOZART 

CTM driven by NCEP meteorological analyses, found that 

ozone is reduced throughout the troposphere at latitudes 

higher than 50°N during the positive phase of the Arctic 

Oscillation (AO, the surface manifestation of the NAM). 

This decrease can be linked to a decrease in ozone transported from the stratosphere. The El Niño-Southern Oscillation (ENSO) may also influence UTLS ozone. Fischer et 

al. (2008), using three different CCMs, found that ENSO 

forces a strengthening of the Brewer-Dobson circulation, 

which leads to lower total ozone in the tropics and higher 

total ozone at higher latitudes. While their study does not 

discuss the vertically resolved impact of the ENSO, the 

change in the Brewer-Dobson circulation is likely to have 

an impact on ozone in the lower stratosphere.

The question remains concerning how much the underlying long-term climate change may have contributed 

to the observed ozone changes. Hegglin and Shepherd 

(2009) use a stratosphere-resolving chemistry-climate 

model (CMAM) to study the effect of climate change on 

the stratospheric ozone distribution and derive expected 

trends of −1%, −3%, and +1% per decade in the SH lower 

stratosphere, the tropical lower stratosphere, and the NH 

lower stratosphere, respectively. They assumed a linear 

ozone response to climate change between 1960 and 2100. 

These numbers are consistent with those derived by Li et 

al. (2009) and are also confirmed in a multi-model comparison by Gettelman et al. (2010) except for the SH lower 

stratosphere, where the results were statistically not significant. These CCM results suggest that climate change 

may have masked around 30% of the ozone decrease due 

to EESC in the NH.

The evaluation of time series of ozone measurements at 200–100 hPa between 1960 and 2010 shows that 

the increase in ozone due to climate change is strongly 
2.59

Stratospheric Ozone and Surface UV

modulated by ozone depletion and recovery, with a larger signature in the SH than in the NH (Gettelman et al., 

2010). This modulation is characterized by strong decreases in lower stratospheric ozone up to around the year 

2000, followed by a strong increase, qualitatively agreeing with observations in this region (compare to Figure 

2-7). However, the decline of lower stratospheric ozone 

in the tropics has not yet been shown unambiguously as 

stated earlier. WMO (2003, 2007) further pointed out the 

importance of decadal scale variability in dynamics that 

modulates the underlying slow dynamical contribution 

from climate change. The degree to which the decrease 

and increase are driven by such decadal scale variability 

is not known.

The changes in lower stratospheric ozone are likely 

to affect upper tropospheric (UT) ozone. Mid- and upper 

tropospheric ozone trends exhibit strong regional differences and are not yet well understood (Staehelin et al., 

2001). CTMs with realistic ozone precursor emissions 

but without realistic stratospheric chemistry fail to reproduce either tropospheric background levels (Parrish et al., 

2009) or observed trends (Fiore et al., 2009). In addition, 

Ordoñez et al. (2007) present observational evidence of a 

significant correlation between ozone at 150 hPa and tropospheric background ozone at two European mountain 

sites over the period 1992–2004, indicating local coupling 

of these air masses through stratosphere-to-troposphere 

exchange. At northern high latitudes, stratospheric ozone 

is estimated to contribute up to 40% to ozone at 500 hPa 

and up to 30% at northern midlatitudes (Terao et al., 

2008). Hsu and Prather (2009) show the strong modulation of stratosphere-troposphere exchange (STE) ozone 

fluxes by the QBO, also pointing to stratospheric control 

of UT ozone.

It is likely that changes in the STE ozone fluxes 

induced by climate change need to be taken into account 

to understand observed UT ozone changes. Hegglin and 

Shepherd (2009) using the CMAM find that the simulated stratospheric ozone flux into the troposphere changes 

over the time period between 1960 and 2000 due to effects 

of both ozone depletion/recovery and climate change. 

Changes in the NH stratospheric ozone flux are around 

6%, while changes in the SH are about −8%. A multimodel comparison of STE ozone fluxes performed during CCMVal-2 agrees with the global mean predictions 

(Chapter 10 of SPARC CCMVal, 2010). While the predicted STE ozone fluxes seem to be biased high by around 

20% when compared to observations, the models’ range 

of STE ozone flux is much smaller than found in previous studies used for the IPCC (2007). The impact of STE 

ozone flux on tropospheric chemistry is discussed further 

in Section 4.4.4 of Chapter 4.

In summary, evaluations of observed ozone changes show evidence of a statistically significant positive 

change in ozone in the lower stratosphere; a part of this 

change may be attributed to decreasing ODSs. Dynamical 

changes are still mainly responsible for these changes, at 

least in the NH. CCMs simulate increases in lower stratospheric ozone and suggest that up to 30% of the ozone 

depletion was masked by the dynamical changes due to 

climate change. While the importance of lower stratospheric ozone changes on upper tropospheric ozone has 

been recognized, understanding of upper tropospheric 

ozone trends remains incomplete.

REFERENCES

Adriani, A., P. Massoli, G. Di Donfrancesco, F. Cairo, 

M.L. Moriconi, and M. Snels, Climatology of polar stratospheric clouds based on lidar observations from 1993 to 2001 over McMurdo Station, 

Antarctica, J. Geophys. Res., 109, D24211, doi: 

10.1029/2004JD004800, 2004.

Andersen, S.B., and B.M. Knudsen, The influence of polar 

vortex ozone depletion on NH mid-latitude ozone 

trends in spring, Atmos. Chem. Phys., 6, 2837-

2845, 2006.

Andreae, M.O., and A. Gelencsér, Black carbon or brown 

carbon? The nature of light-absorbing carbonaceous aerosols, Atmos. Chem. Phys., 6, 3131-3148, 

doi: 10.5194/acp-6-3131-2006, 2006.

Angell, J.K., and M. Free, Ground-based observations of 

the slowdown in ozone decline and onset of ozone 

increase, J. Geophys. Res., 114, D07303, doi: 

10.1029/2008JD010860, 2009.

Antón, M., D. Loyola, M. López, J.M. Vilaplana, M. 

Bañón, W. Zimmer, and A. Serrano, Comparison 

of GOME-2/MetOp total ozone data with Brewer 

spectroradiometer data over the Iberian Peninsula, 

Ann. Geophys., 27, 1377-1386, 2009.

Arola, A., S. Kazadzis, N. Krotkov, A. Bais, J. Gröbner, 

and J.R. Herman, Assessment of TOMS UV bias 

due to absorbing aerosols, J. Geophys. Res., 110, 

D23211, doi: 10.1029/2005JD005913, 2005.

Arola, A., S. Kazadzis, A. Lindfors, N. Krotkov, J. Kujanpää, J. Tamminen, A. Bais, A. di Sarra, J.M. 

Villaplana, C. Brogniez, A.M. Siani, J. Janouch, 

P. Weihs, A. Webb, T. Koskela, N. Kouremeti, 

D. Meloni, V. Buchard, F. Auriol, I. Ialongo, M. 

Staneck, S. Simic, A. Smedley, and S. Kinne, A 

new approach to correct for absorbing aerosols in 

OMI UV, Geophys. Res. Lett., 36, L22805, doi: 

10.1029/2009GL041137, 2009.

Atkinson, R., D.L. Baulch, R.A. Cox, J.N. Crowley, R.F. 

Hampson, R.G. Hynes, M.E. Jenkin, M.J. Rossi, 

and J. Troe, Evaluated kinetic and photochemical 

data for atmospheric chemistry: Volume III - gas 

phase reactions of inorganic halogens, Atmos. 
2.60

Chapter 2

Chem. Phys., 7 (4), 981-1191, 2007. (Cited as 

IUPAC, 2007.)

Austin, J., K. Tourpali, E. Rozanov, H. Akiyoshi, S. Bekki, G. Bodeker, C. Brühl, N. Butchart, M. Chipperfield, M. Deushi, V.I. Formichev, M.A. Giorgetta, 

L. Gray, K. Kodera, F. Lott, E. Manzini, D. Marsh, 

K. Matthes, T. Nagashima, K. Shibata, R.S. Stolarski, H. Struthers, and W. Tian, Coupled chemistry climate model simulations of the solar cycle 

in ozone and temperature, J. Geophys. Res., 113

(D11), doi: 10.1029/2007JD009391, 2008.

Avallone, L.M., and D.W. Toohey, Tests of halogen photochemistry using in situ measurements of ClO and 

BrO in the lower polar stratosphere, J. Geophys. 

Res., 106 (D10), 10411-10422, 2001.

Badosa, J., R.L. McKenzie, M. Kotkamp, J. Calbó, J.A. 

González, P.V. Johnston, M. O’Neill, and D.J. Anderson, Towards closure between measured and 

modelled UV under clear skies at four diverse sites, 

Atmos. Chem. Phys., 7 (11), 2817-2837, 2007.

Bais, A.F., and D. Lubin (Lead Authors), 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, and C.S. Zerefos, Surface ultraviolet radiation: Past, present, and future, Chapter 7 in Scientific Assessment of Ozone Depletion: 2006, Global 

Ozone Research and Monitoring Project–Report 

No. 50, 572 pp., World Meteorological Organization, Geneva, Switzerland, 2007.

Balis, D., J.-C. Lambert, M. Van Roozendael, R. Spurr, 

D. Loyola, Y. Livschitz, P. Valks, V. Amiridis, P. 

Gerard, J. Granville, and C. Zehner, Ten years of 

GOME/ERS2 total ozone data—the new GOME 

data processor (GDP) version 4: 2. Groundbased validation and comparisons with TOMS 

V7/V8, J. Geophys. Res., 112, D07307, doi: 

10.1029/2005JD006376, 2007.

Barnard, J.C., R. Volkamer, and E.I. Kassianov, Estimation of the mass absorption cross section of the organic carbon component of aerosols in the Mexico 

City Metropolitan Area, Atmos. Chem. Phys., 8

(22), 6665-6679, 2008.

Barret, B., M. De Mazière, and P. Demoulin, Retrieval and 

characterization of ozone profiles from solar infrared spectra at the Jungfraujoch, J. Geophys. Res., 

107, 4788, doi: 10.1029/2001JD001298, 2002.

Beig, G., and V. Singh, Trends in tropical tropospheric 

column ozone from satellite data and MOZART 

model, Geophys. Res. Lett., 34, L17801, doi: 

10.1029/2007GL030460, 2007.

Bergstrom, R.W., P. Pilewskie, P.B. Russell, J. Redemann, 

T.C. Bond, P.K. Quinn, and B. Sierau, Spectral absorption properties of atmospheric aerosols, Atmos. 

Chem. Phys., 7 (23), 5937-5943, 2007.

Bergstrom, R.W., K.S. Schmidt, O. Coddington, P. 

Pilewskie, H. Guan, J.M. Livingston, J. Redemann, 

and P.B. Russell, Aerosol spectral absorption in the 

Mexico City area: Results from airborne measurements during MILAGRO/INTEX B, Atmos. Chem. 

Phys., 10 (13), 6333-6343, doi: 10.5194/acp-10-

6333-2010, 2010.

Bernhard, G., C.R. Booth, and J.C. Ehramjian, Comparison of UV irradiance measurements at Summit, 

Greenland; Barrow, Alaska; and South Pole, Antarctica, Atmos. Chem. Phys., 8 (16), 4799-4810, 

doi: 10.5194/acp-8-4799-2008, 2008.

Bernhard, G., C.R. Booth, and J.C. Ehramjian, Climatology of ultraviolet radiation at high latitudes 

derived from measurements of the National Science Foundation’s ultraviolet spectral irradiance 

monitoring network, Chapter 3 in UV Radiation 

in Global Climate Change: Measurements, Modeling and Effects on Ecosystems, edited by W. 

Gao, D.L. Schmoldt, and J.R. Slusser, 550 pp., 

Springer-Verlag and Tsinghua University Press, 

ISBN: 978-3-642-03312-4, 48-72, 2010.

Bhartia, P.K., C.G. Wellemeyer, S.L. Taylor, N. Nath, 

and A. Gopolan, Solar Backscatter Ultraviolet 

(SBUV) version 8 profile algorithm, in Ozone 

Vol. I, Proceedings of the XX Quadrennial Ozone 

Symposium, 1-8 June 2004, Kos, Greece, edited by 

C.S. Zerefos, 295-296, International Ozone Commission, Athens, Greece, 2004.

Biele, J., A. Tsias, B.P. Luo, K.S. Carslaw, R. Neuber, G. 

Beyerle, and T. Peter, Nonequilibrium coexistence 

of solid and liquid particles in Arctic stratospheric 

clouds, J. Geophys. Res., 106 (D19), 22991-23007, 

doi: 10.1029/2001JD900188, 2001.

Blum, U., K.H. Fricke, K.P. Müller, J. Siebert, and G. 

Baumgarten, Long-term lidar observations of polar 

stratospheric clouds at Esrange in northern Sweden, 

Tellus, 57B, 412-422, 2005.

Bodeker, G.E., H. Shiona, and H. Struthers, The NIWA 

assimilated total column ozone data base and assimilated trace gas profile data base for validation 

of CCMs, paper presented at CCMVal 2005 Workshop, October 2005, Boulder, Colo., available: 

http://www.pa.op.dlr.de/workshops/CCMVal2005/

CCMVal2005_abstract.html, 2005.

Booth, C.R., and S. Madronich, Radiation amplification 

factors: Improved formulation accounts for large 

increases in ultraviolet radiation associated with 

Antarctic ozone depletion, in Ultraviolet Radiation 

in Antarctica: Measurement and Biological Effects, 

edited by C.S. Weiler and P.S. Penhale, AGU Antarct. Res. Ser., 62, 39-52, 1994.

Borkowski, J.L., Modelling UV radiation variations at 
2.61

Stratospheric Ozone and Surface UV

different time scales, Ann. Geophys., 26, 441-

446, 2008.

Brunner, D., J. Staehelin, J.A. Maeder, I. Wohltmann, and 

G.E. Bodeker, Variability and trends in total and 

vertically resolved stratospheric ozone based on the 

CATO ozone data set, Atmos. Chem. Phys., 6 (12), 

4985-5008, doi: 10.5194/acp-6-4985-2006, 2006.

Buchard, V., C. Brogniez, F. Auriol, B. Bonnel, J. Lenoble, A. Tanskanen, B. Bojkov, and P. Veefkind, 

Comparison of OMI ozone and UV irradiance data 

with ground-based measurements at two French 

sites, Atmos. Chem. Phys., 8 (16), 4517-4528, doi: 

10.5194/acp-8-4517-2008, 2008.

Burkholder, J.B., J.J. Orlando, and C.J. Howard, Ultravioletabsorption cross sections of Cl2O2 between 210 and 

410 nm, J. Phys. Chem., 94, 687-695, 1990.

Burrows, J.P., M. Weber, M. Buchwitz, V. Rozanov, A. 

Ladstädter-Weiβenmayer, A. Richter, R. de Beek, 

R. Hoogen, K. Bramstedt, K.-U. Eichmann, M. Eisinger, and D. Perner, The Global Ozone Monitoring Experiment (GOME): Mission concept and first 

scientific results, J. Atmos. Sci., 56, 151-175, 1999.

Butchart, N., and A.A. Scaife, Removal of chlorofluorocarbons by increased mass exchange between the 

stratosphere and troposphere in a changing climate, 

Nature, 410 (6830), 799-802, 2001.

Butchart, N., A.A. Scaife, M. Bourqui, J. de Grandpré, 

S.H.E. Hare, J. Kettleborough, U. Langematz, E. 

Manzini, F. Sassi, K. Shibata, D. Shindell, and M. 

Sigmond, Simulations of anthropogenic change 

in the strength of the Brewer-Dobson circulation, Clim. Dyn., 27 (7-8), 727-741, doi: 10.1007/

s00382-006-0162-4, 2006.

Butkovskaya, N., A. Kukui, and G. Le Bras, HNO3 forming channel of the HO2 + NO reaction as a function 

of pressure and temperature in the ranges of 72–600 

Torr and 223–323 K, J. Phys. Chem. A, 111, 9047-

9053, doi: 10.1021/jp074117m, 2007.

Butkovskaya, N., M.-T. Rayez, J.-C. Rayez, A. Kukui, 

and G. Le Bras, Water vapor effect on the HNO3

yield in the HO2 + NO reaction: Experimental and 

theoretical evidence, J. Phys. Chem. A, 113, 11327-

11342, doi: 10.1021/jp811428p, 2009.

Canty, T., E.D. Rivière, R.J. Salawitch, G. Berthet, 

J.-B. Renard, K. Pfeilsticker, M. Dorf, A. Butz, 

H. Bösch, R.M. Stimpfle, D.M. Wilmouth, E.C. 

Richard, D.W. Fahey, P.J. Popp, M.R. Schoeberl, L.R. Lait, and T.P. Bui, Nighttime OClO in 

the winter Arctic vortex, J. Geophys. Res., 110, 

D01301, doi: 10.1029/2004JD005035, 2005.

Cariolle, D., M.J. Evans, M.P. Chipperfield, N. Butkovskaya, A. Kukui, and G. Le Bras, Impact of the new 

HNO3-forming channel of the HO2 + NO reaction on tropospheric HNO3, NOx, HOx and ozone, 

Atmos. Chem. Phys., 8 (14), 4061-4068, doi: 

10.5194/acp-8-4061-2008, 2008.

Carl, S.A., A highly sensitive method for time-resolved 

detection of O(1D) applied to precise determination 

of absolute O(1D) reaction rate constants and O(3P) 

yields, Phys. Chem. Chem. Phys., 7, 4051-4053, 

doi: 10.1039/b513576c, 2005.

Carslaw, K.S., B.P. Luo, S.L. Clegg, Th. Peter, P. Brimblecombe, and P.J. Crutzen, Stratospheric aerosol 

growth and HNO3 gas phase depletion from coupled HNO3 and water uptake by liquid particles, 

Geophys. Res. Lett., 21 (23), 2479-2482, 1994.

Cede, A., J. Herman, A. Richter, N. Krotkov, and J. Burrows, Measurements of nitrogen dioxide total column amounts using a Brewer double spectrophotometer in direct sun mode, J. Geophys. Res., 111, 

D05304, doi: 10.1029/2005JD006585, 2006.

Chen, H.-Y., C.-Y. Lien, W.-Y. Lin, Y.T. Lee, and J.J. 

Lin, UV absorption cross sections of ClOOCl are 

consistent with ozone degradation models, Science, 

324, 781-784, doi: 10.1126/science.1171305, 2009.

Chipperfield, M.P., and V.E. Fioletov (Lead Authors), 

B. Bregman, J. Burrows, B.J. Connor, J.D. Haigh, 

N.R.P. Harris, A. Hauchecorne, L.L. Hood, S.R. 

Kawa, J.W. Krzyścin, J.A. Logan, N.J. Muthama, 

L. Polvani, W.J. Randel, T. Sasaki, J. Stähelin, R.S. 

Stolarski, L.W. Thomason, and J.M. Zawodny, 

Global ozone: Past and present, Chapter 3 in Scientific Assessment of Ozone Depletion: 2006, Global 

Ozone Research and Monitoring Project–Report 

No. 50, 572 pp., World Meteorological Organization, Geneva, Switzerland, 2007.

Chipperfield, M., W. Feng, and M. Rex, Arctic ozone loss 

and climate sensitivity: Updated three-dimensional 

model study, Geophys. Res. Lett., 32, L11813, doi: 

10.1029/2005GL022674, 2005.

Chubarova, N.Y., UV variability in Moscow according to 

the long-term UV measurements and reconstruction model, Atmos. Chem. Phys., 8 (12), 3025-3031, 

doi: 10.5194/acp-8-3025-2008, 2008.

Clain, G., J.L. Baray, R. Delmas, R. Diab, J. Leclair de 

Bellevue, P. Keckhut, F. Posny, J.M. Metzger, and 

J.P. Cammas, Tropospheric ozone climatology 

at two Southern Hemisphere tropical/subtropical 

sites, (Reunion Island and Irene, South Africa) from 

ozonesondes, LIDAR, and in situ aircraft measurements, Atmos. Chem. Phys., 9 (5), 1723-1734, doi: 

10.5194/acp-9-1723-2009, 2009.

Coldewey-Egbers, M., M. Weber, L.N. Lamsal, R. deBeck, M. Buchwitz, and J.P. Burrows, Total ozone 

retrieval from GOME UV spectral data using the 

weighting function DOAS approach, Atmos. Chem. 

Phys., 5 (4), 1015-1025, doi: 10.5194/acp-5-1015-

2005, 2005.
2.62

Chapter 2

Cordero, R.R., G. Seckmeyer, D. Pissulla, L. DaSilva, F. 

Labbe, Uncertainty evaluation of the spectral UV 

irradiance evaluated by using the UVSPEC Radiative Transfer Model, Opt. Commun., 276, 44-53, 

doi: 10.1016/j.optcom.2007.04.008, 2007.

Cordero, R.R., G. Seckmeyer, D. Pissulla, L. DaSilva, 

and F. Labbe, Uncertainty evaluation of spectral 

UV irradiance measurements, Meas. Sci. Technol., 19 (4), 045104, 1-15, doi: 10.1088/0957-

0233/19/4/045104, 2008a.

Cordero, R.R., G. Seckmeyer, D. Pissulla, and F. Labbe, 

Uncertainty of experimental integrals: Application 

to the UV index calculation, Metrologia, 45 (1), 

1-10, doi: 10.1088/0026-1394/45/1/001, 2008b.

Corr, C.A., N. Krotkov, S. Madronich, J.R. Slusser, B. 

Holben, W. Gao, J. Flynn, B. Lefer, and S.M. Kreidenweis, Retrieval of aerosol single scattering albedo at ultraviolet wavelengths at the T1 site during 

MILAGRO, Atmos. Chem. Phys., 9 (15), 5813-

5827, doi: 10.5194/acp-9-5813-2009, 2009.

COST 726 (European Cooperation in Science and Technology), Final report of COST action 726 – Long 

Term Changes and Climatology of UV Radiation 

over Europe, edited by Z. Lityńska, P. Koepke, 

H. De Backer, J. Gröbner, A. Schmalwieser, and 

L. Vuilleumier, COST Earth System Science and 

Environmental Management, Luxemburg: Office 

for Official Publications of the European Communities, 137 pp., available: http://www.cost726.org/, 

in press, 2010.

Cox, R.A., and G.D. Hayman, The stability and photochemistry of dimers of the ClO radical and implications for Antarctic ozone depletion, Nature, 332 

(6167), 796-800, doi: 10.1038/332796a0, 1988.

Crutzen, P.J., Albedo enhancement by stratospheric sulfur injections: A contribution to resolve a policy 

dilemma?, Clim. Change, 77 (3-4), 211-220, doi: 

10.1007/s10584-006-9101-y, 2006.

Curylo, A., J. Krzyścin, and B. Bogdańska, UV reconstruction with long time series global radiation 

measurements, in Proceedings of the UV Conference “One Century of UV Radiation Research,”

18-20 September 2007, Davos, Switzerland, edited 

by J. Gröbner, 211-212, 2007.

David, C., S. Bekki, S. Godin, G. Mégie, and M.P. Chipperfield, Polar stratospheric clouds climatology 

over Dumont d’Urville between 1989 and 1993 and 

the influence of volcanic aerosols on their formation, J. Geophys. Res., 103, 22163-22180, 1998.

David, C., P. Keckhut, A. Armetta, J. Jumelet, M. Snels, 

M. Marchand, and S. Bekki, Radiosonde stratospheric temperatures at Dumont d’Urville (Antarctica): Trends and link with polar stratospheric 

clouds, Atmos. Chem. Phys., 10 (8), 3813-3825, 

doi: 10.5194/acp-10-3813-2010, 2010.

DeMore, W.B., and E. Tschuikow-Roux, Ultraviolet spectrum and chemical reactivity of the chlorine monoxide dimer, J. Phys. Chem., 94 (15), 5856-5860, 

doi: 10.1021/j100378a046, 1990.

den Outer, P.N., H. Slaper, and R.B. Tax, UV radiation in the Netherlands: Assessing long-term 

variability and trends in relation to ozone and 

clouds, J. Geophys. Res., 110, D02203, 1-11, doi: 

10.1029/2004JD004824, 2005.

den Outer, P.N., H. Slaper, J. Kaurola, A. Lindfors, A. 

Kazantzidis, A.F. Bais, U. Feister, J. Junk, M. Janouch, and W. Josefsson, Reconstructing erythemal ultraviolet radiation levels in Europe for the 

past 4 decades, J. Geophys. Res., 115, D10102, doi: 

10.1029/2009JD012827, 2010.

Deshler, T., B.J. Johnson, and W.R. Rozier, Changes in 

the character of polar stratospheric clouds over 

Antarctica in 1992 due to the Pinatubo volcanic 

aerosol, Geophys. Res. Lett., 21 (4), 273-276, 1994.

Deshler, T., J.L. Mercer, H.G.J. Smit, R. Stubi, G. Levrat, 

B.J. Johnson, S.J. Oltmans, R. Kivi, A.M. Thompson, J. Witte, J. Davies, F.J. Schmidlin, G. Brothers, and T. Sasaki, Atmospheric comparison of 

electrochemical cell ozonesondes from different 

manufacturers, and with different cathode solution 

strengths: The Balloon Experiment on Standards 

for Ozonesondes, J. Geophys. Res., 113, D04307, 

doi: 10.1029/2007JD008975, 2008.

Dessler, A.E., D.B. Considine, J.E. Rosenfield, S.R. 

Kawa, A.R. Douglass, and J.M. Russell III, Lower 

stratospheric chlorine partitioning during the decay 

of the Mt. Pinatubo aerosol cloud, Geophys. Res. 

Lett., 24 (13), 1623-1626, 1997.

Dhomse, S., M. Weber, I. Wohltmann, M. Rex, and J.P. 

Burrows, On the possible causes of recent increases in NH total ozone from a statistical analysis of 

satellite data from 1979 to 2003, Atmos. Chem. 

Phys., 6 (5), 1165-1180, doi: 10.5194/acp-6-1165-

2006, 2006.

Diaz, S., C. Camilión, G. Deferrari, H. Fuenzalida, R. 

Armstrong, C. Booth, A. Paladini, S. Cabrera, C. 

Casiccia, C. Lovengreen, J. Pedroni, A. Rosales, 

H. Zagarese, and M. Vernet, Ozone and UV radiation over southern South America: Climatology and 

anomalies, Photochem. Photobiol., 82 (4), 834-843, 

doi: 10.1562/2005-09-26-RA-697, 2006.

Dillon, T.J., A. Horowitz, and J.N. Crowley, The atmospheric chemistry of sulphuryl fluoride, SO2F2, Atmos. Chem. Phys., 8 (6), 1547-1557, doi: 10.5194/

acp-8-1547-2008, 2008.

Drdla, K., B.W. Gandrud, D. Baumgardner, J.C. Wilson, 

T.P Bui, D. Hurst, S.M. Schauffler, H. Jost, J.B. 

Greenblatt, and C.R. Webster, Evidence for the 
2.63

Stratospheric Ozone and Surface UV

widespread presence of liquid-phase particles during the 1999-2000 Arctic winter, J. Geophys. Res., 

108 (D5), 8318, doi: 10.1029/2001JD001127, 2003.

Dunlea, E.J., and A.R. Ravishankara, Measurement of the 

rate coefficient for the reaction of O(1D) with H2O 

and re-evaluation of the atmospheric OH production rate, Phys. Chem. Chem. Phys., 6, 3333-3340, 

doi: 10.1039/b402483d, 2004.

Eckermann, S.D., L. Hoffmann, M. Höpfner, D.L. Wu, 

and M.J. Alexander, Antarctic NAT PSC belt of 

June 2003: Observational validation of the mountain wave seeding hypothesis, Geophys. Res. Lett., 

36, L02807, doi: 10.1029/2008GL036629, 2009.

Engel, A., T. Möbius, H. Bönisch, U. Schmidt, R. Heinz, 

I. Levin, E. Atlas, S. Aoki, T. Nakazawa, S. Sugawara, F. Moore, D. Hurst, J. Elkins, S. Schauffler, 

A. Andrews, and K. Berine, Age of stratospheric 

air unchanged within uncertainties over the past 

30 years, Nature Geosci., 2, 28-31, doi: 10.1038/

GEO388, 2009.

Eskes, H.J., R.J. van der A, E.J. Brinksma, J.P. Veefkind, 

J.F. de Haan, and P.J.M. Valks, Retrieval and validation of ozone columns derived from measurements of SCIAMACHY on Envisat, Atmos. Chem. 

Phys. Discuss., 5, 4429- 4475, doi: 10.5194/acpd-5-

4429-2005, 2005.

Eyring, V., N. Butchart, D.W. Waugh, H. Akiyoshi, J. 

Austin, S. Bekki, G.E. Bodeker, B.A. Boville, C. 

Brühl, M.P. Chipperfield, E. Cordero, M. Dameris, 

M. Deushi, V.E. Fioletov, S.M. Frith, R.R. Garcia, 

A. Gettelman, M.A. Giorgetta, V. Grewe, L. Jourdain, D.E. Kinnison, E. Mancini, E. Manzini, M. 

Marchand, D.R. Marsh, T. Nagashima, P.A. Newman, J.E. Nielsen, S. Pawson, G. Pitari, D.A. Plummer, E. Rozanov, M. Schraner, T.G. Shepherd, K. 

Shibata, R.S. Stolarski, H. Struthers, W. Tian, and 

M. Yoshiki, Assessment of temperature, trace species, and ozone in chemistry-climate model simulations of the recent past, J. Geophys. Res., 111, 

D22308, doi: 10.1029/2006JD007327, 2006.

Fahey, D.W., S.R. Kawa, E.L. Woodbridge, P. Tin, J.C. 

Wilson, H.H. Jonsson, J.E. Dye, D. Baumgardner, 

S. Borrmann, D.W. Toohey, L.M. Avallone, M.H. 

Proffitt, J. Margitan, M. Loewenstein, J.R. Podolske, R.J. Salawitch, S.C. Wofsy, M.K.W. Ko, D.E. 

Anderson, M.R. Schoeberl, and K.R. Chan, In situ 

measurements constraining the role of sulfate aerosols in midlatitude ozone depletion, Nature, 363

(6429), 509-514, doi: 10.1038/363509a0, 1993.

Fahey, D.W., R.S. Gao, K.S. Carslaw, J. Kettleborough, 

P.J. Popp, M.J. Northway, J.C. Holecek, S.C. 

Ciciora, R.J. McLaughlin, T.L. Thompson, R.H. 

Winkler, D.G. Baumgardner, B. Gandrud, P.O. 

Wennberg, S. Dhaniyala, K. McKinney, Th. Peter, 

R.J. Salawitch, T.P. Bui, J.W. Elkins, C.R. Webster, E.L. Atlas, H. Jost, J.C. Wilson, R.L. Herman, 

A. Kleinböhl, and M. von König, The detection of 

large HNO3-containing particles in the winter Arctic stratosphere, Science, 291, 1026-1031, 2001.

Farman, J.C., B.G. Gardiner, and J.D. Shanklin, Large 

losses of total ozone in Antarctica reveal seasonal 

ClOx/NOx interaction, Nature, 315, 207-210, 1985.

Feister, U., J. Junk, M. Woldt, A. Bais, A. Helbig, M. Janouch, W. Josefsson, A. Kazantzidis, A. Lindfors, 

P.N. den Outer, and H. Slaper, Long-term solar UV 

radiation reconstructed by ANN modelling with 

emphasis on spatial characteristics of input data, Atmos. Chem. Phys., 8 (12), 3107-3118, doi: 10.5194/

acp-8-3107-2008, 2008.

Felton, M.A. Jr., T.A. Kovacs, A.H. Omar, and C.A. 

Hostetler, Classification of Polar Stratospheric 

Clouds using LIDAR Measurements from the SAGE 

III Ozone Loss and Validation Experiment, Tech. 

Rep. ARL-TR-4154, 36 pp., U.S. Army Research 

Laboratory, Adelphi, Md., 2007.

Feng, W., M.P. Chipperfield, M. Dorf, K. Pfeilsticker, 

and P. Ricaud, Mid-latitude ozone changes: Studies 

with a 3-D CTM forced by ERA-40 analyses, Atmos. Chem. Phys., 7 (9), 2357-2369, doi: 10.5194/

acp-7-2357-2007, 2007.

Fioletov, V.E., Estimating the 27-day and 11-year solar cycle variations in tropical upper stratospheric ozone, J. Geophys. Res., 114, D02302, doi: 

10.1029/2008JD010499, 2009.

Fioletov, V.E., L.J.B. McArthur, J.B. Kerr, and D.I. Wardle, Long-term variations of UV-B irradiance over 

Canada estimated from Brewer observations and derived from ozone and pyranometer measurements, 

J. Geophys. Res., 106 (D19), 23009-23027, 2001.

Fioletov, V.E., G.E. Bodeker, A.J. Miller, R.D. McPeters, 

and R. Stolarski, Global and zonal total ozone variations estimated from ground-based and satellite 

measurements: 1964–2000, J. Geophys. Res., 107

(D22), 4647, doi: 10.1029/2001JD001350, 2002.

Fioletov, V.E., G. Labow, R. Evans, E.W. Hare, U. 

Köhler, C.T. McElroy, K. Miyagawa, A. Redondas, V. Savastiouk, A.M. Shalamyansky, J. Staehelin, K. Vanicek, and M. Weber, Performance of 

ground-based total ozone network assessed using 

satellite data, J. Geophys. Res., 113, D14313, doi: 

10.1029/2008JD009809, 2008.

Fiore, A.M., F.J. Dentener, O. Wild, C. Cuvelier, M.G. 

Schultz, P. Hess, C. Textor, M. Schulz, R.M. Doherty, 

L.W. Horowitz, I.A. MacKenzie, M.G. Sanderson, 

D.T. Shindell, D.S. Stevenson, S. Szopa, R. Van 

Dingenen, G. Zeng, C. Atherton, D. Bergmann, I. 

Bey, G. Carmichael, W.J. Collins, B.N. Duncan, G. 

Faluvegi, G. Folberth, M. Gauss, S. Gong, D. Hau-
2.64

Chapter 2

glustaine, T. Holloway, I.S.A. Isaksen, D.J. Jacob, 

J.E. Jonson, J.W. Kaminski, T.J. Keating, A. Lupu, 

E. Marmer, V. Montanaro, R.J. Park, G. Pitari, K.J. 

Pringle, J.A. Pyle, S. Schroeder, M.G. Vivanco, P. 

Wind, G. Wojcik, S. Wu, and A. Zuber, Multimodel 

estimates of intercontinental source-receptor relationships for ozone pollution, J. Geophys. Res., 114, 

D04301, doi: 10.1029/2008JD010816, 2009.

Fischer, A.M., D.T. Shindell, B. Winter, M.S. Bourqui, 

G. Faluvegi, E. Rozanov, M. Schraner, and S. 

Brönnimann, Stratospheric winter climate response to ENSO in three chemistry-climate models, Geophys. Res. Lett., 35, L13819, doi: 10.1029/ 

2008GL034289, 2008.

Fleming, E.L., C.H. Jackman, D.K. Weisenstein, and 

M.K.W. Ko, The impact of interannual variability on multidecadal total ozone simulations, J. 

Geophys. Res., 112, D10310, doi: 10.1029/2006 

JD007953, 2007.

Flynn, L.E., (editor), Solar Backscatter Ultraviolet Instrument (SBUV/2) Version 8 Ozone Retrieval Algorithm Theoretical Basis Document (V8 ATBD), 

NOAA National Environmental, Satellite, Data, 

and Information Service, Center for Satellite Applications and Research, available: http://www.star.

nesdis.noaa.gov/smcd/spb/calibration/icvs/sbuv/

doc/SBUV2_V8_ATBD_020207.pdf, 2007.

Frieler, K., M. Rex, R.J. Salawitch, T. Canty, M. Streibel, 

R.M. Stimpfle, K. Pfeilsticker, M. Dorf, D.K. 

Weisenstein, and S. Godin-Beekmann, Toward a 

better quantitative understanding of polar stratospheric ozone loss, Geophys. Res. Lett., 33, L10812, 

doi: 10.1029/2005GL025466, 2006.

Frith, S., R. Stolarski, and P.K. Bhartia, Implications of 

Version 8 TOMS and SBUV data for long-term 

trend analysis, in Ozone Vol. I, Proceedings of the 

XX Quadrennial Ozone Symposium, 1-8 June 2004, 

Kos, Greece, edited by C.S. Zerefos, 65-66, International Ozone Commission, Athens, Greece, 2004.

Fromm, M., J. Alfred, and M. Pitts, A unified, long-term, 

high-latitude stratospheric aerosol and cloud database using SAM II, SAGE II, and POAM II/III 

data: Algorithm description, database definition, 

and climatology, J. Geophys. Res., 108 (D12), 

4366, doi: 10.1029/2002JD002772, 2003.

Fusco, A.C., and M.L. Salby, Interannual variations of 

total ozone and their relationship to variations of 

planetary wave activity, J. Clim., 12 (6), 1619-

1629, 1999.

Gadhavi, H., R.T. Pinker, and I. Laszlo, Estimates of surface ultraviolet radiation over north America using 

Geostationary Operational Environmental Satellites observations, J. Geophys. Res., 113, D21205, 

doi: 10.1029/2007JD009308, 2008.

Gardiner, T., A. Forbes, M. de Mazière, C. Vigouroux, E. 

Mahieu, P. Demoulin, V. Velazco, J. Notholt, T. 

Blumenstock, F. Hase, I. Kramer, R. Sussmann, W. 

Stremme, J. Mellqvist, A. Strandberg, K. Ellingsen, 

and M. Gauss, Trend analysis of greenhouse gases 

over Europe measured by a network of groundbased remote FTIR instruments, Atmos. Chem. 

Phys., 8 (22), 6719-6727, doi: 10.5194/acp-8-6719-

2008, 2008.

Gettelman, A., T. Birner, V. Eyring, H. Akiyoshi, S. 

Bekki, C. Brühl, M. Dameris, D.E. Kinnison, 

F. Lefevre, F. Lott, E. Mancini, G. Pitari, D.A. 

Plummer, E. Rozanov, K. Shibata, A. Stenke, H. 

Struthers, and W. Tian, The tropical tropopause 

layer 1960–2100, Atmos. Chem. Phys., 9 (5), 

1621-1637, doi: 10.5194/acp-9-1621-2009, 2009.

Gettelman, A., M.I. Hegglin, S.-W. Son, J. Kim, M. Fujiwara, T. Birner, S. Kremser, M. Rex, J.A. Añel, 

H. Akiyoshi, J. Austin, S. Bekki, P. Braesike, C. 

Brühl, N. Butchart, M. Chipperfield, M. Dameris, 

S. Dhomse, H. Garny, S.C. Hardiman, P. Jöckel, 

D.E. Kinnison, J.F. Lamarque, E. Mancini, M. 

Marchand, M. Michou, O. Morgenstern, S. Pawson, G. Pitari, D. Plummer, J.A. Pyle, E. Rozanov, 

J. Scinocca, T.G. Shepherd, K. Shibata, D. Smale, 

H. Teyssèdre, and W. Tian, Multimodel assessment of the upper troposphere and lower stratosphere: Tropics and trends, J. Geophys. Res., 115, 

D00M08, doi: 10.1029/2009JD013638, 2010.

Gray, L.J., J. Beer, M. Geller, J.D. Haigh, M. Lockwood, 

K. Matthes, U. Cubasch, D. Fleitmann, G. Harrison, L. Hood, J. Luterbacher, G.A. Meehl, D. 

Shindell, B. van Geel, and W. White, Solar influence on climate, Rev. Geophys., 48, RG4001, doi: 

10.1029/2009RG000282, 2010.

Gröbner, J., M. Blumthaler, S. Kazadzis, A. Bais, A. 

Webb, J. Schreder, G. Seckmeyer, and D. Rembges, Quality assurance of spectral solar UV measurements: Results from 25 UV monitoring sites 

in Europe, 2002 to 2004, Metrologia, 43, S66-S71, 

doi: 10.1088/0026-1394/43/2/S14, 2006.

Gröbner, J., G. Hülsen, L. Vuilleumier, M. Blumthaler, 

J.M. Vilaplana, D. Walker, and J.E. Gil, Report 

of the PMOD/WRC-COST Calibration and Intercomparison of Erythemal Radiometers, Davos, 

Switzerland, 28 July–23 August 2006, 108 pp., 

European Cooperation in Science and Technology 

(COST), available: ftp.pmodwrc.ch/pub/publications/PMOD_COST726_BBreport.pdf, 2007.

Hall, T.M., D.W. Waugh, K.A. Boering, and R.A. Plumb, 

Evaluation of transport in stratospheric models, J. 

Geophys. Res., 104 (D15), 18815-18839, 1999.

Harris, N.R.P., E. Kyrö, J. Staehelin, D. Brunner, S.-B. 

Andersen, S. Godin-Beekmann, S. Dhomse, P. 
2.65

Stratospheric Ozone and Surface UV

Hadjinicolaou, G. Hansen, I. Isaksen, A. Jrrar, A. 

Karpetchko, R. Kivi, B. Knudsen, P. Krizan, J. Lastovicka, J. Maeder, Y. Orsolini, J.A. Pyle, M. Rex, 

K. Vanicek, M. Weber, I. Wohltmann, P. Zanis, 

and C. Zerefos, Ozone trends at northern mid- and 

high latitudes – a European perspective, Ann. Geophys., 26, 1207-1220, 2008.

Hase, F., Inversion von Spurengasprofilen aus Hochaufgelösten Bodengebundenen FTIR-Messungen in 

Absorption, Wissenschaftliche Berichte Forschungszentrum Karlsruhe, FZKA 6512; Ph.D. thesis, 

156 pp., ISSN 0947-8620, available: http://wwwimk.fzk.de/asf/ftir/disshase.pdf, 2000.

Hassler, B., G.E. Bodeker, and M. Dameris, Technical 

Note: A new global database of trace gases and 

aerosols from multiple sources of high vertical resolution measurements, Atmos. Chem. Phys., 8 (17), 

5403-5421, doi: 10.5194/acp-8-5403-2008, 2008.

Hegglin, M.I., and T.G. Shepherd, Large climate-induced 

changes in stratosphere-to-troposphere ozone flux 

and ultraviolet index, Nature Geosci., 2, 687-691, 

doi: 10.1038/NGEO604, 2009.

Herman, J.R., Global increase in UV irradiance during 

the past 30 years (1979–2008) estimated from satellite data, J. Geophys. Res., 115, D04203, doi: 

10.1029/2009JD012219, 2010.

Herman, J.R., G. Labow, N.C. Hsu, and D. Larko, Changes 

in cloud cover derived from reflectivity time series 

using SeaWiFS, N7-TOMS, EP-TOMS, SBUV2, and OMI radiance data, J. Geophys. Res., 114, 

D01201, doi: 10.1029/2007JD009508, 2009.

Hess, P.G., and J.-F. Lamarque, Ozone source attribution 

and its modulation by the Arctic oscillation during 

the spring months, J. Geophys. Res., 112, D11303, 

doi: 10.1029/2006JD007557, 2007.

Hicke, J.A., J. Slusser, K. Lantz, and F.G. Pascual, Trends 

and interannual variability in surface UVB radiation over 8 to 11 years observed across the United States, J. Geophys. Res., 113, D21302, doi: 

10.1029/2008JD009826, 2008.

Hickson, K.M., L.F. Keyser, and S.P. Sander, Temperature dependence of the HO2 + ClO reaction. 2. Reaction kinetics using the discharge flow-resonance 

fluorescence technique, J. Phys. Chem. A, 111, 

8126-8138, doi: 10.1021/ jp0689464, 2007.

Hitchcock, P., T.G. Shepherd, and C. McLandress, Past 

and future conditions for polar stratospheric cloud 

formation simulated by the Canadian middle atmosphere model, Atmos. Chem. Phys., 9 (2), 483-495, 

doi: 10.5194/acp-9-483-2009, 2009.

Hofmann, D.J., and S. Solomon, Ozone destruction 

through heterogeneous chemistry following the 

eruption of El Chichόn, J. Geophys. Res., 94 (D4), 

5029-5041, 1989.

Hofmann, D.J., and S.J. Oltmans, Anomalous Antarctic 

ozone during 1992: Evidence for Pinatubo volcanic 

aerosol effects, J. Geophys. Res., 98 (D10), 18555-

18562, 1993.

Hofmann, D.J., B.J. Johnson, and S.J. Oltmans, Twentytwo years of ozonesonde measurements at the South 

Pole, Int. J. Remote Sens., 30 (15-16), 3995-4008, 

doi: 10.1080/01431160902821932, 2009a.

Hofmann, D., J. Barnes, M. O’Neill, M. Trudeau, and R. 

Neely, Increase in background stratospheric aerosol observed with lidar at Mauna Loa Observatory 

and Boulder, Colorado, Geophys. Res. Lett., 36, 

L15808, doi: 10.1029/2009GL039008, 2009b.

Holton, J.R., Meridional distribution of stratospheric 

trace constituents, J. Atmos. Sci., 43 (12), 1238-

1242, 1986.

Höpfner, M., B.P. Luo, P. Massoli, F. Cairo, R. Spang, M. 

Snels, G. Di Donfrancesco, G. Stiller, T. von Clarmann, H. Fischer, and U. Biermann, Spectroscopic 

evidence for NAT, STS, and ice in MIPAS infrared 

limb emission measurements of polar stratospheric 

clouds, Atmos. Chem. Phys., 6 (5), 1201-1219, doi: 

10.5194/acp-6-1201-2006, 2006a.

Höpfner, M., N. Larsen, R. Spang, B. Luo, J. Ma, S.H. 

Svendsen, S.D. Eckermann, B. Knudsen, P. Massoli, F. Cairo, G. Stiller, T. v. Clarmann, and H. 

Fischer, MIPAS detects Antarctic stratospheric belt 

of NAT PSCs caused by mountain waves, Atmos. 

Chem. Phys., 6 (5), 1221-1230, doi: 10.5194/acp-6-

1221-2006, 2006b.

Höpfner, M., M.C. Pitts, and L.R. Poole, Comparison 

between CALIPSO and MIPAS observations of 

polar stratospheric clouds, J. Geophys. Res., 114, 

D00H05, doi: 10.1029/2009JD012114, 2009.

Hsu, J., and M.J. Prather, Stratospheric variability and tropospheric ozone, J. Geophys. Res., 114, D06102, 

doi: 10.1029/2008JD010942, 2009.

Huder, K.J., and W.B. DeMore, Absorption cross-sections 

of the ClO dimer, J. Phys. Chem., 99, 3905-3908, 

1995.

Hudman, R.C., D.J. Jacob, S. Turquety, E.M. Leibensperger, L.T. Murray, S. Wu, A.B. Gilliland, M. 

Avery, T.H. Bertram, W. Brune, R.C. Cohen, J.E. 

Dibb, F.M. Flocke, A. Fried, J. Holloway, J.A. Neuman, R. Orville, A. Perring, X. Ren, G.W. Sachse, 

H.B. Singh, A. Swanson, and P.J. Wooldrige, Surface and lightning sources of nitrogen oxides over 

the United States: Magnitudes, chemical evolution, 

and outflow, J. Geophys. Res., 112, D12S05, doi: 

10.1029/2006JD007912, 2007.

Hülsen, G., J. Gröbner, A. Bais, M. Blumthaler, P. Disterhoft, B. Johnsen, K.O. Lantz, C. Meleti, J. Schreder, J.M. Vilaplana Guerrero, and L. Ylianttila, Intercomparison of erythemal broadband radiometers 
2.66

Chapter 2

calibrated by seven UV calibration facilities in 

Europe and the USA, Atmos. Chem. Phys., 8 (16), 

4865-4875, doi: 10.5194/acp-8-4865-2008, 2008.

Ialongo, I., G.R. Casale, and A.M. Siani, Comparison of 

total ozone and erythemal UV data from OMI with 

ground-based measurements at Rome station, Atmos. Chem. Phys., 8 (12), 3283-3289, doi: 10.5194/

acp-8-3283-2008, 2008.

Ialongo, I., V. Buchard, C. Brogniez, G.R. Casale, and 

A.M. Siani, Aerosol Single Scattering Albedo retrieval in the UV range: An application to OMI satellite validation, Atmos. Chem. Phys., 10 (2), 331-

340, doi: 10.5194/acp-10-331-2010, 2010.

Innis, J.L., and A.R. Klekociuk, Planetary wave and 

gravity wave influence on the occurrence of polar stratospheric clouds over Davis Station, Antarctica, seen in lidar and radiosonde observations, 

J. Geophys. Res., 111, D22102, doi: 10.1029/ 

2006JD007629, 2006.

IPCC (Intergovernmental Panel on Climate Change), Climate Change 2007: Synthesis Report, Contribution 

of Working Groups I, II and III to the Fourth Assessment Report of the Intergovernmental Panel 

on Climate Change, edited by Core Writing Team, 

R.K. Pachauri, and A. Reisinger, IPCC, Geneva, 

Switzerland, 104 pp., 2007.

IUPAC (International Union of Pure and Applied Chemistry), 2007: see Atkinson et al., 2007.

Janjai, S., S. Buntung, R. Wattan, and I. Masiri, Mapping 

solar ultraviolet radiation from satellite data in a 

tropical environment, Remote Sens. Environ., 114 

(3), 682-691, doi: 10.1016/j.rse.2009.11.008, 2010.

Janouch, M., and L. Metelka, Modeling UV spectra with 

help of neural networks, in Proceedings of the UV 

Conference “One Century of UV Radiation Research,” 18-20 September 2007, Davos, Switzerland, edited by J. Gröbner, 207-208, 2007.

Jin, B., I.-C. Chen, W.-T. Huang, C.-Y. Lien, N. Guchhait, and J.J. Lin, Photodissociation cross section of 

ClOOCl at 330 nm, J. Phys. Chem. A., 114, 4791-

4797, doi: 10.1021/jp909374k, 2010.

JMA (Japan Meteorological Agency), Climate Change 

Monitoring Report 2008, 87 pp., Japan Meteorological Agency, Tokyo, available: http://ds.data.jma.

go.jp/tcc/tcc/products/gwp/CCMR2008.pdf, 2009.

Johnsen, B., B. Kjeldstad, T.N. Aalerud, L.T. Nilsen, J. 

Schreder, M. Blumthaler, G. Bernhard, C. Topaloglou, O. Meinander, A. Bagheri, J.R. Slusser, 

and J. Davis, Intercomparison and harmonization 

of UV Index measurements from multiband filter 

radiometers, J. Geophys. Res., 113, D15206, doi: 

10.1029/2007JD009731, 2008.

Jones, A., J. Urban, D.P. Murtagh, P. Eriksson, S. Brohede, C. Haley, D. Degenstein, A. Bourassa, C. 

von Savigny, T. Sonkaew, A. Rozanov, H. Bovensmann, and J. Burrows, Evolution of stratospheric 

ozone and water vapour time series studied with 

satellite measurements, Atmos. Chem. Phys., 9 (16), 

6055-6075, doi: 10.5194/acp-9-6055-2009, 2009.

JPL (Jet Propulsion Laboratory) 02-25: see Sander et al, 

2003.

JPL (Jet Propulsion Laboratory) 06-2: see Sander et al, 

2006.

JPL (Jet Propulsion Laboratory) 09-31: see Sander et al, 

2009.

Junk, J., U. Feister, and A. Helbig, Reconstruction of daily 

solar UV radiation from 1893 to 2002 in Potsdam, 

Germany, Int. J. Biometeorol., 51 (6), 505-512, doi: 

10.1007/s00484-007-0089-4, 2007.

Kawa, S.R., R.S. Stolarski, P.A. Newman, A.R. Douglass, 

M. Rex, D.J. Hofmann, M.L. Santee, and K. Frieler, Sensitivity of polar stratospheric ozone loss to 

uncertainties in chemical reaction kinetics, Atmos. 

Chem. Phys., 9 (22), 8651-8660, doi: 10.5194/acp9-8651-2009, 2009.

Kazadzis, S., A. Bais, V. Amiridis, D. Balis, C. Meleti, 

N. Kouremeti, C.S. Zerefos, S. Rapsomanikis, M. 

Petrakakis, A. Kelesis, P. Tzoumaka, and K. Kelektsoglou, Nine years of UV aerosol optical depth 

measurements at Thessaloniki, Greece, Atmos. 

Chem. Phys., 7 (8), 2091-2101, doi: 10.5194/acp7-2091-2007, 2007.

Kazadzis, S., A. Bais, A. Arola, N. Krotkov, N. Kouremeti, and C. Meleti, Ozone Monitoring Instrument 

spectral UV irradiance products: comparison with 

ground based measurements at an urban environment, Atmos. Chem. Phys., 9 (2), 585-594, doi: 

10.5194/acp-9-585-2009, 2009a.

Kazadzis, S., A. Bais, D. Balis, N. Kouremeti, M. Zempila, A. Arola, E. Giannakaki, V. Amiridis, and A. 

Kazantzidis, Spatial and temporal UV irradiance 

and aerosol variability within the area of an OMI 

satellite pixel, Atmos. Chem. Phys., 9 (14), 4593-

4601, doi: 10.5194/acp-9-4593-2009, 2009b.

Kazantzidis, A., A. Bais, K. Garane, S. Kazadzis, and C. 

Meleti, Estimation of UV irradiance from ancillary 

data and comparison with measurements at Thessaloniki, Greece (40.5°N, 23°E), in Remote Sensing 

of Clouds and the Atmosphere XI, edited by J.R. 

Slusser, K. Schäfer, and A. Comerón, Proc. SPIE, 

6362, 636228, doi: 10.1117/12.689813, 2006.

Kivi, R., E. Kyrö, T. Turunen, N.R.P. Harris, P. von der 

Gathen, M. Rex, S.B. Andersen, and I. Wohltmann, Ozonesonde observations in the Arctic 

during 1989–2003: Ozone variability and trends 

in the lower stratosphere and free troposphere, 

J. Geophys. Res., 112, D08306, doi: 10.1029/ 

2006JD007271, 2007.
2.67

Stratospheric Ozone and Surface UV

Kovalenko, L.J., K.W. Jucks, R.J. Salawitch, G.C. Toon, 

J.-F. Blavier, D.G. Johnson, A. Kleinböhl, N.J. Livesey, J.J. Margitan, H.M. Pickett, M.L. Santee, B. 

Sen, R.A. Stachnik, and J.W. Waters, Observed and 

modeled HOCl profiles in the midlatitude stratosphere: Implication for ozone loss, Geophys. Res. 

Lett., 34, L19801, doi: 10.1029/2007GL031100, 

2007.

Krzyścin, J.W., and J.L. Borkowski, Variability of the total ozone trend over Europe for the period 1950-

2004 derived from reconstructed data, Atmos. 

Chem. Phys., 8 (11), 2847-2857, doi: 10.5194/acp8-2847-2008, 2008.

Krzyścin, J.W., and B. Rajewska-Więch, Trends in the ozone 

vertical distribution from the Umkehr observations 

at Belsk 1963-2007, Int. J. Remote Sens., 30 (15), 

3917-3926, doi: 10.1080/014311609002821866, 

2009.

Krzyścin, J.W., P.S. Sobolewski, J. Jarosławski, J. Podgórski, and B. Rajewska-Więch, Erythemal UV observations at Belsk, Poland, in the period 1976-2008: 

Data homogenization, climatology, and trends, 

Acta Geophysica, 59 (1) 155-182, 2011.

Kuttippurath, J., S. Godin-Beekmann, F. Lefèvre, and 

A. Pazmiño, Ozone depletion in the Arctic winter 

2007–2008, Int. J. Remote Sens., 30 (15-16), 4071-

4082, doi: 10.1080/01431160902821965, 2009.

Kvalevåg, M.M., G. Myhre, and C.E. Lund Myhre, Extensive reduction of surface UV radiation since 1750 

in world’s populated regions, Atmos. Chem. Phys., 

9 (20), 7737-7751, doi: 10.5194/acp-9-7737-2009, 

2009.

Lee-Taylor, J., S. Madronich, C. Fischer, and B. Mayer, 

A climatology of UV radiation, 1979-2000, 65S65N, Chapter 1 in UV Radiation in Global Climate 

Change: Measurements, Modeling and Effects on 

Ecosystems, edited by W. Gao, D.L. Schmoldt, and 

J. Slusser, 550 pp., Springer-Verlag and Tsinghua 

University Press, ISBN: 978-3-642-03312-4, 1-20, 

2009.

Lerot, C., M. Van Roozendael, J. van Geffen, J. van Gent, 

C. Fayt, R. Spurr, G. Lichtenberg, and A. von Bargen, Six years of total ozone column measurements 

from SCIAMACHY nadir observations, Atmos. 

Meas. Tech., 2, 87-98, doi: 10.5194/amt-2-87-

2009, 2009.

Li, F., R.S. Stolarski, and P.A. Newman, Stratospheric 

ozone in the post-CFC era, Atmos. Chem. Phys., 

9 (6), 2207-2213, doi: 10.5194/acp-9-2207-2009, 

2009.

Lien, C.Y., W.Y. Lin, H.Y. Chen, W.T. Huang, B. Jin, I.C. 

Chen, and J.J. Lin, Photodissociation cross sections 

of ClOOCl at 248.4 and 266 nm, J. Chem. Phys.,

131 (17), 174301, 2009.

Lindfors, A., J. Kaurola, A. Arola, T. Koskela, K. Lakkala, W. Josefsson, J.A. Olseth, and B. Johnsen, 

A method for reconstruction of past UV radiation 

based on radiative transfer modeling: Applied to 

four stations in northern Europe, J. Geophys. Res., 

112, D23201, doi: 10.1029/2007JD008454, 2007.

Logan, J., An analysis of ozonesonde data for the lower 

stratosphere: Recommendations for testing models, 

J. Geophys. Res., 104 (D13), 16151-16170, 1999.

Lowe, D., and A.R. MacKenzie, Polar stratospheric cloud 

microphysics and chemistry, J. Atmos. Solar-Terr. 

Phys., 70, 13-40, doi: 10.1016/j.jastp.2007.09.011, 

2008.

Loyola, D.G., R.M. Coldewey-Egbers, M. Dameris, H. 

Garny, A. Stenke, M. Van Roozendael, C. Lerot, 

D. Balis, and M. Koukouli, Global long-term monitoring of the ozone layer — a prerequisite for predictions, Int. J. Remote Sens., 30, 4295-4318, doi: 

10.1080/01431160902825016, 2009a.

Loyola, D.G., R.M. Coldewey-Egbers, W. Zimmer, M. 

Koukouli, D. Balis, C. Lerot, M. Van Roozendael, and M. Dameris, Total ozone trends derived 

from the 14-years merged GOME/SCIAMACHY/

GOME-2 data record, Proc. ESA Atmospheric Science Conference, 7-11 September 2009, Barcelona, Spain, European Space Agency, ESA SP-676, 

available: http://uv-vis.aeronomie.be/publications/

proceedings/2009a_loyola.pdf, 2009b.

Mäder, J.A., J. Staehelin, D. Brunner, W.A. Stahel, I. 

Wohltmann, and T. Peter, Statistical modeling of 

total ozone: Selection of appropriate explanatory 

variables, J. Geophys. Res., 112, D11108, doi: 

10.1029/2006JD007694, 2007.

Marley, N.A., J.S. Gaffney, M. Tackett, N.C. Sturchio, L. 

Heraty, N. Martinez, K.D. Hardy, A. MarchanyRivera, T. Guilderson, A. MacMillan, and K. Steelman, The impact of biogenic carbon sources on 

aerosol absorption in Mexico City, Atmos. Chem. 

Phys., 9 (5), 1537-1549, doi: 10.5194/acp-9-1537-

2009, 2009.

Marsh, D.R., and R.R. Garcia, Attribution of decadal 

variability in lower-stratospheric tropical 

ozone, Geophys. Res. Lett., 34, L21807, doi: 

10.1029/2007GL030935, 2007.

Martins, J.V., P. Artaxo, Y.J. Kaufman, A.D. Castanho, 

and L.A. Remer, Spectral absorption properties of 

aerosol particles from 350–2500 nm, Geophys. Res. 

Lett., 36, L13810, doi: 10.1029/2009GL037435, 

2009.

Massoli, P., M. Maturilli, and R. Neuber, Climatology of Arctic polar stratospheric clouds as measured by lidar in Ny-Ålesund, Spitsbergen (79°N, 

12°E), J. Geophys. Res., 111, D09206, doi: 

10.1029/2005JD005840, 2006.
2.68

Chapter 2

Maturilli, M., R. Nueber, P. Massoli, F. Cairo, A. Adriani, M.L. Moriconi, and G. Di Donfrancesco, Differences in Arctic and Antarctic PSC occurrence 

as observed by lidar in Ny-Ålesund (79°N, 12°E) 

and McMurdo (78°S, 167°E), Atmos. Chem. Phys.,

5 (8), 2081-2090, doi: 10.5194/acp-5-2081-2005, 

2005.

Matus, M.H., M.T. Nguyen, D.A. Dixon, K.A. Peterson, 

and J.S. Francisco, ClClO2 is the most stable isomer 

of Cl2O2. Accurate coupled cluster energetics and 

electronic spectra of Cl2O2 isomers, J. Phys. Chem. 

A, 112, 9623-9627, doi: 10.1021/jp806220r, 2008.

McCormack, J.P., D.E. Siskind, and L.L. Hood, SolarQBO interaction and its impact on stratospheric 

ozone in a zonally averaged photochemical transport model of the middle atmosphere, J. Geophys. 

Res., 112, D16109, doi: 10.1029/2006JD008369, 

2007.

McDonald, A.J., S.E. George, and R.M. Woollands, Can 

gravity waves significantly impact PSC occurrence in the Antarctic?, Atmos. Chem. Phys., 9 (22), 

8825-8840, doi: 10.5194/acp-9-8825-2009, 2009.

McElroy, M.B., R.J. Salawitch, S.C. Wofsy, and J.A. Logan, Reductions of Antarctic ozone due to synergistic interactions of chlorine and bromine, Nature, 

321, 759-762, doi: 10.1038/321759a0, 1986.

McKenzie, R., G. Bodeker, G. Scott, J. Slusser, and K. 

Lantz, Geographical differences in erythemallyweighted UV measured at mid-latitude USDA 

sites, Photochem. Photobiol. Sci., 5 (3), 343-352, 

doi: 10.1039/B510943D, 2006.

McKenzie, R.L., C. Weinreis, P.V. Johnston, B. Liley, H. 

Shiona, M. Kotkamp, D. Smale, N. Takegawa, and 

Y. Kondo, Effects of urban pollution on UV spectral irradiances, Atmos. Chem. Phys., 8 (18), 5683-

5697, doi: 10.5194/acp-8-5683-2008, 2008.

McKenzie, R.L., J.B. Liley, and L.O. Björn, UV radiation: Balancing risks and benefits, Photochem. 

Photobiol., 85, 88-98, doi: 10.1111/j.1751-1097. 

2008.00400.x, 2009.

McLandress, C., and T.G. Shepherd, Simulated anthropogenic changes in the Brewer-Dobson circulation, including its extension to high latitudes, J. Clim., 22, 

1516-1540, doi: 10.1175/2008JCLI2679.1, 2009.

McLinden, C.A., S. Tegtmeier, and V. Fioletov, Technical 

note: A SAGE-corrected SBUV zonal-mean ozone 

data set, Atmos. Chem. Phys., 9 (20), 7963-7972, 

doi: 10.5194/acp-9-7963-2009, 2009.

Meinander, O., A. Kontu, K. Lakkala, A. Heikkilä, L. Ylianttila, and M. Toikka, Diurnal variations in the UV 

albedo of arctic snow, Atmos. Chem. Phys., 8 (21), 

6551-6563, doi: 10.5194/acp-8-6551-2008, 2008.

Miller, A.J., A. Cai, G. Taio, D.J. Wuebbles, L.E. Flynn, 

S.-K. Yang, E.C. Weatherhead, V. Fioletov, I. 

Petropavlovskikh, X.-L. Meng, S. Guillas, R.M. 

Nagatani, and G.C. Reinsel, Examination of ozonesonde data for trends and trend changes incorporating solar and Arctic oscillation signals, J. Geophys. 

Res., 111, D13305, doi: 10.1029/2005JD006684, 

2006.

Molina, L.T., and M.J. Molina, Production of Cl2O2 from 

the self-reaction of the ClO radical, J. Phys. Chem.,

91, 433-436, 1987.

Moore, T.A., M. Okumura, J.W. Seale, and T.K. Minton, 

UV photolysis of ClOOCl, J. Phys. Chem. A, 103

(12), 1691-1695, 1999.

Müller, R., J.-U. Grooß, C. Lemmen, D. Heinze, M. Dameris, and G. Bodeker, Simple measures of ozone 

depletion in the polar stratosphere, Atmos. Chem. 

Phys., 8 (2), 251-264, doi: 10.5194/acp-8-251-

2008, 2008.

Murata, I., K. Sato, S. Okano, and Y. Tomikawa, Measurements of stratospheric ozone with a balloon-borne 

optical ozone sensor, Int. J. Remote Sens., 30, 3961-

3966, doi: 10.1080/01431160902822823, 2009.

Newchurch, M.J., E.-S. Yang, D.M. Cunnold, G.C. Reinsel, J.M. Zawodny, and J.M. Russell III, Evidence 

for slowdown in stratospheric ozone loss: First 

stage of ozone recovery, J. Geophys. Res., 108 

(D16), 4507, doi: 10.1029/2003JD003471, 2003.

Newman, P.A., and M. Rex (Lead Authors), 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, and S. Yoden, 

Polar ozone: Past and present, Chapter 4 in Scientific Assessment of Ozone Depletion: 2006, Global 

Ozone Research and Monitoring Project–Report 

No. 50, 572 pp., World Meteorological Organization, Geneva, Switzerland, 2007.

Newman, P.A., E.R. Nash, S.R. Kawa, S.A. Montzka, and 

S.M. Schauffler, When will the Antarctic ozone 

hole recover?, Geophys. Res. Lett., 33, L12814, 

doi: 10.1029/2005GL025232, 2006.

Noel, V., A. Hertzog, H. Chepfer, and D.M. Winker, Polar stratospheric clouds over Antarctica from the 

CALIPSO spaceborne lidar, J. Geophys. Res., 113, 

D02205, doi: 10.1029/2007JD008616, 2008.

Noel, V., A. Hertzog, and H. Chepfer, CALIPSO observations of wave-induced PSCs with near-unity optical 

depth over Antarctica in 2006–2007, J. Geophys. 

Res., 114, D05202, doi: 10.1029/2008JD010604, 

2009.

Ordóñez, C., D. Brunner, J. Staehelin, P. Hadjinicolaou, 

J.A. Pyle, M. Jonas, H. Wernli, and A.S.H. Prévôt, 

Strong influence of lowermost stratospheric ozone 

on lower tropospheric background ozone changes 

over Europe, Geophys. Res. Lett., 34, L07805, doi: 
2.69

Stratospheric Ozone and Surface UV

10.1029/2006GL029113, 2007.

Panicker, A.S., G. Pandithurai, T. Takamura, and R.T. 

Pinker, Aerosol effects in the UV-B spectral region 

over Pune, an urban site in India, Geophys. Res. 

Lett., 36, L10802, doi: 10.1029/2009GL037632, 

2009.

Papanastasiou, D.K., V.C. Papadimitriou, D.W. Fahey, 

and J.B. Burkholder, UV absorption spectrum of 

the ClO dimer (Cl2O2) between 200 and 420 nm, 

J. Phys. Chem. A, 113, 13711-13726, doi: 10.1021/

jp9065345, 2009.

Parrington, M., D.B.A. Jones, K.W. Bowman, L.W. 

Horowitz, A.M. Thompson, D.W. Tarasick, and 

J.C. Witte, Estimating the summertime tropospheric ozone distribution over North America through 

assimilation of observations from the Tropospheric Emission Spectrometer, J. Geophys. Res., 113, 

D18307, doi: 10.1029/2007JD009341, 2008.

Parrish, D.D., D.B. Millet, and A.H. Goldstein, Increasing 

ozone in marine boundary layer inflow at the west 

coasts of North America and Europe, Atmos. Chem. 

Phys., 9 (4), 1303-1323, doi: 10.5194/acp-9-1303-

2009, 2009.

Permien, T., R. Vogt, and R.N. Schindler, Absorption 

spectra of HOCl and Cl2O2, in Mechanisms of Gas 

Phase-Liquid Phase Chemical Transformations, 

edited by R.A. Cox, Air Pollution Report #17, 149-

153, Environmental Research Program of the CEC, 

EUr 12035 EN, Brussels, Belgium, 1988.

Petkov, B., V. Vitale, C. Tomasi, U. Bonafé, S. Scaglione, 

D. Flori, R. Santaguida, M. Gausa, G. Hansen, 

and T. Colombo, Narrowband filter radiometer for 

ground-based measurements of global ultraviolet 

solar irradiance and total ozone, Appl. Opt., 45 (18), 

4383-4395, doi: 10.1364/AO.45.004383, 2006.

Petropavlovskikh, I., C. Ahn, P.K. Bhartia, and L.E. 

Flynn, Comparison and covalidation of ozone 

anomalies and variability observed in SBUV(/2) 

and Umkehr northern midlatitude ozone profile 

estimates, Geophys. Res. Lett., 32, L06805, doi: 

10.1029/2004GL022002, 2005a.

Petropavlovskikh, I., P.K. Bhartia, and J.J. DeLuisi, New 

Umkehr ozone profile retrieval algorithm optimized 

for climatological studies, Geophys. Res. Lett., 32, 

L16808, doi: 10.1029/2005GL023323, 2005b.

Pitts, M.C., L.W. Thomason, L.R. Poole, and D.M. Winker, Characterization of Polar Stratospheric Clouds 

with spaceborne lidar: CALIPSO and the 2006 Antarctic season, Atmos. Chem. Phys., 7 (19), 5207-

5228, doi: 10.5194/acp-7-5207-2007, 2007.

Pitts, M.C., L.R. Poole, and L.W. Thomason, CALIPSO 

polar stratospheric cloud observations: Secondgeneration detection algorithm and composition 

discrimination, Atmos. Chem. Phys., 9 (19), 7577-

7589, doi: 10.5194/acp-9-7577-2009, 2009.

Plenge, J., R. Flesch, S. Kühl, B. Vogel, R. Müller, F. 

Stroh, and E. Rühl, Ultraviolet photolysis of the ClO 

dimer, J. Phys. Chem. A, 108, 4859-4863, 2004.

Pommereau, J.P., and F. Goutail, O3 and NO2 groundbased measurements by visible spectrometry during Arctic winter and spring 1988, Geophys. Res. 

Lett., 15, (8) 891-894, 1988.

Poole, L.R., and M.C. Pitts, Polar stratospheric cloud climatology based on Stratospheric Aerosol Measurement II observations from 1978 to 1989, J. Geophys. Res., 99 (D6), 13083-13089, 1994.

Pope, F.D., J.C. Hansen, K.D. Bayes, R.R. Friedl, and S.P. 

Sander, Ultraviolet absorption spectrum of chlorine peroxide, ClOOCl, J. Phys. Chem., 111, 4322-

4332, 2007.

Portmann, R.W., S. Solomon, R.R. Garcia, L.W. Thomason, L.R. Poole, and M.P. McCormick, Role of 

aerosol variations in anthropogenic ozone depletion 

in the polar regions, J. Geophys. Res., 101 (D17), 

22991-23006, 1996.

Pougatchev, N.S., B.J. Connor, and C.P. Rinsland, Infrared measurements of the ozone vertical distribution above Kitt Peak, J. Geophys. Res., 100 (D8), 

16689-16697, 1995.

Pribullová, A., and M. Chmelík, Typical distribution of 

the solar erythemal UV radiation over Slovakia, Atmos. Chem. Phys., 8 (17), 5393-5401, doi: 10.5194/

acp-8-5393-2008, 2008.

Price, C., and D. Rind, A simple lightning parameterization for calculating global lightning distributions, J. 

Geophys. Res., 97 (D9), 9919-9933, 1992.

Randel, W.J., and F. Wu, A stratospheric ozone profile data 

set for 1979–2005: Variability, trends, and comparisons with column ozone data, J. Geophys. Res., 112, 

D06313, doi: 10.1029/2006JD007339, 2007.

Randel, W.J., K.P. Shine, J. Austin, J. Barnett, C. Claud, 

N.P. Gillett, P. Keckhut, U. Langematz, R. Lin, 

C. Long, C. Mears, A. Miller, J. Nash, D.J. Seidel, D.W.J. Thompson, F. Wu, and S. Yoden, 

An update of observed stratospheric temperature trends, J. Geophys. Res., 114, D02107, doi: 

10.1029/2008JD010421, 2009.

Rasch, P.J., P.J. Crutzen, and D.B. Coleman, Exploring the 

geoengineering of climate using stratospheric sulfate aerosols: The role of particle size, Geophys. Res. 

Lett., 35, L02809, doi: 10.1029/2007GL032179, 

2008.

Reinsel, G.C., E.C. Weatherhead, G.C. Tiao, A.J. Miller, 

R.M. Nagatani, D.J. Wuebbles, and L.E. Flynn, 

On detection of turnaround and recovery in trend 

for ozone, J. Geophys. Res., 107 (D10), 4078, doi: 

10.1029/2001JD000500, 2002.

Reinsel, G.C., A.J. Miller, E.C. Weatherhead, L.E. Flynn, 
2.70

Chapter 2

R. Nagatani, G.C. Tiao, and D.J. Wuebbles, Trend 

analysis of total ozone data for turnaround and 

dynamical contributions, J. Geophys. Res., 110, 

D16306, doi: 10.1029/2004JD004662, 2005.

Remsberg, E.E., On the response of Halogen Occultation Experiment (HALOE) stratospheric 

ozone and temperature to the 11-year solar cycle 

forcing, J. Geophys. Res., 113, D22304, doi: 

10.1029/2008JD010189, 2008.

Rex, M., R.J. Salawitch, P. von der Gathen, N.R.P. Harris, 

M.P. Chipperfield, and B. Naujokat, Arctic ozone 

loss and climate change, Geophys. Res. Lett., 31, 

L04116, doi: 10.1029/2003GL018844, 2004.

Rex, M., R.J. Salawitch, H. Deckelmann, P. von der Gathen, N.R.P. Harris, M.P. Chipperfield, B. Naujokat, 

E. Reimer, M. Allaart, S.B. Andersen, R. Bevilacqua, G.O. Braathen, H. Claude, J. Davies, H. De 

Backer, H. Dier, V. Dorokhov, H. Fast, M. Gerding, S. Godin-Beekmann, K. Hoppel, B. Johnson, 

E. Kyrö, Z. Litynska, D. Moore, H. Nakane, M.C. 

Parrondo, A.D. Risley Jr., P. Skrivankova, R. Stübi, P. Viatte, V. Yushkov, and C. Zerefos, Arctic 

winter 2005: Implications for stratospheric ozone 

loss and climate change, Geophys. Res. Lett., 33, 

L23808, doi: 10.1029/2006GL026731, 2006.

Rieder, H.E., F. Holawe, S. Simic, M. Blumthaler, J.W. 

Krzyścin, J.E. Wagner, A.W. Schmalwieser, and P. 

Weihs, Reconstruction of erythemal UV-doses for 

two stations in Austria: A comparison between alpine and urban regions, Atmos. Chem. Phys., 8 (20), 

6309-6323, doi: 10.5194/acp-8-6309-2008, 2008.

Robock, A., T. Adams, M. Moore, L. Oman, G. 

Stenchikov, Southern Hemisphere atmospheric 

circulation effects of the 1991 Mount Pinatubo 

eruption, Geophys. Res. Lett., 34, L23710, doi: 

10.1029/2007GL031403, 2007.

Rodgers, C.D., Inverse Methods for Atmospheric Sounding: Theory and Practice, 240 pp., ISBN 981-02-

2740X, World Scientific Publishing Co. Ltd., London, UK, 2000.

Rosenfield, J.E., S.M. Frith, and R.S. Stolarski, Version 8 SBUV ozone profile trends compared 

with trends from a zonally averaged chemical model, J. Geophys. Res., 110, D12302, doi: 

10.1029/2004JD005466, 2005.

Rösevall, J.D., D.P. Murthagh, and J. Urban, Ozone depletion in the 2006/2007 Arctic winter, Geophys. Res. 

Lett., 34, L21809, doi: 10.1029/2007GL030620, 

2007.

Ruckstuhl, C., R. Philipona, K. Behrens, M. Collaud 

Coen, B. Dürr, A. Heimo, C. Mätzler, S. Nyeki, 

A. Ohmura, L. Vuilleumier, M. Weller, C. Wehrli, and A. Zelenka, Aerosol and cloud effects on 

solar brightening and the recent rapid warming, 

Geophys. Res. Lett., 35, L12708, doi: 10.1029/ 

2008GL034228, 2008.

Salawitch, R.J., D.K. Weisenstein, L.J. Kovalenko, C.E. 

Sioris, P.O. Wennberg, K. Chance, M.K.W. Ko, and 

C.A. McLinden, Sensitivity of ozone to bromine in 

the lower stratosphere, Geophys. Res. Lett., 32 (5), 

L05811, doi: 10.1029/2004GL021504, 2005.

Sander, S.P., R.R. Friedl, D.M. Golden, M.J. Kurylo, R.E. 

Huie, V.L. Orkin, G.K. Moortgat, A.R. Ravishankara, C.E. Kolb, M.J. Molina, and B.J. FinlaysonPitts, Chemical Kinetics and Photochemical Data 

for Use in Atmospheric Studies: Evaluation No. 14, 

JPL Publication 02-25, Jet Propulsion Laboratory, 

Pasadena, Calif., 2003. (Cited as JPL 02-25.)

Sander, S.P., R.R. Friedl, D.M. Golden, M.J. Kurylo, G.K. 

Moortgat, H. Keller-Rudek, P.H. Wine, A.R. Ravishankara, C.E. Kolb, M.J. Molina, B.J. FinlaysonPitts, R.E. Huie, and V.L. Orkin, Chemical Kinetics 

and Photochemical Data for Use in Atmospheric 

Studies: Evaluation No. 15, JPL Publication 06-2, 

Jet Propulsion Laboratory, Pasadena, Calif., 2006. 

(Cited as JPL 06-2.)

Sander, S.P., J. Abbatt, J.R. Barker, J.B. Burkholder, R.R. 

Friedl, D.M. Golden, R.E. Huie, C.E. Kolb, M.J. 

Kurylo, G.K. Moorgat, V.L. Orkin, and P.H. Wine, 

Chemical Kinetics and Photochemical Data for 

Use in Atmospheric Studies: Evaluation Number 

16, JPL Publication 09-31, Jet Propulsion Laboratory, Pasadena, Calif., available at http://jpldataeval.jpl.nasa.gov, 2009. (Cited as JPL 09-31.)

Santee, M.L., I.A. MacKenzie, G.L. Manney, M.P. Chipperfield, P.F. Bernath, K.A. Walker, C.D. Boone, 

L. Froidevaux, N.J. Livesey, and J.W. Waters, A 

study of stratospheric chlorine partitioning based 

on new satellite measurements and modeling, 

J. Geophys. Res., 113, D12307, doi: 10.1029/ 

2007JD009057, 2008.

Santee, M.L., S.P. Sander, N.J. Livesey, and L. Froidevaux, 

Constraining the chlorine monoxide (ClO)/chlorine 

peroxide (ClOOCl) equilibrium constant from Aura 

Microwave Limb Sounder measurements of nighttime ClO, Proc. Nat. Acad. Sci., 107 (15), 6588-

6593, doi: 10.1073/pnas.0912659107, 2010.

Sauvage, B., R.V. Martin, A. van Donkelaar, X. Liu, K. 

Chance, L. Jaeglé, P.I. Palmer, S. Wu, and T.-M. 

Fu, Remote sensed and in situ constraints on processes affecting tropical tropospheric ozone, Atmos. Chem. Phys., 7 (3), 815-838, doi: 10.5194/

acp-7-815-2007, 2007.

Scarnato, B., J. Staehelin, T. Peter, J. Gröbner, and R. Stübi, Temperature and slant path effects in Dobson 

and Brewer total ozone measurements, J. Geophys. 

Res., 114, D24303, doi: 10.1029/2009JD012349, 

2009.
2.71

Stratospheric Ozone and Surface UV

Scarnato, B., J. Staehelin, R. Stübi, and H. Schill, Long 

term total ozone observations at Arosa (Switzerland) with Dobson and Brewer instruments 

(1988–2007), J. Geophys. Res., 115, D13306, doi: 

10.1029/2009JD011908, 2010.

Schallhart, B., M. Blumthaler, J. Schreder, and J. Verdebout, A method to generate near real time UV-Index maps of Austria, Atmos. Chem. Phys., 8 (24), 

7483-7491, doi: 10.5194/acp-8-7483-2008, 2008.

Schiermeier, Q., Chemists poke holes in ozone theory, Nature, 449, 382383, 2007.

Schmidt, H., and G.P. Brasseur, The response of the middle 

atmosphere to solar cycle forcing in the Hamburg 

Model of the neutral and ionized atmosphere, Space 

Sci. Rev., 125, 345-356, doi: 10.1007/s11214-006-

9068-z, 2006.

Schmidt, H., G.P. Brasseur, and M.A. Giorgetta, Solar 

cycle signal in a general circulation and chemistry model with internally generated quasi-biennial 

oscillation, J. Geophys. Res., 115, D00I14, doi: 

10.1029/2009JD012542, 2010.

Schneider, M., A. Redondas, F. Hase, C. Guirado, T. Blumenstock, and E. Cuevas, Comparison of groundbased Brewer and FTIR total O3 monitoring techniques, Atmos. Chem. Phys., 8 (18), 5535-5550, 

doi: 10.5194/acp-8-5535-2008, 2008.

Schofield, R., K. Frieler, I. Wohltmann, M. Rex, M. von 

Hobe, F. Stroh, G. Koch, T. Peter, T. Canty, R. Salawitch, and C.M. Volk, Polar stratospheric chlorine 

kinetics from a self-match flight during SOLVE-II/

EUPLEX, Geophys. Res. Lett., 35, L01807, doi: 

10.1029/2007GL031740, 2008.

Seckmeyer G., M. Glandorf, C. Wichers, R. McKenzie, D. 

Henriques, F. Carvalho, A. Webb, A.M. Siani, A. 

Bais, B. Kjeldstad, C. Brogniez, P. Werle, T. Koskela, K. Lakkala, J. Gröbner, H. Slaper, P. denOuter, and U. Feister, Europe’s darker atmosphere in 

the UV-B, Photochem. Photobiol. Sci., 7 (8), 925- 

930, 2008a.

Seckmeyer, G., D. Pissulla, M. Glandorf, D. Henriques, 

B. Johnsen, A. Webb, A.-M. Siani, A. Bais, B. 

Kjeldstad, C. Brogniez, J. Lenoble, B. Gardiner, 

P. Kirsch, T. Koskela, J. Kaurola, B. Uhlmann, 

H. Slaper, P. den Outer, M. Janouch, P. Werle, J. 

Gröbner, B. Mayer, A. de la Casiniere, S. Simic, 

and F. Carvalho, Variability of UV irradiance in 

Europe, Photochem. Photobiol., 84, 172-179, doi: 

10.1111/j.1751-1097.20007.00216.x, 2008b.

Shepherd, T.G., Dynamics, stratospheric ozone, and climate change, Atmos.-Ocean, 46 (1), 117-138, doi: 

10.3137/ao.460106, 2008.

Shepherd, T.G., and A.I. Jonsson, On the attribution of 

stratospheric ozone and temperature changes to 

changes in ozone-depleting substances and wellmixed greenhouse gases, Atmos. Chem. Phys., 8 (5), 

1435-1444, doi: 10.5194/acp-8-1435-2008, 2008.

Shindell, D.T., and R.L. de Zafra, Chlorine monoxide 

in the Antarctic spring vortex 2. A comparison of 

measured and modeled diurnal cycling over McMurdo Station, 1993, J. Geophys. Res., 101 (D1), 

1475-1487, 1996.

Sinnhuber, B.-M., D.W. Arlander, H. Bovensmann, J.P. 

Burrows, M.P. Chipperfield, C.-F. Enell, U. Frieß, 

F. Hendrick, P.V. Johnston, R.L. Jones, K. Kreher, 

N. Mohamed-Tahrin, R. Müller, K. Pfeilsticker, U. 

Platt, J.-P. Pommereau, I. Pundt, A. Richter, A.M. 

South, K.K. Tørnkvist, M. Van Roozendael, T. 

Wagner, and F. Wittrock, Comparison of measurements and model calculations of stratospheric bromine monoxide, J. Geophys. Res., 107 (D19), 4398, 

doi: 10.1029/2001JD000940, 2002.

Sitnov, S., Influence of the 11-year solar cycle on the effects of the equatorial quasi-biennial oscillation, 

manifesting in the extratropical northern atmosphere, Clim. Dyn., 32, 1-17, doi: 10.1007/s00382-

007-0362-6, 2009.

Slusser, J., J. Gibson, D. Bigelow, D. Kolinski, P. Disterhoft, K. Lantz, and A. Beaubien, Langley method 

of calibrating UV filter radiometers, J. Geophys. 

Res., 105 (D4), 4841-4849, 2000.

Smit, H.G.J., W. Straeter, B.J. Johnson, S.J. Oltmans, J. 

Davies, D.W. Tarasick, B. Goegger, R. Stubi, F.J. 

Schmidlin, T. Northam, A.M. Thompson, J.C. 

Witte, I. Boyd, and F. Posny, Assessment of the performance of ECC ozonesondes under quasi-flight 

conditions in the environmental simulation chamber: Insights from the Juelich Ozone Sonde Intercomparison Experiment (JOSIE), J. Geophys. Res., 

112, D19306, doi: 10.1029/2006JD007308, 2007.

Smith, A.K., and K. Matthes, Decadal-scale periodicities 

in the stratosphere associated with the solar cycle 

and the QBO, J. Geophys. Res., 113, D05311, doi: 

10.1029/2007JD009051, 2008.

Soller, R., J.M. Nicovich, and P.H. Wine, Temperaturedependent rate coefficients for the reactions of 

Br(2P3/2), Cl(2P3/2), and O(3PJ) with BrONO2, J. 

Phys. Chem. A, 105 (9), 1416-1422, 2001.

Solomon, P., B. Connor, J. Barrett, T. Mooney, A. Lee, 

and A. Parrish, Measurements of stratospheric ClO 

over Antarctica in 1996–2000 and implications for 

ClO dimer chemistry, Geophys. Res. Lett., 29 (15), 

1708, doi: 10.1029/2002GL015232, 2002.

Solomon, S., R.R. Garcia, F.S. Rowland, and D.J. Wuebbles, On the depletion of Antarctic ozone, Nature, 

321, 755-758, doi: 10.1038/321755a0, 1986.

Solomon, S., R.W. Portmann, T. Sasaki, D.J. Hofmann, 

and D.W.J. Thompson, Four decades of ozonesonde 

measurements over Antarctica, J. Geophys. Res., 
2.72

Chapter 2

110, D21311, doi: 10.1029/2005JD005917, 2005.

Solomon, S., R.W. Portmann, and D.W.J. Thompson, 

Contrasts between Antarctic and Arctic ozone 

depletion, Proc. Nat. Acad. Sci., 104 (2), 445-449, 

doi: 10.1073/pnas.0604895104, 2007.

SPARC (Stratospheric Processes And their Role in Climate), SPARC/IOC/GAW Assessment of Trends in 

the Vertical Distribution of Ozone, edited by N. 

Harris, R. Hudson, and C. Phillips, SPARC Report 

No. 1, WMO Global Ozone Research and Monitoring Project Report No. 43, 289 pp., Verrières le 

Buisson, France, 1998.

SPARC (Stratospheric Processes And their Role in Climate), The Role of Halogen Chemistry in Polar 

Stratospheric Ozone Depletion, Report from the 

June 2008 Cambridge UK Workshop for an Initiative under the Stratospheric Processes and Their 

Role in Climate (SPARC) Project of the World 

Climate Research Programme, authored by M.J. 

Kurylo, B.-M. Sinnhuber, N.R.P. Harris, M. von 

Hobe, P.A. Newman, D.W. Fahey, R.S. Gao, R.J. 

Salawitch, M.P. Chipperfield, J.G. Anderson, M.L. 

Santee, T.P. Canty, R. Müller, R. Schofield, R.M. 

Stimpfle, F. Stroh, D.W. Toohey, J. Urban, S.R. 

Kawa, D.J. Hofmann, K.W. Hoppel, M. Rex, K.D. 

Bayes, D.A. Dixon, K.W. Jucks, S.P. Sander, J.-U. 

Grooß, and D.E. Kinnison, 48 pp., available: http://

www.atmosp.physics.utoronto.ca/SPARC/HalogenChem_Final_20090213.pdf, 2009.

SPARC CCMVal (Stratospheric Processes And their Role 

in Climate), SPARC Report on the Evaluation of 

Chemistry-Climate Models, edited by V. Eyring, 

T.G. Shepherd, and D.W. Waugh, SPARC Report 

No. 5, WCRP-132, WMO/TD-No. 1526, 478 pp., 

available: http://www.atmosp.physics.utoronto.ca/

SPARC, 2010.

Staehelin, J., N.R.P. Harris, C. Appenzeller, and J. Eberhard, Ozone trends: A review, Rev. Geophys., 39, 

231-290, 2001.

Staiger, H., P.N. den Outer, A.F. Bais, H. Feister, B. Johnsen, and L. Vuilleumier, Hourly resolved cloud 

modification factors in the ultraviolet, Atmos. 

Chem. Phys., 8 (9), 2493-2508, doi: 10.5194/acp8-2493-2008, 2008.

Steinbrecht, W., H. Claude, F. Schönenborn, I.S. McDermid, T. Leblanc, S. Godin-Beekmann, P. Keckhut, 

A. Hauchecorne, J.A.E. Van Gijsel, D.P.J. Swart, 

G.E. Bodeker, A. Parrish, I.S. Boyd, N. Kämpfer, 

K. Hocke, R.S. Stolarski, S.M. Frith, L.W. Thomason, E.E. Remsberg, C. Von Savigny, A. Rozanov, 

and J.P. Burrows, Ozone and temperature trends in 

the upper stratosphere at five stations of the Network for the Detection of Atmospheric Composition Change, Int. J. Remote Sens., 30, 3875-3886, 

doi: 10.1080/01431160902821841, 2009.

Stenchikov, G., A. Robock, V. Ramaswamy, M.D. 

Schwarzkopf, K. Hamilton, and S. Ramachandran, 

Arctic Oscillation response to the 1991 Mount Pinatubo eruption: Effects of volcanic aerosols and 

ozone depletion, J. Geophys. Res., 107 (D24), 

4803, doi: 10.1029/2002JD002090, 2002.

Stimpfle, R.M., R.A. Perry, and C.J. Howard, Temperature-dependence of the reaction of ClO and HO2

radicals, J. Chem. Phys., 71 (12), 5183-5190, doi: 

10.10631/1.438293, 1979.

Stimpfle, R., D.M. Wilmouth, R.J. Salawitch, and J.G. 

Anderson, First measurements of ClOOCl in the 

stratosphere: The coupling of ClOOCl and ClO 

in the Arctic polar vortex, J. Geophys. Res., 109, 

D03301, doi: 10.1029/2003JD003811, 2004.

Stolarski, R.S., and S.M. Frith, Search for evidence of 

trend slow-down in the long-term TOMS/SBUV 

total ozone data record: The importance of instrument drift uncertainty, Atmos. Chem. Phys., 6, 

4057-4065, 2006.

Stolarski, R.S., A.R. Douglass, S. Steenrod, and S. Pawson, Trends in stratospheric ozone: Lessons learned 

from a 3D Chemical Transport Model, J. Atmos. 

Sci., 63 (3), 1028-1041, 2006.

Stübi, R., G. Levrat, B. Hoegger, P. Viatte, J. Staehelin, 

and F. J. Schmidlin, In-flight comparison of Brewer-Mast and electrochemical concentration cell 

ozonesondes, J. Geophys. Res., 113, D13302, doi: 

10.1029/2007JD009091, 2008.

Tabazadeh, A., R.P. Turco, K. Drdla, M.Z. Jacobson, and 

O.B. Toon, A study of Type I polar stratospheric 

cloud formation, Geophys. Res. Lett., 21 (15), 

1619-1622, 1994.

Tabazadeh, A., Y.S. Djikaev, P. Hamill, and H. Reiss, 

Laboratory evidence for surface nucleation of solid 

polar stratospheric cloud particles, J. Phys. Chem. 

A, 106, 10238-10246, 2002.

Takahashi, K., Y. Takeuchi, and Y. Matsumi, Rate constants of the O(1

D) reactions with N2, O2, N2O, and 

H2O at 295 K, Chem. Phys. Lett., 410, 196-200, 

2005.

Tanskanen, A., N.A. Krotkov, J.R. Herman, and A. Arola, 

Surface ultraviolet irradiance from OMI, Geosci. 

Remote Sens., 44 (5), 1267-1271, doi: 10.1109/

TGRS.2005.862203, 2006.

Tanskanen, A., A. Lindfors, A. Määttä, N. Krotkov, J. 

Herman, J. Kaurola, T. Koskela, K. Lakkala, V. 

Fioletov, G. Bernhard, R. McKenzie, Y. Kondo, 

M. O’Neill, H. Slaper, P. den Outer, A.F. Bais, and 

J. Tamminen, Validation of daily erythemal doses 

from Ozone Monitoring Instrument with groundbased UV measurement data, J. Geophys. Res., 

112, D24S44, doi: 10.1029/2007JD008830, 2007.
2.73

Stratospheric Ozone and Surface UV

Tatarov, B., H. Nakane, Ch. B. Park, N. Sugimoto, and 

I. Matsui, Lidar observation of long-term trends 

and variations of stratospheric ozone and temperature over Tsukuba, Japan, Int. J. Remote 

Sens., 30 (15-16), 3951-3960, doi: 10.1080/ 

01431160902821882, 2009.

Tegtmeier, S., M. Rex, I. Wohltmann, and K. Krüger, 

Relative importance of dynamical and chemical contributions to Arctic wintertime ozone, 

Geophys. Res. Lett., 35, L17801, doi: 10.1029/ 

2008GL034250, 2008.

Telford, P., P. Braesicke, O. Morgenstern, and J. Pyle, Reassessment of causes of ozone column variability 

following the eruption of Mount Pinatubo using a 

nudged CCM, Atmos. Chem. Phys., 9 (13), 4251-

4260, doi: 10.5194/acp-9-4251-2009, 2009.

Terao, Y., and J.A. Logan, Consistency of time series and 

trends of stratospheric ozone as seen by ozonesonde, 

SAGE II, HALOE, and SBUV(/2), J. Geophys. Res., 

112, D06310, doi: 10.1029/2006JD007667, 2007.

Terao, Y., J.A. Logan, A.R. Douglass, and R.S. Stolarski, 

Contribution of stratospheric ozone to the interannual variability of tropospheric ozone in the northern extratropics, J. Geophys. Res., 113, D18309, 

doi: 10.1029/2008JD009854, 2008.

Thiel, S., L. Ammannato, A. Bais, B. Bandy, M. Blumthaler, B. Bohn, O. Engelsen, G.P. Gobbi, J. Gröbner, 

E. Jäkel, W. Junkermann, S. Kazadzis, R. Kift, B. 

Kjeldstad, N. Kouremeti, A. Kylling, B. Mayer, 

P.S. Monks, C.E. Reeves, B. Schallhart, R. Scheirer, S. Schmidt, R. Schmitt, J. Schreder, R. Silbernagl, C. Topaloglou, T.M. Thorseth, A.R. Webb, 

M. Wendisch, and P. Werle, Influence of clouds on 

the spectral actinic flux density in the lower troposphere (INSPECTRO): Overview of the field campaigns, Atmos. Chem. Phys., 8 (6), 1789-1812, doi: 

10.5194/acp-8-1789-2008, 2008.

Thompson, A.M., J.B. Stone, J.C. Witte, S.K. Miller, R.B. 

Pierce, R.B. Chatfield, S.J. Oltmans, O.R. Cooper, 

A.L. Loucks, B.F. Taubman, B.J. Johnson, E. Joseph, T.L. Kucsera, J.T. Merrill, G.A. Morris, S. 

Hersey, G. Forbes, M.J. Newchurch, F.J. Schmidlin, 

D.W. Tarasick, V. Thouret, and J.-P. Cammas, Intercontinental Chemical Transport Experiment Ozonesonde Network Study (IONS) 2004: 1. Summertime 

upper troposphere/lower stratosphere ozone over 

northeastern North America, J. Geophys. Res., 112, 

D12S12, doi: 10.1029/2006JD007441, 2007.

Thouret, V., J.-P. Cammas, B. Sauvage, G. Athier, R. 

Zbinden, P. Nédélec, P. Simon, and F. Karcher, 

Tropopause referenced ozone climatology and 

inter-annual variability (1994–2003) from the 

MOZAIC programme, Atmos. Chem. Phys., 6 (4), 

1033-1051, doi: 10.5194/acp-6-1033-2006, 2006.

Tie, X.X., and G.P. Brasseur, The response of stratospheric ozone to volcanic eruptions: Sensitivity to 

atmospheric chlorine loading, Geophys. Res. Lett., 

22 (22), 3035-3038, 1995.

Tie, X.X., G.P. Brasseur, B. Briegleb, and C. Granier, 

Two-dimensional simulation of Pinatubo aerosol 

and its effect on stratospheric ozone, J. Geophys. 

Res., 99 (D10), 20545-20562, 1994.

Tilmes, S., R. Müller, J.-U. Grooß, and J.M. Russell III, 

Ozone loss and chlorine activation in the Arctic 

winters 1991–2003 derived with the tracer-tracer 

correlations, Atmos. Chem. Phys., 4 (8), 2181-2213, 

doi: 10.5194/acp-4-2181-2004, 2004.

Tilmes, S., R. Müller, A. Engel, M. Rex, and J. M. 

Russell, Chemical ozone loss in the Arctic and 

Antarctic stratosphere between 1992 and 2005, 

Geophys. Res. Lett., 33, L20812, doi: 10.1029/ 

2006GL026925, 2006.

Tilmes, S., R. Müller, and R. Salawitch, The sensitivity of 

polar ozone depletion to proposed geoengineering 

schemes, Science, 320, 1201-1204, 2008a.

Tilmes, S., R. Müller, R.J. Salawitch, U. Schmidt, C.R. 

Webster, H. Oelhaf, C.C. Camy-Peyret, and J.M. 

Russell III, Chemical ozone loss in the Arctic winter 1991–1992, Atmos. Chem. Phys., 8 (7), 1897-

1910, doi: 10.5194/acp-8-1897-2008, 2008b.

Toohey, D.W., L.M. Avallone, L.R. Lait, P.A. Newman, 

M.R. Schoeberl, D.W. Fahey, E.L. Woodbridge, 

and J.G. Anderson, The seasonal evolution of reactive chlorine in the northern hemisphere stratosphere, Science, 261 (5125), 1134-1136, doi: 

10.1126/science.261-5125.1134, 1993.

Toon, O., A. Tabazadeh, E. Browell, and J. Jordan, Analysis of lidar observations of Arctic polar stratospheric clouds during January 1989, J. Geophys. Res., 

105 (D16), 20589-20615, 2000.

Tourpali, K., C.S. Zerefos, D.S. Balis, and A.F. Bais, The 

11-year solar cycle in stratospheric ozone: Comparison between Umkehr and SBUVv8 and effects 

on surface erythemal irradiance, J. Geophys. Res., 

112, D12306, doi: 10.1029/2006JD007760, 2007.

Tsutsui, J., K. Nishizawa, and F. Sassi, Response of the 

middle atmosphere to the 11-year solar cycle simulated with the Whole Atmosphere Community Climate Model, J. Geophys. Res., 114, D02111, doi: 

10.1029/2008JD010316, 2009.

Van Roozendael, M., D. Loyola, R. Spurr, D. Balis, J.-

C. Lambert, Y. Livschitz, P. Valks, T. Ruppert, 

P. Kenter, C. Fayt, and C. Zehner, Ten years of 

GOME/ERS-2 total ozone data – The new GOME 

data processor (GDP) version 4: 1. Algorithm description, J. Geophys. Res., 111, D14311, doi: 

10.1029/2005JD006375, 2006.

Verdebout, J., A method to generate surface UV radiation 
2.74

Chapter 2

maps over Europe using GOME, Meteosat, and 

ancillary geophysical data, J. Geophys. Res., 105

(D4), 5049-5058, 2000.

Verdebout, J., A European satellite-derived UV climatology available for impact studies, Radiat. Prot. Dosim., 111 (4), 407-411, 2004a.

Verdebout, J., A satellite-derived UV radiation climatology over Europe to support impact studies, Arct. 

Antarct. Alp. Res., 36 (3), 357-363, 2004b.

Vernet, M., S.B. Diaz, H.A. Fuenzalida, C. Camilion, 

C.R. Booth, S. Cabrera, C. Casiccia, G. Deferrari, 

C. Lovengreen, A. Paladini, J. Pedroni, A. Rosales, 

and H.E. Zagarese, Quality of UVR exposure for 

different biological systems along a latitudinal gradient, Photochem. Photobiol., 8, 1329-1345, doi: 

10.1039/b904540f, 2009.

Vernier, J.P., J.P. Pommereau, A. Garnier, J. Pelon, N. 

Larsen, J. Nielsen, T. Christensen, F. Cairo, L.W. 

Thomason, T. Leblanc, and I.S. McDermid, Tropical stratospheric aerosol layer from CALIPSO lidar 

observations, J. Geophys. Res., 114, D00H10, doi: 

10.1029/2009JD011946, 2009.

Vigouroux C., M. De Mazière, P. Demoulin, C. Servais, 

F. Hase, T. Blumenstock, I. Kramer, M. Schneider, 

J. Mellqvist, A. Strandberg, V. Velazco, J. Notholt, R. Sussmann, W. Stremme, A. Rockmann, T. 

Gardiner, M. Coleman, and P. Woods, Evaluation 

of tropospheric and stratospheric ozone trends over 

Western Europe from ground-based FTIR network 

observations, Atmos. Chem. Phys., 8 (23), 6865-

6886, doi: 10.5194/acp-8-6865-2008, 2008.

Vogel, B., R. Müller, T. Deshler, J.-U. Grooβ, J. Karhu, 

D.S. McKenna, M. Müller, D. Toohey, G.C. Toon, 

and F. Stroh, Vertical profiles of activated ClO 

and ozone loss in the Arctic vortex in January and 

March 2000: In situ observations and model simulations, J. Geophys. Res., 108 (D22), 8334, doi: 

10.1029/2002JD002564, 2003.

Voigt, C., J. Schreiner, A. Kohlman, P. Zink, K. Mauersberger, N. Larsen, T. Deshler, C. Kröger, J. 

Rosen, A. Adriani, F. Cairo, G. Di Donfrancesco, 

M. Viterbini, J. Ovarlez, H. Ovarlez, C. David, and 

A. Dörnbrack, Nitric acid trihydrate (NAT) in polar stratospheric clouds, Science, 290, 1756-1758, 

2000.

von Hobe, M., Revisiting ozone depletion, Science, 318, 

18781879, 2007.

von Hobe, M., R.J. Salawitch, T. Canty, H. Keller-Rudek, 

G.K. Moortgat, J.-U. Grooß, R. Müller, and F. 

Stroh, Understanding the kinetics of the ClO dimer 

cycle, Atmos. Chem. Phys., 7 (12), 3055-3069, doi: 

10.5194/acp-7-3055-2007, 2007.

von Hobe, M., F. Stroh, H. Beckers, T. Benter, and H. 

Willner, The UV/Vis absorption spectrum of matrix-isolated dichlorine peroxide, ClOOCl, Phys. 

Chem. Chem. Phys., 11, 1571-1580, doi: 10.1039/

b814373k, 2009.

Vranckx, S., J. Peeters, and S.A. Carl, Absolute rate constant and O(3

P) yield for the O(1D) + N2O reaction 

in the temperature range 227 K to 719 K, Atmos. 

Chem. Phys., 8 (20), 6261-6272, doi: 10.5194/acp8-6261-2008, 2008.

Vyushin, D., V.E. Fioletov, and T.G. Shepherd, Impact 

of long-range correlations on trend detection in 

total ozone, J. Geophys. Res., 112, D14307, doi: 

10.1029/2006JD008168, 2007.

Vyushin, D., T.G. Shepherd, and V.E. Fioletov, On the 

statistical modeling of persistence in total ozone 

anomalies, J. Geophys. Res., 115, D16306, doi: 

10.1029/2009JD013105, 2010.

Waibel, A.E., T. Peter, K.S. Carslaw, H. Oelhaf, G. Wetzel, P.J. Crutzen, U. Pöschl, A. Tsias, E. Reimer, 

and H. Fischer, Arctic ozone loss due to denitrification, Science, 283, 2064-2069, 1999.

Wang, K., R.E. Dickinson, and S. Liang, Clear sky visibility has decreased over land globally from 1973 

to 2007, Science, 323 (5921), 1468-1470, doi: 

10.1126/science.1167549, 2009.

Wang, Z., G. Stephens, T. Deshler, C. Trepte, T. Parish, D. 

Vane, D. Winker, D. Liu, and L. Adhikari, Association of Antarctic polar stratospheric cloud formation 

on tropospheric cloud systems, Geophys. Res. Lett.,

35, L13806, doi: 10.1029/2008GL034209, 2008.

Waugh, D., Atmospheric dynamics: The age of stratospheric air, Nature Geosci., 2, 14-16, doi: 10.1038/

nge0397, 2009.

Waugh, D.W., and T.M. Hall, Age of stratospheric air: 

Theory, observations, and models, Rev. Geophys.,

40, doi: 10.1029/2000RG000101, 2000.

Waugh, D.W., and V. Eyring, Quantitative performance 

metrics for stratospheric-resolving chemistryclimate models, Atmos. Chem. Phys., 8 (18), 5699-

5713, doi: 10.5194/acp-8-5699-2008, 2008.

Waugh, D.W., L. Oman, S.R. Kawa, R.S. Stolarski, S. 

Pawson, A.R. Douglass, P.A. Newman, and J.E. 

Nielsen, Impacts of climate change on stratospheric 

ozone recovery, Geophys. Res. Lett., 36, L03805, 

doi: 10.1029/2008GL036223, 2009.

Weatherhead, E.C., G.C. Reinsel, G.C. Tiao, C.H. Jackman, L. Bishop, S.M.H. Frith, J. DeLuisi, T. Keller, 

S.J. Oltmans, E.L. Fleming, D.J. Wuebbles, J.B. 

Kerr, A.J. Miller, J. Herman, R. McPeters, R.M. 

Nagatani, and J.E. Frederick, Detecting the recovery of total column ozone, J. Geophys. Res., 105

(D17), 22201-22210, 2000.

Webb, A.R., B.G. Gardine, T.J. Martin, K. Leszcynsk, J. 

Metzdorf, and V.A. Mohnen, Guidelines for Site 

Quality Control of UV Monitoring, WMO/GAW 
2.75

Stratospheric Ozone and Surface UV

No. 126, WMO TD 884, World Meteorological Organization, Geneva, 39 pp., 1998.

Webb, A.R., J. Gröbner, and M. Blumthaler, A Practical 

Guide to Operating Broadband Instruments Measuring Erythemally Weighted Irradiance, publication of COST 726 and World Meteorological Organization, 21 pp., European Cooperation in Science 

and Technology (COST) Office, Luxembourg, 

available: http://www.cost726.org/, 2006.

Weber, M., L.N. Lamsal, M. Coldewey-Egbers, K. 

Bramstedt, and J.P. Burrows, Pole-to-pole validation of GOME WFDOAS total ozone with groundbased data, Atmos. Chem. Phys., 5 (5), 1341-1355, 

doi: 10.5194/acp-5-1341-2005, 2005.

Weber, M., L.N. Lamsal, and J.P. Burrows, Improved 

SCIAMACHY WFDOAS total ozone retrieval: Steps towards homogenising long-term total 

ozone datasets from GOME, SCIAMACHY, and 

GOME2, in Proceedings of the Envisat Symposium 

2007, Montreux, Switzerland, 23-27 April 2007, 

ESA SP-636, available: http://envisat.esa.int/envisatsymposium/proceedings/posters/3P4/463281we.

pdf, 2007.

Weihs, P., M. Blumthaler, H.E. Rieder, A. Kreuter, S. 

Simic, W. Laube, A.W. Schmalwieser, J.E. Wagner, and A. Tanskanen, Measurements of UV irradiance within the area of one satellite pixel, Atmos. 

Chem. Phys., 8 (18), 5615-5626, doi: 10.5194/acp8-5615-2008, 2008.

Weiss, A.K., J. Staehelin, C. Appenzeller, and N.R.P. Harris, Chemical and dynamical contributions to ozone 

profile trends of the Payerne (Switzerland) balloon 

soundings, J. Geophys. Res., 106 (D19), 22685-

22694, 2001.

Wetzel, G., H. Oelhaf, O. Kirner, R. Ruhnke, F. FriedlVallon, A. Kleinert, G. Maucher, H. Fischer, M. 

Birk, G. Wagner, and A. Engel, First remote sensing measurements of ClOOCl along with ClO and 

ClONO2 in activated and deactivated Arctic vortex 

conditions using new ClOOCl IR absorption cross 

sections, Atmos. Chem. Phys., 10 (3), 931-945, doi: 

10.5194/acp-10-931-2010, 2010.

Wilmouth, D.M., R.M. Stimpfle, J.G. Anderson, J.W. 

Elkins, D.F. Hurst, R.J. Salawitch, and L.R. Lait, 

Evolution of inorganic chlorine partitioning in the 

Arctic polar vortex, J. Geophys. Res., 111, D16308, 

doi: 10.1029/2005JD006951, 2006.

Wilmouth, D.M., T.F. Hanisco, R.M. Stimpfle, and J.G. 

Anderson, Chlorine-catalyzed ozone destruction: 

Cl atom production from ClOOCl photolysis, J. 

Phys. Chem. A, 113, 14099-14108, doi: 10.1021/

jp9053204, 2009.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 1994, Global 

Ozone Research and Monitoring Project–Report 

No. 37, Geneva, Switzerland, 1995.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 1998, Global 

Ozone Research and Monitoring Project–Report 

No. 44, Geneva, Switzerland, 1999.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 2002, Global 

Ozone Research and Monitoring Project–Report 

No. 47, Geneva, Switzerland, 2003.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 2006, Global 

Ozone Research and Monitoring Project–Report 

No. 50, 572 pp., Geneva, Switzerland, 2007.

WMO (World Meteorological Organization), Instruments 

to Measure Solar Ultraviolet Radiation – Part 2: 

Broadband Instruments Measuring Erythemally 

Weighted Solar Irradiance, WMO TD No. 1289, G. 

Seckmeyer, A. Bais, G. Bernhard, M. Blumthaler, 

C.R. Booth, K. Lantz, and R.L. McKenzie, 55 pp., 

Geneva, Switzerland, available: http://www.wmo.

int/pages/prog/arep/gaw/gaw-reports.html, 2008.

Wohltmann, I., R. Lehmann, M. Rex, D. Brunner, and J. 

Mäder, A process-oriented regression model for 

column ozone, J. Geophys. Res., 112, D12304, doi: 

10.1029/2006JD007573, 2007.

Wuttke, S., J. Verdebout, and G. Seckmeyer, An improved 

algorithm for satellite-derived UV radiation, Photochem. Photobiol., 77 (1), 52-57, 2003.

Wuttke, S., G. Seckmeyer, and G. König-Langlo, Measurements of spectral snow albedo at Neumayer, 

Antarctica, Ann. Geophys., 24, 7-21, 2006a.

Wuttke, S., G. Seckmeyer, G. Bernhard, J. Ehramjian, 

R. McKenzie, P. Johnston, and M. O’Neill, New 

spectroradiometers complying with the NDSC 

standards, J. Atmos. Oceanic Tech., 23, 241-251, 

2006b.

Yang, E.-S., D.M. Cunnold, R.J. Salawitch, M.P. McCormick, J. Russell III, J.M. Zawodny, S. Oltmans, 

and M.J. Newchurch, Attribution of recovery in 

lower-stratospheric ozone, J. Geophys. Res., 111, 

D17309, doi: 10.1029/2005JD006371, 2006.

Yang, E.-S., D.M. Cunnold, M.J. Newchurch, R.J. Salawitch, M.P. McCormick, J.M. Russell III, J.M. Zawodny, and S.J. Oltmans, First stage of Antarctic 

ozone recovery, J. Geophys. Res., 113, D20308, 

doi: 10.1029/2007JD009675, 2008.

Yang, M., S.G. Howell, J. Zhuang, and B.J. Huebert, Attribution of aerosol light absorption to black carbon, 

brown carbon, and dust in China – interpretations 

of atmospheric measurements during EASTAIRE, Atmos. Chem. Phys., 9 (6), 2035-2050, doi: 

10.5194/acp-9-2035-2009, 2009.

Yang, S.-K., C.S. Long, A.J. Miller, X. He, Y. Yang, 
2.76

Chapter 2

D.J. Wuebbles, and G. Tiao, Modulation of natural variability on a trend analysis of updated cohesive SBUV(/2) total ozone, Int. J. Remote Sens., 

30, 3975-3986, doi: 10.1080/01431160902821924, 

2009.

Zanis, P., E. Maillard, J. Staehelin, C. Zerefos, E. Kosmidis, K. Tourpali, and I. Wohltmann, On the 

turnaround of stratospheric ozone trends deduced 

from the reevaluated Umkehr record of Arosa, 

Switzerland, J. Geophys. Res., 111, D22307, doi: 

10.1029/2005JD006886, 2006.

Zhao, J., R.P. Turco, and O.B. Toon, A model simulation 

of Pinatubo volcanic aerosols in the stratosphere, J. 

Geophys. Res., 100 (D4), 7315-7328, 1995.

Ziemke, J.R., S. Chandra, and P.K. Bhartia, A 25-year 

data record of atmospheric ozone in the Pacific 

from Total Ozone Mapping Spectrometer (TOMS) 

cloud slicing: Implications for ozone trends in the 

stratosphere and troposphere, J. Geophys. Res., 

110, D15105, doi: 10.1029/2004JD005687, 2005.
Chapter 3

Future Ozone and Its Impact on Surface UV

Coordinating Lead Authors: Contributors:

S. Bekki H. Akiyoshi

G.E. Bodeker J. Austin

M.P. Chipperfield

Lead Authors: M. Dameris

A.F. Bais S. Dhomse

N. Butchart S.M. Frith

V. Eyring R.R. Garcia

D.W. Fahey H. Garny

D.E. Kinnison A. Gettelman

U. Langematz S.C. Hardiman

B. Mayer P. Jöckel

R.W. Portmann A.I. Jonsson

E. Rozanov A. Kazantzidis

A. Kubin

Coauthors: J.-F. Lamarque

P. Braesicke E. Mancini

A.J. Charlton-Perez M. Marchand

N.E. Chubarova M. Michou

I. Cionni O. Morgenstern

S.B. Diaz L.D. Oman

N.P. Gillett S. Pawson

M.A. Giorgetta G. Pitari

N. Komala D. Plummer

F. Lefèvre J.A. Pyle

C. McLandress D. Saint-Martin

J. Perlwitz J.F. Scinocca

T. Peter T.G. Shepherd

K. Shibata D. Smale

R.S. Stolarski

H. Teyssèdre

S. Tilmes

Chapter 3

FUTURE OZONE AND ITS IMPACT ON SURFACE UV

Contents

SCIENTIFIC SUMMARY .............................................................................................................................................1

3.1 INTRODUCTION ...................................................................................................................................................5

Box 3-1. Measures of Atmospheric Halogens......................................................................................................5

Box 3-2. Ozone Return Dates and Full Ozone Recovery.....................................................................................7

3.2 FACTORS AFFECTING FUTURE OZONE AND SURFACE UV......................................................................7

3.2.1 Stratospheric Halogen Loading.................................................................................................................8

3.2.2 Stratospheric Reactive Nitrogen and Hydrogen Levels............................................................................8

3.2.3 Stratospheric Temperatures.......................................................................................................................9

3.2.4 Transport and Dynamics .........................................................................................................................10

3.2.4.1 Brewer-Dobson Circulation......................................................................................................10

3.2.4.2 Vortex Integrity and Mixing .....................................................................................................10

3.2.5 Background and Volcanic Stratospheric Aerosols..................................................................................12

3.2.6 Geoengineering by Sulfate Aerosol Injection .........................................................................................13

3.2.7 Effects of Ozone on Future Surface UV .................................................................................................15

3.2.8 Factors Other Than Stratospheric Ozone Affecting Surface UV............................................................16

3.2.8.1 Clouds.......................................................................................................................................16

3.2.8.2 Aerosols ....................................................................................................................................17

3.2.8.3 Surface Albedo and Sea Ice Cover ...........................................................................................17

3.2.8.4 Tropospheric Gases...................................................................................................................17

3.3 PROJECTIONS OF OZONE THROUGH THE 21ST CENTURY .......................................................................18

3.3.1 Model Descriptions and Scenarios..........................................................................................................18

3.3.2 Model Evaluation and Multi-Model Mean Analysis...............................................................................19

3.3.2.1 Model Evaluation......................................................................................................................19

3.3.2.2 Analysis Method for Multi-Model Time Series........................................................................22

3.3.3 Tropical Ozone........................................................................................................................................22

3.3.3.1 Long-Term Projections of Tropical Ozone...............................................................................22

3.3.3.2 Processes Determining Future Tropical Ozone ........................................................................25

3.3.4 Midlatitude Ozone...................................................................................................................................28

3.3.4.1 Long-Term Projections of Midlatitude Ozone..........................................................................28

3.3.4.2 Processes Determining Future Midlatitude Ozone ...................................................................28

3.3.5 Polar Ozone .............................................................................................................................................30

3.3.5.1 Long-Term Projections of Polar Ozone....................................................................................30

3.3.5.2 Processes Determining Future Polar Ozone .............................................................................34

3.3.6 Ozone Return Dates and Ozone Recovery..............................................................................................36

3.3.7 Uncertainties in Model Projections and Open Questions........................................................................41

3.3.7.1 Uncertainty in Future Emissions Scenarios..............................................................................41

3.3.7.2 Future CCM Development........................................................................................................42

3.4 PROJECTIONS OF UV CHANGES RELATED TO OZONE CHANGES 

THROUGH THE 21ST CENTURY .......................................................................................................................42

3.4.1 Midlatitude and Tropical UV ..................................................................................................................42

3.4.2 Polar UV..................................................................................................................................................44

3.4.3 Link to UNEP Environmental Effects Panel Assessment.......................................................................46
3.5 CONCLUSIONS ...................................................................................................................................................46

REFERENCES .............................................................................................................................................................49

APPENDIX 3A: CONSTRUCTING CORRELATIVE TIME SERIES PLOTS........................................................59
3.1

Future Ozone and UV

SCIENTIFIC SUMMARY

Globally averaged total column ozone has declined over recent decades due to the release of ozone-depleting substances (ODSs) into the atmosphere. Now, as a result of the Montreal Protocol, ozone is expected to recover from the 

effects of ODSs as ODS abundances decline in the coming decades. However, a number of factors in addition to ODSs 

have led to and will continue to lead to changes in ozone. Discriminating between the causes of past and projected ozone 

changes is necessary, not only to identify the progress in ozone recovery from ODSs, but also to evaluate the effectiveness 

of climate and ozone protection policy options.

Factors Affecting Future Ozone and Surface Ultraviolet Radiation

•		 At least for the next few decades, the decline of ODSs is expected to be the major factor affecting the

anticipated increase in global total column ozone. However, several factors other than ODS will affect the

future evolution of ozone in the stratosphere. These include changes in (i) stratospheric circulation and temperature due to changes in long-lived greenhouse gas (GHG) abundances, (ii) stratospheric aerosol loading, and 

(iii) source gases of highly reactive stratospheric hydrogen and nitrogen compounds. Factors that amplify the 

effects of ODSs on ozone (e.g., stratospheric aerosols) will likely decline in importance as ODSs are gradually 

eliminated from the atmosphere.

•		 Increases in GHG emissions can both positively and negatively affect ozone. Carbon dioxide (CO2)-induced 

stratospheric cooling elevates middle and upper stratospheric ozone and decreases the time taken for ozone to 

return to 1980 levels, while projected GHG-induced increases in tropical upwelling decrease ozone in the tropical lower stratosphere and increase ozone in the extratropics. Increases in nitrous oxide (N2O) and methane 

(CH4) concentrations also directly impact ozone chemistry but the effects are different in different regions.

•		 The Brewer-Dobson circulation (BDC) is projected to strengthen over the 21st century and thereby affect

ozone amounts. Climate models consistently predict an acceleration of the BDC or, more specifically, of the 

upwelling mass flux in the tropical lower stratosphere of around 2% per decade as a consequence of GHG abundance increases. A stronger BDC would decrease the abundance of tropical lower stratospheric ozone, increase 

poleward transport of ozone, and could reduce the atmospheric lifetimes of long-lived ODSs and other trace 

gases. While simulations showing faster ascent in the tropical lower stratosphere to date are a robust feature of 

chemistry-climate models (CCMs), this has not been confirmed by observations and the responsible mechanisms 

remain unclear.

•		 Substantial ozone losses could occur if stratospheric aerosol loading were to increase in the next few

decades, while halogen levels are high. Stratospheric aerosol increases may be caused by sulfur contained in 

volcanic plumes entering the stratosphere or from human activities. The latter might include attempts to geoengineer the climate system by enhancing the stratospheric aerosol layer. The ozone losses mostly result from 

enhanced heterogeneous chemistry on stratospheric aerosols. Enhanced aerosol heating within the stratosphere 

also leads to changes in temperature and circulation that affect ozone.

•		 Surface ultraviolet (UV) levels will not be affected solely by ozone changes but also by the effects of climate change and by air quality change in the troposphere. These tropospheric effects include changes in 

clouds, tropospheric aerosols, surface reflectivity, and tropospheric sulfur dioxide (SO2) and nitrogen dioxide 

(NO2). The uncertainties in projections of these factors are large. Projected increases in tropospheric ozone are 

more certain and may lead to reductions in surface erythemal (“sunburning”) irradiance of up to 10% by 2100. 

Changes in clouds may lead to decreases or increases in surface erythemal irradiance of up to 15% depending 

on latitude.
3.2

Chapter 3

Expected Future Changes in Ozone

Full ozone recovery from the effects of ODSs and return of ozone to historical levels are not synonymous. In this 

chapter a key target date is chosen to be 1980, in part to retain the connection to previous Ozone Assessments. Noting, 

however, that decreases in ozone may have occurred in some regions of the atmosphere prior to 1980, 1960 return dates 

are also reported.

The projections reported on in this chapter are taken from a recent compilation of CCM simulations. The ozone 

projections, which also form the basis for the UV projections, are limited in their representativeness of possible futures 

since they mostly come from CCM simulations based on a single GHG emissions scenario (scenario A1B of Emissions 

Scenarios. A Special Report of Working Group III of the Intergovernmental Panel on Climate Change, Cambridge 

University Press, 2000) and a single ODS emissions scenario (adjusted A1 of the previous (2006) Ozone Assessment).

Throughout this century, the vertical, latitudinal, and seasonal structure of the ozone distribution will be different from 

what it was in 1980. For this reason, ozone changes in different regions of the atmosphere are considered separately.

•		 The projections of changes in ozone and surface clear-sky UV are broadly consistent with those reported

on in the 2006 Assessment.

•		 The capability of making projections and attribution of future ozone changes has been improved since the

2006 Assessment. Use of CCM simulations from an increased number of models extending through the entire 

period of ozone depletion and recovery from ODSs (1960–2100) as well as sensitivity simulations have allowed 

more robust projections of long-term changes in the stratosphere and of the relative contributions of ODSs and 

GHGs to those changes.

•		 Global annually averaged total column ozone is projected to return to 1980 levels before the middle of

the century and earlier than when stratospheric halogen loading returns to 1980 levels. CCM projections 

suggest that this early return is primarily a result of GHG-induced cooling of the upper stratosphere because the 

effects of circulation changes on tropical and extratropical ozone largely cancel. Global (90°S–90°N) annually 

averaged total column ozone will likely return to 1980 levels between 2025 and 2040, well before the return of 

stratospheric halogens to 1980 levels between 2045 and 2060.

•		 Simulated changes in tropical total column ozone from 1960 to 2100 are generally small. The evolution of 

tropical total column ozone in models depends on the balance between upper stratospheric increases and lower 

stratospheric decreases. The upper stratospheric increases result from declining ODSs and a slowing of ozone 

destruction resulting from GHG-induced cooling. Ozone decreases in the lower stratosphere mainly result from 

an increase in tropical upwelling. From 1960 until around 2000, a general decline is simulated, followed by 

a gradual increase to values typical of 1980 by midcentury. Thereafter, although total column ozone amounts 

decline slightly again toward the end of the century, by 2080 they are no longer expected to be affected by ODSs. 

Confidence in tropical ozone projections is compromised by the fact that simulated decreases in column ozone 

to date are not supported by observations, suggesting that significant uncertainties remain.

•		 Midlatitude total column ozone is simulated to evolve differently in the two hemispheres. Over northern 

midlatitudes, annually averaged total column ozone is projected to return to 1980 values between 2015 and 2030, 

while for southern midlatitudes the return to 1980 values is projected to occur between 2030 and 2040. The 

more rapid return to 1980 values in northern midlatitudes is linked to a more pronounced strengthening of the 

poleward transport of ozone due to the effects of increased GHG levels, and effects of Antarctic ozone depletion 

on southern midlatitudes. By 2100, midlatitude total column ozone is projected to be above 1980 values in both 

hemispheres.

•		 October-mean Antarctic total column ozone is projected to return to 1980 levels after midcentury, later

than in any other region, and yet earlier than when stratospheric halogen loading is projected to return to

1980 levels. The slightly earlier return of ozone to 1980 levels (2045–2060) results primarily from upper stratospheric cooling and resultant increases in ozone. The return of polar halogen loading to 1980 levels (2050–2070) 
3.3

Future Ozone and UV

in CCMs is earlier than in empirical models that exclude the effects of GHG-induced changes in circulation. Our 

confidence in the drivers of changes in Antarctic ozone is higher than for other regions because (i) ODSs exert a 

strong influence on Antarctic ozone, (ii) the effects of changes in GHG abundances are comparatively small, and 

(iii) projections of ODS emissions are more certain than those for GHGs. Small Antarctic ozone holes (areas of 

ozone <220 Dobson units, DU) could persist to the end of the 21st century.

•		 March-mean Arctic total column ozone is projected to return to 1980 levels two to three decades before

polar halogen loading returns to 1980 levels, and to exceed 1980 levels thereafter. While CCM simulations 

project a return to 1980 levels between 2020 and 2035, most models tend not to capture observed low temperatures and thus underestimate present-day Arctic ozone loss such that it is possible that this return date is biased 

early. Since the strengthening of the Brewer-Dobson circulation through the 21st century leads to increases in 

springtime Arctic column ozone, by 2100 Arctic ozone is projected to lie well above 1960 levels.

Uncertainties in Projections

•		 Conclusions dependent on future GHG levels are less certain than those dependent on future ODS levels

since ODS emissions are controlled by the Montreal Protocol. For the six GHG scenarios considered by a 

few CCMs, the simulated differences in stratospheric column ozone over the second half of the 21st century are 

largest in the northern midlatitudes and the Arctic, with maximum differences of 20–40 DU between the six 

scenarios in 2100.

•		 There remain sources of uncertainty in the CCM simulations. These include the use of prescribed ODS mixing ratios instead of emission fluxes as lower boundary conditions, the range of sea surface temperatures and sea 

ice concentrations, missing tropospheric chemistry, model parameterizations, and model climate sensitivity.

•		 Geoengineering schemes for mitigating climate change by continuous injections of sulfur-containing compounds into the stratosphere, if implemented, would substantially affect stratospheric ozone, particularly

in polar regions. Ozone losses observed following large volcanic eruptions support this prediction. However, 

sporadic volcanic eruptions provide limited analogs to the effects of continuous sulfur emissions. Preliminary 

model simulations reveal large uncertainties in assessing the effects of continuous sulfur injections.

Expected Future Changes in Surface UV

While a number of factors, in addition to ozone, affect surface UV irradiance, the focus in this chapter is on the 

effects of changes in stratospheric ozone on surface UV. For this reason, clear-sky surface UV irradiance is calculated 

from ozone projections from CCMs.

•		 Projected increases in midlatitude ozone abundances during the 21st century, in the absence of changes

in other factors, in particular clouds, tropospheric aerosols, and air pollutants, will result in decreases in

surface UV irradiance. Clear-sky erythemal irradiance is projected to return to 1980 levels on average in 2025 

for the northern midlatitudes, and in 2035 for the southern midlatitudes, and to fall well below 1980 values by 

the second half of the century. However, actual changes in surface UV will be affected by a number of factors 

other than ozone.

•		 In the absence of changes in other factors, changes in tropical surface UV will be small because changes in

tropical total column ozone are projected to be small. By the middle of the 21st century, the model projections 

suggest surface UV to be slightly higher than in the 1960s, very close to values in 1980, and slightly lower than 

in 2000. The projected decrease in tropical total column ozone through the latter half of the century will likely 

result in clear-sky surface UV remaining above 1960 levels. Average UV irradiance is already high in the tropics 

due to naturally occurring low total ozone columns and high solar elevations.
3.4

Chapter 3

•		 The magnitude of UV changes in the polar regions is larger than elsewhere because ozone changes in polar

regions are larger. For the next decades, surface clear-sky UV irradiance, particularly in the Antarctic, will 

continue to be higher than in 1980. Future increases in ozone and decreases in clear-sky UV will occur at slower 

rates than those associated with the ozone decreases and UV increases that occurred before 2000. In Antarctica, 

surface clear-sky UV is projected to return to 1980 levels between 2040 and 2060, while in the Arctic this is projected to occur between 2020 and 2030. By 2100, October surface clear-sky erythemal irradiance in Antarctica 

is likely to be between 5% below to 25% above 1960 levels, with considerable uncertainty. This is consistent 

with multi-model-mean October Antarctic total column ozone not returning to 1960 levels by 2100. In contrast, 

by 2100, surface clear-sky UV in the Arctic is projected to be 0–10% below 1960 levels.
3.5

Future Ozone and UV

3.1 INTRODUCTION

The primary goal of the Montreal Protocol and its 

Amendments and Adjustments is to avoid depletion of 

the ozone layer. Ozone depletion elevates surface ultraviolet (UV) radiation, thereby posing a threat to Earth’s 

biosphere. The mechanism used to achieve this goal is 

the control of the production and consumption of anthropogenic ozone-depleting substances (ODSs: see Box 3-1). 

Chapter 5 in this Assessment describes the Montreal Protocol mechanism in more detail while Chapters 1, 2 and 5 

summarize the success of the Montreal Protocol to date. 

At present, throughout most of the stratosphere, equivalent 

stratospheric chlorine (ESC; see Box 3-1) is significantly 

elevated above natural levels due to the cumulative effect 

of historical emissions of ODSs. Elevated ESC remains 

the most important anthropogenic perturbation to stratospheric ozone. However, as halogen loading declines in 

the future, other factors are expected to displace ODSs as 

the dominant influence on ozone. Perhaps the most important of these are continued and increasing emissions 

of long-lived greenhouse gases (GHGs). For example, 

GHGs affect ozone directly because they act as stratospheric source gases for ozone-destroying radicals (e.g., 

methane (CH4) and nitrous oxide (N2O)) and indirectly 

because they change temperatures (predominantly carbon 

dioxide (CO2)).

This chapter builds on and extends Chapter 5 of the 

2006 Ozone Assessment (“Climate-Ozone Connections,” 

Baldwin and Dameris et al., 2007) by assessing the most 

Box 3-1. Measures of Atmospheric Halogens

In this chapter, as in the 2006 Ozone Assessment (hereafter also cited as WMO (2007)), ozone-depleting substances (ODSs) are defined as those gases of anthropogenic origin controlled under the Montreal Protocol. Principal 

ODSs are chlorofluorocarbons (CFCs), hydrochlorofluorocarbons (HCFCs), halons, carbon tetrachloride, 1,1,1-trichloroethane (methyl chloroform), and methyl bromide. These gases are useful in meeting global application demands, for 

example, in refrigeration, air conditioning, insulating foams, and fumigation, but contain the halogens (chlorine and/

or bromine atoms) which, when released, react to destroy stratospheric ozone. Other gases emitted in human activities 

(e.g., methane (CH4) and nitrous oxide (N2O)) also influence ozone (Section 3.2.2), but have not been considered ODSs 

under the Montreal Protocol.

When chlorine and bromine atoms are released from the degradation of ODSs in the stratosphere, they combine 

to form the inorganic chlorine- and bromine-containing compounds that belong to the chemical groups called total inorganic chlorine (Cly) and inorganic bromine (Bry). The combination of Cly and Bry amounts represents the potential for 

halogens to destroy ozone in a stratospheric air mass. A measure of this potential is defined as equivalent stratospheric 

chlorine (ESC; Eyring et al., 2007), “equivalent” since it weights Bry with respect to Cly, according to

ESC = Cly + α × Bry (1)

where the units are stratospheric mixing ratio and α is the weighting factor that accounts for the greater effectiveness of 

bromine in ozone destruction compared to the effectiveness of chlorine on a per-atom basis. In general, α varies with altitude, latitude, and time, following changes in Cly and Bry which in turn follow total emissions of organic halogen gases 

at Earth’s surface. In 2010, α is estimated to be ~60 in the lower stratosphere and ~5 in the upper stratosphere (Sinnhuber 

et al., 2009). ESC can be calculated as a function of latitude, longitude, altitude, and time from distributions of Cly and 

Bry simulated with a chemistry-climate model (CCM). Some ODSs, such as methyl chloride and methyl bromide, and 

a number of other halogen-containing gases, have natural sources. These natural emissions are responsible for the low 

background level of ESC.

Deriving ESC directly from observations is hindered by the lack of direct measurements of Cly and Bry. As an 

alternative, the spatial and temporal distributions of Cly and Bry can be approximated from time series measurements 

of ODS surface concentrations combined with estimated rates at which individual gases release their halogens into the 

stratosphere (so-called fractional release rates) and estimates of the age of air parcels (i.e., the time elapsed since air 

parcels entered the stratosphere at the tropical tropopause; see Box 8-1 of WMO (2007)). Using these estimates in equation (1) yields equivalent effective stratospheric chlorine (EESC; see Section 1.4.4 of Chapter 1), which has been widely 

used before the availability of ESC from CCM outputs. ESC and EESC are similar in that they both represent aggregate 

quantities that quantify the combined potential of Cly and Bry in a particular air mass to destroy ozone. They differ in 

the way in which the Cly and Bry inputs are obtained. For CCMs, ESC is calculated directly from the simulated fields of 

chlorine- and bromine-containing species. In contrast, EESC is estimated from measured source gas surface concentrations and assumptions about their transport and conversion to Cly and Bry in the stratosphere.
3.6

Chapter 3

recent literature on the likely effects of increasing GHG 

emissions on future ozone amounts and resultant changes 

in climate parameters. This chapter also builds on and 

extends Chapter 6 of the 2006 Ozone Assessment (“The 

Ozone Layer in the 21st Century,” Bodeker and Waugh et 

al., 2007) by assessing chemistry-climate model (CCM) 

projections of the evolution of ozone through the 21st

century. The number of available CCMs and CCM sensitivity simulations has significantly increased since the 

previous Assessment. In contrast to WMO (2007), which 

assessed two- and three-dimensional model simulations 

of future ozone, this chapter assesses ozone projections 

from the current generation of three-dimensional CCMs. 

The simulations used in this chapter are based on those 

from 17 CCMs that participated in the second round of 

a coordinated model intercomparison organized by the 

Chemistry-Climate Model Validation (CCMVal) Activity (Eyring et al., 2005) of WCRP’s (World Climate Research Programme) SPARC (Stratospheric Processes And 

their Role in Climate) project (hereafter referred to as 

CCMVal-2). The participating CCMs are listed in Table 

3-1 (see Section 3.3.1) and described in detail in the cited 

literature, in Chapter 2 of SPARC CCMVal (2010), and 

in Morgenstern et al. (2010). An extensive assessment of 

these CCMs was made in the SPARC CCMVal Report 

(2010), which provides a strong foundation for this chapter. Results presented in that report are augmented here 

by including additional reference model simulations, as 

well as sensitivity simulations in which either greenhouse 

gas (GHG) or ODS levels are kept fixed at 1960 values or 

in which alternative GHG scenarios are used. The reference simulations include the most important forcings of 

the long-term evolution of ozone and provide the primary 

data set of ozone projections used here. The sensitivity 

simulations, by fixing a single forcing at a constant baseline level, permit a quantitative separation of the effects of 

different factors on future ozone. Thus details of the full 

recovery of ozone from the effects of ODSs (see Box 3-2) 

were determined by quantifying the contribution of ESC 

to future ozone changes and contrasting it with that from 

climate change.

Although interest in the factors affecting surface 

UV has focused primarily on ozone to date, other factors, 

some related to changes in climate, will play an increasingly important role in modulating future surface UV levels as the ozone layer recovers. This chapter builds on 

and extends the prognostic aspects of Chapter 7 of the 

2006 Ozone Assessment (“Surface Ultraviolet Radiation: 

Past, Present, and Future,” Bais and Lubin et al., 2007), 

with a primary focus on the role of ozone in affecting the 

future evolution of UV. While a brief discussion of the 

non-ozone factors affecting surface UV is included in this 

chapter, it primarily serves as a link to the UNEP (United 

Nations Environment Programme) Environmental Effects 

Assessment Panel report (UNEP, 2010).

While ESC is projected to return close to historical 

levels during the 21st century, global ozone amounts are 

not necessarily expected to return close to their respective 

historical levels over the same period because factors other 

than ESC will increasingly affect future ozone amounts. 

Therefore, this chapter reports both the dates when ESC, 

ozone, and UV return to some respective historical values, 

and the dates when ozone is no longer expected to be influenced by ODSs (Box 3-2). Annual mean total column 

ozone and surface UV may not necessarily return to historical levels at the same time. Summertime ozone exerts a 

greater influence on annual mean surface UV than ozone in 

other seasons, and so it is the return of summertime ozone 

to historical levels that is more relevant for surface UV.

The scope of this chapter includes providing a discussion framework that accommodates a wide variety of 

factors beyond ODSs that affect ozone. In some regions of 

the atmosphere, for example in the Antarctic lower stratosphere, the steep decline in springtime ozone until ~2000 

is projected to be followed by a slow increase back to 1980 

levels during the latter half of the 21st century, mirroring 

what is expected from changes in ESC over that period. In 

the tropical upper stratosphere, where the effects of GHGinduced cooling are significant, ozone concentrations 

could significantly exceed 1960 or 1980 values by the end 

of the 21st century. These changes are necessarily larger 

than what would be expected from ESC declines alone. In 

contrast, in the tropical lower stratosphere, ozone shows 

little sensitivity to stratospheric halogen loading and is 

more strongly influenced by long-term changes in atmospheric circulation. This chapter, in addition to addressing 

the full recovery of ozone from ODSs, explores a suite 

of scenarios for future ozone in different regions of the 

atmosphere, and the impact of projected ozone changes 

on surface UV.

Section 3.2 details how future changes in stratospheric halogen loading will affect the future evolution 

of ozone (the primary focus in this chapter), describes 

other factors that will affect future ozone, and discusses 

how ozone and other factors will affect future surface UV. 

Section 3.3 begins by noting the recent improvements in 

ozone projections compared to those presented in WMO 

(2007) and by describing the CCMs, the emissions scenarios to which they were applied, and their evaluation. 

Section 3.3 focuses mainly on long-term CCM projections of stratospheric ozone based primarily on a common 

reference scenario. In addition to this reference scenario, 

a number of sensitivity simulations allow the two most 

important factors affecting future ozone (i.e., ODSs and 

GHGs) to be quantitatively disentangled. These results 

form the basis for statements regarding the milestones of 
3.7

Future Ozone and UV

full ozone recovery from the effects of ODSs. Section 3.4 

builds on the ozone projections in Section 3.3 to detail how 

future surface UV radiation will be affected by changes 

in ozone and provides a point of connection between this 

Ozone Assessment and the UNEP Environmental Effects 

Assessment Panel report (UNEP, 2010).

3.2 FACTORS AFFECTING FUTURE OZONE 

AND SURFACE UV

This section discusses the factors that will affect the 

future evolution of stratospheric ozone and consequently 

UV reaching the surface. While the primary focus remains 

on ODS emissions and the associated increase in stratospheric halogen loading (Section 3.2.1), developments 

since WMO (2007) have highlighted the potentially important roles of nitrous oxide (N2O) emissions (Section 3.2.2) 

and anthropogenic enhancements of the stratospheric 

sulfate aerosol layer, from both increased surface sulfur 

emissions (Section 3.2.5) and possible geoengineering 

actions (Section 3.2.6). New research has also advanced 

our understanding of how the Brewer-Dobson circulation 

(BDC) is likely to change (Section 4.2.2 of Chapter 4 of 

this Assessment) and affect future ozone (Section 3.2.4.1). 

The increase in GHG emissions and subsequent changes 

in climate, including cooling of the stratosphere (Section 

3.2.3), is an overarching theme linking many of the factors likely to affect ozone in the future. Evaluating the 

recovery of ozone from ODSs involves all of these factors, 

in addition to ODSs. As ESC levels decline in the future, 

the factors that act to amplify the effects of ESC on ozone 

Box 3-2. Ozone Return Dates and Full Ozone Recovery

The context for the discussion of future ozone in this chapter is provided by defining two distinct milestones in the 

future evolution of ozone. These quantitative milestones are motivated by the need to answer questions often put forth 

by policymakers and the public, viz.:

1. When do we expect ozone to return to levels typical of some earlier time?

2. When do we expect ozone to no longer be significantly affected by ozone-depleting substances (ODSs)?

Identifying the first milestone requires no attribution to the separate factors affecting ozone and, hence, can be 

evaluated directly from time series of observed or simulated ozone. If ozone has already returned to levels typical of 

some target year, measurements alone can be used to address this milestone and answer the related question. Since this 

is not the case for the target years considered here (i.e., 1960 and 1980), the expected return of ozone to these historical 

levels is evaluated here instead from multi-model mean chemistry-climate model (CCM) projections. The selection of 

a target year, and its associated ozone level, should not be interpreted as selecting past states defined by the absence of 

significant ozone depletion from ODSs.

In this chapter a key target date is chosen to be 1980 in order to retain a connection to previous Ozone Assessments. Noting, however, that decreases in ozone may have occurred in some regions of the atmosphere prior to 1980, 

1960 return dates are also presented and discussed. Evaluating return dates is relevant for gauging when the adverse 

impacts of enhanced surface ultraviolet radiation on human health and ecosystems caused by ozone depletion are likely 

to become negligible. In addition, the return dates of ozone to historical levels in some regions are also valuable in demonstrating the effectiveness of policies that have abated anthropogenic ODS emissions if ODSs have been the dominant 

driver of ozone changes in those regions.

Identifying the second milestone is equivalent to assessing when the third stage of ozone recovery, so-called “full 

ozone recovery from ODSs” as defined in WMO (2007), has or is expected to occur. This, by definition, requires an 

attribution of projected changes in ozone to different factors, since ozone is not affected by ODSs alone. The required 

attribution can be obtained using idealized CCM simulations that hold certain model forcings fixed in conjunction with 

reference simulations that include the most important forcings for the long-term evolution of ozone. The reference and 

idealized simulations also provide a method, based on a subjective statistical test, for quantifying what is meant by “no 

longer significantly affected by ODSs” (Section 3.3.6). When analyzing CCM simulations for full ozone recovery from 

ODSs as defined here, the choice of 1960 as the start date implicitly assumes that the contribution of anthropogenic emissions of ODSs to equivalent stratospheric chlorine (ESC) prior to 1960 was negligible. However, ESC in 1960 was not 

zero because of the contribution of natural sources of halogens (see Box 3-1). If the sensitivity of ozone destruction to 

stratospheric halogen loading does not change with time, this milestone is equivalent to a return of ESC to 1960 levels. 

This definition has the advantage that it directly assesses the effectiveness of the Montreal Protocol in model projections 

of ozone but requires the additional step of attributing observed or modeled changes in ozone to ODS and other factors.

Both milestones defined above can be evaluated for total column ozone or vertically resolved ozone, averaged 

globally or averaged over some region of the atmosphere.
3.8

Chapter 3

(e.g., stratospheric aerosols) will decline in importance, 

while other factors, especially those related to GHGs and 

climate change, will emerge to dominate the long-term 

evolution of ozone. Attributing past and future changes 

in ozone to these different factors is essential to identify 

the stages of ozone recovery. Furthermore, because ozone 

protection policy governs only ESC (by limiting emissions 

of ODSs), evaluating the effectiveness of such policy requires attributing observed and projected ozone changes 

to ESC and other factors. While the discussion of UV in 

this section is primarily focused on the effects of future 

changes in stratospheric ozone on surface UV (Section 

3.2.7), research since WMO (2007) has mostly dealt with 

other factors affecting surface UV, which are summarized 

briefly in Section 3.2.8.

3.2.1 Stratospheric Halogen Loading

At present, and for much of the 21st century, elevated stratospheric halogen loading is expected to remain the 

most important factor affecting stratospheric ozone. In the 

model simulations of 21st century ozone made in support 

of WMO (2007), ODSs were prescribed at the surface according to the Ab “best guess” scenario of WMO (2003) 

(Eyring et al., 2007). This scenario was superseded by the 

“baseline” halogen scenario A1 defined in WMO (2007), 

which prescribes halogen loadings higher than in Ab. 

There are notable differences between the two scenarios. 

The A1 scenario includes larger emissions of chlorofluorocarbon-11 (CFC-11) and CFC-12 after 2010 due to their 

larger estimated banks. Increases in projected hydrochlorofluorocarbon-22 (HCFC-22) emissions, due to expectations of greater future use in A1 than in the Ab scenario, 

are also important. However, at the 2007 Meeting of the 

Parties to the Montreal Protocol, the Parties agreed to an 

earlier phase-out of HCFCs, with nearly a full phase-out 

in developing countries (Article 5) by 2030. Scenario A1 

does not include this phase-out. Hence, a new scenario 

has been developed that includes this phase-out. This 

adjusted A1 scenario has been used in the model simulations performed in support of the present Assessment. In 

this new scenario, only HCFCs have been adjusted, while 

emissions of CFCs, halons, and other non-HCFC species 

remain identical to the original A1 scenario.

Due to the uncertainty in their future trends, brominated very short-lived substances (VSLS, atmospheric 

lifetime <0.5 year) are not considered in the standard scenarios. As a result, the bromine loading in most CCMs 

used in support of the previous and current Assessments is 

only determined by the projected evolution of long-lived 

organic source gases methyl bromide (CH3Br) and halons 

(halon-1211, halon-1202, halon-1301, and halon-2402). 

However, observations suggest that the stratospheric bromine loading is 5+

–

3

2 parts per trillion (ppt) higher than can 

be explained by tropospheric levels of CH3Br and the 

halons. This additional bromine (~5 ppt out of the current total of ~20 ppt) may likely come from short-lived 

bromine-containing compounds (see Chapter 1; WMO, 

2007). The degradation of those species could release a 

substantial amount of bromine into the tropical tropopause 

layer (TTL) and lowermost stratosphere. The average 

photochemical lifetime of VSLS is comparable with, or 

shorter than, the average transport timescales in the TTL, 

which vary from rapid in deep convection to slow outside 

of regions of deep convection. This may lead to a highly 

variable distribution of bromine injections in the form of 

VSLS into the tropical lower stratosphere, as some surface and balloon observations suggest (see Section 1.3.1 

and references therein). There is also uncertainty in how 

dehydration and wet removal of degradation products of 

VSLS occur (Sinnhuber and Folkins, 2006). Current and 

future trends in brominated VSLS are difficult to determine. VSLS, and their rapid vertical transport through the 

troposphere, are also very difficult to describe in global 

models and are not explicitly treated in the CCMs used 

here. Together, these gaps in our understanding of the 

stratospheric bromine budget are a source of uncertainty 

in long-term projections of stratospheric halogen loading 

and hence ozone.

3.2.2 Stratospheric Reactive Nitrogen and 

Hydrogen Levels

Natural ozone loss in the absence of chlorine is primarily due to the chemical effects of nitrogen (NOx) and 

hydrogen (HOx) radicals. NOx and HOx levels are controlled by the amount of the long-lived source gases N2O, 

H2O, and CH4, which are also GHGs. Since the levels 

of these source gases are affected by anthropogenic emissions and/or climate change, their future evolution will 

influence ozone through changes in NOx and HOx.

The amount of NOx in the stratosphere is largely 

controlled by the amount of N2O entering at the tropical 

tropopause (of which approximately 10% is converted to 

NOx). The preindustrial level of N2O was ~270 parts per 

billion (ppb) and the present level is ~320 ppb. Future 

changes in ozone due to NOx increases have been evaluated using model simulations based on scenarios for the 21st

century that include projected N2O emissions (Randeniya 

et al., 2002; Chipperfield and Feng, 2003; Portmann and 

Solomon, 2007). It was shown that several percent of global total ozone loss is possible by 2100 due to increases in 

anthropogenic N2O. The effects of increasing N2O levels 

on ozone depend on altitude. The resulting NOx increase 

causes ozone losses in the middle stratosphere, centered 

just above the ozone concentration maximum, while chlorine and hydrogen radicals destroy ozone predominately 
3.9

Future Ozone and UV

in both the lower and upper stratosphere (see Figure 3-1). 

Interactions between halogen and nitrogen species cause 

nitrogen species to be less effective at destroying ozone 

in the lower stratosphere while halogen levels remain elevated. In contrast, temperature decreases cause a faster 

chemical loss of total reactive nitrogen (NOy) and hence 

NOx (nitrogen oxides, NO + NO2) in the upper stratosphere (Rosenfield and Douglass, 1998). Because of these 

effects, the changes in active nitrogen species levels and 

in the amount of ozone destroyed by those species are not 

expected to follow the change in N2O levels.

The efficiency of N2O emissions in destroying 

stratospheric ozone can also be compared to that of ODSs. 

Ravishankara et al. (2009) computed the Ozone Depletion 

Potential (ODP, see Chapter 5) of N2O in the same way 

that it is computed for halogen source gases. Surprisingly, 

they found that the ODP-weighted anthropogenic emissions of N2O were larger than those of any chlorinecontaining source gas emitted in 2008. Moreover, they 

also found that the ODP-weighted anthropogenic emissions of N2O were already significant in 1987 when CFC 

emissions were peaking. This result is partly due to the 

long lifetime of N2O (approximately 125 years) since the 

ODP is weighted by the ozone depletion over the lifetime 

of the gas. Further study is needed to characterize the 

effects of N2O emissions and carefully examine the 

tradeoffs with ODSs (see also Section 5.4.2.2).

Future levels of HOx will be mostly determined by 

changes in CH4 emissions and by the amount of H2O entering the tropical stratosphere. In terms of ozone depletion, these two sources of HOx have differing effects on 

ozone. Increasing H2O tends to enhance the chemical 

ozone loss in the upper stratosphere where H2O is converted to HOx (HOx is the dominant ozone loss cycle in 

the upper stratosphere and mesosphere), whereas increasing H2O leads to a reduction in ozone loss in the middle 

stratosphere because H2O enhances conversion of nitrogen dioxide (NO2), the ozone-destroying nitrogen radical, 

into HNO3. However, when all the radiative and dynamical feedbacks associated with the H2O increase are accounted for in a CCM, the effects on column ozone exhibit a strong hemispheric asymmetry, especially at high 

latitudes (Tian et al., 2009). Changes in the abundance of 

H2O also have the potential to affect the sulfuric acid aerosol size distribution. However, these H2O-driven aerosol 

changes have been estimated to be negligible (SPARC, 

2006). Polar stratospheric cloud (PSC) formation and the 

heterogeneous reaction rates on liquid ternary aerosol are 

more sensitive to H2O changes. H2O increases may be 

expected to enhance PSC formation and hence accelerate 

chemical polar ozone destruction as long as chlorine levels are high enough. CH4 increases also enhance H2O in 

the stratosphere but additionally cause ozone increases in 

the troposphere and lowermost stratosphere due to direct 

ozone production from methane oxidation. The net effect 

of increases in CH4 levels on global ozone is expected to 

be ozone production (Randeniya et al., 2002). Thus it 

is important to assess the sources of H2O changes in the 

stratosphere to predict the net effect on ozone.

H2O increased in the stratosphere in the latter part 

of the 20th century but showed a sustained decrease after 

2000 (Randel et al., 2006). Increases in CH4 levels have 

caused part of this H2O increase in the middle and upper 

stratosphere but not in the lower stratosphere (Rohs et al., 

2006). In addition, the separation of the climate change 

signal from natural variability has proved difficult (Garcia 

et al., 2007; Austin et al., 2007; Oman et al., 2008).

3.2.3 Stratospheric Temperatures

Because of the temperature dependence of gasphase chemical reaction rates and the formation of PSCs, 

stratospheric temperatures have a large impact on ozone 

abundances. A cooling of the middle and upper stratosphere increases ozone by slowing gas-phase destruction 

rates. In contrast, a cooling of the polar lower stratosphere 

is expected to enhance PSC formation which, when halogen levels are elevated, favors ozone destruction. In the 

Antarctic, where temperatures are already well below the 

Change in Ozone due to Perturbation

−0.15 −0.10 −0.05 −0.00 0.05

DU / km

0

10

20

30

40

50

Approximate Altitude (km)

CFC−11

N2O

Global Mean Ozone

0 5 10 15

0

10

20

30

40

50

Approximate Altitude (km)

DU / km

Figure 3-1. The change in global average ozone 

concentrations (Dobson units per kilometer) due to 

a 100 ppt increase in CFC-11 and a 20 ppb increase 

in N2O for year 2000 levels of source gases and 

background aerosol conditions computed with the 

NOCAR 2-D model (Portmann et al., 1999). The 

global average ozone profile for the same conditions 

is also shown. Note that in order to illustrate differences in the ozone loss profiles, the relative sizes of 

the CFC-11 and N2O perturbations were chosen to 

achieve ozone loss peaks of the same magnitude. 

Adapted from Figure S1 of the supplementary material from Ravishankara et al. (2009).
3.10

Chapter 3

thresholds of PSC formation, additional cooling is likely 

to have less of an effect than in the Arctic, where stratospheric temperatures are closer to the thresholds of PSC 

formation. Predictions of future changes in ozone must 

therefore account for future changes in temperature.

Stratospheric temperatures are controlled by a combination of radiative and dynamical processes (see also 

Section 4.3.1). Since regional heating and cooling by dynamical processes tend to cancel out in the global mean, 

the global mean temperature is in radiative equilibrium to 

a good approximation (e.g., Fomichev et al., 2002). Dynamical processes, on the other hand, lead to latitudinal 

variations in heating and cooling. For example, the projected strengthening of the BDC in response to increasing 

GHG levels (Section 4.3.2) causes adiabatic cooling of 

the tropical lower and middle stratosphere, and adiabatic 

warming at high latitudes (Section 3.2.4).

Current CCMs are reasonably good at reproducing 

the Stratospheric Sounding Unit/Mesospheric Sounding 

Unit (SSU/MSU) record of observed changes in stratospheric temperature since 1979 (Figure 4-10, and Section 

4.3; Gillett et al., 2010). The multi-model ensemble approximately captures the magnitude of observed global 

mean cooling in the upper and lower stratosphere, as well 

as the volcanic warming and the recent leveling off of 

lower stratospheric temperatures (Section 4.3.1). Nonetheless, discrepancies between individual models and the 

observations may still be significant, and there are some 

clear areas of disagreement between the models and observations, such as in SSU channel 26 after 1995. However, 

the uncertainties in the observations remain poorly quantified (Randel et al., 2009). Thus it is possible that this 

discrepancy results from observational errors rather than 

from a common bias in the CCMVal models. Over the 21st

century CCMVal-2 models simulate a continued strong 

cooling of the middle to upper stratosphere of 4–10 K in 

the tropics, due to increasing GHG concentrations (Figure 

3-2(e)). Weaker cooling is simulated lower down in the 

stratosphere, and little change in temperature is simulated 

in the lowermost stratosphere due to the competing effects 

of ozone recovery and continued GHG increases.

3.2.4 Transport and Dynamics

3.2.4.1 Brewer-Dobson Circulation

As described in Section 4.2.2, the BDC is the stratospheric overturning circulation that transports air upward 

in the tropics, poleward at midlatitudes, and downward 

at middle and high latitudes, and so plays a crucial role 

in determining the meridional distribution of ozone and 

long-lived trace gases. Climate models and CCMs consistently predict an acceleration of the BDC in response 

to climate change, producing a trend of ~2% per decade 

in net upward mass flux in the tropical lower stratosphere 

in the multi-model mean (Section 4.3.2). An acceleration 

of the BDC would increase the rate at which ODSs are removed from the stratosphere (Butchart and Scaife, 2001), 

thus advancing ozone recovery. However, this removal 

mechanism is not represented in current CCMs as a consequence of the lower boundary condition for ODSs where 

mixing ratios, not fluxes, are specified (Section 3.3.7). An 

accelerated BDC would also increase the rate at which 

other gases such as N2O and CH4 get into the stratosphere, 

which would in turn affect the lifetimes of these gases and 

the evolution of ozone. In addition, the projected increase 

in the strength of the BDC would decrease ozone in the 

tropical lower stratosphere where ozone-poor air of tropospheric origin enters and rises slowly in relative isolation 

within this region. During the transit through the tropical lower stratosphere, ozone is continuously produced 

by molecular oxygen (O2) photolysis within the rising air. 

As a result, tropical lower stratospheric ozone content is 

mostly determined by a balance between the rate of ozone 

production (i.e., from photolysis of O2) and the rate at 

which the air is transported through and out of the tropical 

lower stratosphere (essentially the rate of ascent and, to a 

lesser extent, mixing with the subtropics) (Avallone and 

Prather, 1996). A faster transit of air through the tropical lower stratosphere from an enhanced BDC would ultimately lead to less time for production of ozone and hence 

lower ozone levels in this region. In contrast, ozone levels 

would increase in the extratropical lower stratosphere due 

to increased downward transport of ozone-rich air from 

above. This latitudinal dependence is illustrated in Figure 3-3, which shows 1970 to 2090 changes in the residual vertical velocity w−* at 70 hPa and the corresponding 

changes in total column ozone from an ensemble of simulations using one CCM, the Canadian Middle Atmosphere 

Model (CMAM) (McLandress and Shepherd, 2009a). The 

close correspondence between changes in w−* and ozone (a 

negative correlation) highlights the potential importance 

of future changes in the strength of the BDC on ozone. 

However, the fact that the modest modeled increases in 

the strength of the BDC to date are difficult to detect in the 

currently available measurements (Section 4.2.2) suggests 

that some caution should be exercised when diagnosing 

projected future changes in ozone in the light of expected 

changes in the BDC.

3.2.4.2 Vortex Integrity and Mixing

Future changes in the shape and strength of the polar 

vortex and in the frequency and strength of sudden stratospheric warmings (SSWs) will likely alter meridional transport which will in turn affect polar ozone. Charlton-Perez 

et al. (2008) diagnose a significant increase in major SSW 
3.11

Future Ozone and UV

(a)

(d)

(e) (f)

(b)

(c)

O3 (ppm) O3 (ppm)

O3 (ppm)

O3 Temp. (K) (ppm)

 Cly + αBry (ppb)

100

10

1

p (hPa)

100

10

1

p (hPa) 100

10

1

p (hPa) 100

10

1

p (hPa)

100

10

1

p (hPa) 100

10

1

p (hPa) Figure 3-2. Vertical profile results of the Multiple Linear Regression (MLR) analysis for the CCMVal-2 models 

(described in Table 3-1) in the latitude band 10°S–10°N. (a) Ozone in the year 2000 (parts per million) versus 

pressure (hectoPascals), (b) Ozone change from 2000 to 2100, (c) ESC change (parts per billion) from 2000 

to 2100, (d) ESC-congruent ozone change. (e) temperature change and (f) temperature-congruent ozone 

change. From Chapter 9 of SPARC CCMVal (2010), their Figure 9.5.
3.12

Chapter 3

frequency of 1 event per decade between 1960 and 2099 

in the Atmospheric Model with TRansport and Chemistry 

(AMTRAC) using a crossing method based on transition 

from westerlies to easterlies to detect major SSWs. However, using a criterion for major SSW occurrence based on 

the amplitude of the Northern Annular Mode and simulations from the CMAM, McLandress and Shepherd (2009b) 

showed that the future increase in the frequency of major 

SSWs in that model was a consequence of changes in the 

underlying climatology (i.e., mean vortex strength) and did 

not mean an increase in stratospheric variability. A recent 

study by Bell et al. (2010), however, found an increase in 

both stratospheric variability and major SSW frequency in 

simulations using enhanced GHG concentrations. Clearly, 

a consensus on the impact of stratospheric climate changes 

on vortex variability has not yet been reached, and cleanly 

separating changes in vortex variability from changes in 

the mean vortex strength remains a challenge.

The isolation of the polar vortex in conjunction with 

mixing across the vortex edge is an important dynamical 

regulator for ozone amounts at high latitudes (e.g., Strahan and Polansky, 2006) and midlatitudes (Braesicke and 

Pyle, 2003; Hadjinicolaou and Pyle, 2004; Wohltmann et 

al., 2007). Assessments of observed and modeled changes in stratospheric mixing are sparse. Garny et al. (2007) 

diagnosed recent trends in stratospheric mixing on selected isentropic surfaces in the lower stratosphere and found 

substantial differences for different heights, horizontal 

regions, and seasons, with, for example, long-term positive trends in mixing in southern midlatitudes at 450K 

nearly year-round and negative trends in southern high 

latitudes at 650K from May to August. In models, the 

choice of vertical resolution (Rind et al., 2007) and advection algorithm (Stenke et al., 2008) plays an important 

role. Modeling the 2002 major warming in the Southern 

Hemisphere, Konopka et al. (2005) diagnosed a larger degree of isolation than for similar events in the Northern 

Hemisphere. Changes in Antarctic vortex isolation at the 

end of the 21st century have been examined for a subset of CCMs (Chapter 5 of the SPARC CCMVal Report, 

2010). Some of the models analyzed indicate increased 

mixing between the vortex and midlatitudes above 1000K 

(~35 km) in the future, suggesting that winter planetary 

wave activity may have increased in the models. In this 

model intercomparison, no consensus could be found as 

to how the Antarctic vortex size and depth may change 

toward 2100. Note that the link between the modeled 

strengthening of the BDC and changes in mixing is not 

straightforward. Therefore, based on current knowledge, 

we cannot judge with high confidence how mixing across 

vortex barriers will change in the future and how this will 

affect ozone.

3.2.5 Background and Volcanic 

Stratospheric Aerosols

A layer of sulfuric acid aerosol is present at all 

latitudes in the lower stratosphere. During volcanically 

quiescent periods (also called “background” conditions), 

the dominant source of this aerosol layer is thought to be 

carbonyl sulfide photolysis and, possibly, sulfur entering the stratosphere in the form of sulfur dioxide (SO2) 

(SPARC, 2006). The other major source of sulfur to the 

stratosphere is volcanoes. Chapters 3 and 6 of WMO 

(2007) (Chipperfield and Fioletov et al., 2007; Bodeker 

and Waugh et al., 2007) contain an in-depth discussion 

of volcanic aerosols and their effects on ozone. Chapter 8 of the CCMVal report (SPARC CCMVal, 2010) 

contains recent modeling intercomparisons of volcanic 

effects on ozone.

Volcanic eruptions can inject large amounts of 

sulfur directly into the stratosphere. This can considerably enhance the stratospheric aerosol layer (or Junge 

layer) for several years. Such an effect was observed 

after the eruptions of El Chichón in 1982 and Mt. 

Delta Vertical Velocity

J F M A M J J A S O N D

Month

-90

-60

-30

0

30

60

90

Latitude (deg)

J F M A M J J A S O N D

-90

-60

-30

0

30

60

90

-0.10

0.00 0.00

0.00

0.00

0.10

(a) Delta Column Ozone

J F M A M J J A S O N D

Month

-90

-60

-30

0

30

60

90

J F M A M J J A S O N D

-90

-60

-30

0

30

60

90

-20

0

0

0

20

40

(b) Figure 3-3. Difference between 

future and past (a) residual vertical velocity at 70 hPa and (b) 

total column ozone (deviation 

about global mean). Contour 

intervals are 0.05 mm/s and 10 

DU. Future and past time periods are computed using 20-year 

means centered about 2090 

and 1970, respectively. From 

McLandress and Shepherd 

(2009a), adapted from panels 

(b) and (f) of their Figure 6.
3.13

Future Ozone and UV

Pinatubo in 1991. The enhanced stratospheric aerosol 

loading can lead to very significant ozone depletion on 

a global scale, as was both observed and modeled after 

these eruptions. The ozone depletion is mostly due to 

heterogeneous reactions on sulfuric acid aerosol particles that convert halogen reservoir species into more reactive forms. The overall chemical changes include an 

increase in halogen radicals at the expense of nitrogen 

radicals. Thus, the net effect on ozone depends primarily on the stratospheric halogen loading, changing for 

instance from ozone decreases when chlorine loading 

is high to small ozone increases when chlorine loading 

is low (Tie and Brasseur, 1995). As a result, the potential for ozone depletion from enhanced aerosol loading is expected to decline as ESC decays toward natural 

levels late in the 21st century. The future recovery of 

ozone would be temporarily interrupted by large volcanic eruptions, especially during the first half of the 

21st century. Small volcanic eruptions can also influence ozone, as was found in aircraft measurements that 

probed the Hekla, Iceland, volcanic plume in the lowermost stratosphere (Millard et al., 2006). The volcanic 

plume contained high SO2 levels and very low ozone 

levels, near zero in places, and was still 30% depleted 

two weeks after. Simulations using a three-dimensional 

(3-D) chemistry-transport model (Chipperfield, 2006) 

showed that increased heterogeneous activation of chlorine due to elevated H2O and HNO3 from the volcanic 

plume was the likely cause of the ozone loss. Events 

such as these would not be expected to cause widespread 

ozone loss because of both the small size of the eruption and the short residence time of volcanic aerosols 

in the lowermost stratosphere. The analyses of these 

events do, however, confirm the link between aerosols 

and ozone depletion.

By virtue of their optical properties, aerosols also 

impact the radiative balance of the atmosphere and, 

hence, can affect stratospheric temperatures and dynamics substantially. Since the previous Assessment, several 

studies have attempted to separate the chemical and dynamical signals in the ozone response following the Mt. 

Pinatubo eruption. Using a two-dimensional model 

forced with meteorological analyses, Fleming et al. 

(2007) showed that, while the chemical ozone destruction due to halogen chemistry on volcanic aerosols took 

place in the lower stratosphere in both hemispheres, dynamical effects acted to decrease (increase) total column 

ozone in the Northern (Southern) Hemisphere. This result was confirmed by Telford et al. (2009) using a CCM 

nudged toward meteorological analyses. They found 

that, in addition to the ~10 Dobson unit (DU) chemical 

ozone loss in both hemispheres following the Mt. Pinatubo eruption, the quasi-biennial oscillation (QBO) (see 

Chapter 2) increased ozone by ~10 DU in the Southern 

Hemisphere and decreased ozone by ~10 DU in the 

Northern Hemisphere. These two modeling studies confirm the earlier work of Hadjinicolaou et al. (1997) that 

demonstrated the existence of a large dynamical effect 

on ozone following the eruption.

From long-term aerosol measurements, it has been 

suggested that background stratospheric aerosol levels 

have been increasing in the last decade possibly due to enhanced tropospheric SO2 background (see Sections 2.4.3.3 

and 4.1.4 in Chapters 2 and 4 of this Assessment). There 

are a few reasons why a change in background aerosol levels could significantly affect ozone while chlorine levels 

remain elevated. First, halogen radical levels (and thus 

ozone loss) increase more strongly with aerosol loading 

when the aerosol loading is small, because the halogen activation saturates at high aerosol loadings due to the reduction of dinitrogen pentoxide (N2O5). Second, depending 

on the cause of this observed aerosol increase, it is possible that the increases may be larger at lower altitudes, 

where halogen radicals impact ozone most strongly. Thus, 

the potential for significant changes in background aerosol 

levels remains an uncertainty in predicting ozone changes 

in the coming decades.

3.2.6 Geoengineering by Sulfate 

Aerosol Injection

Doubts regarding the effectiveness of current international agreements to restrict emissions of GHGs to 

the atmosphere to mitigate climate change have led to a 

debate about the possibility of intentionally modifying 

climate through large-scale geoengineering actions; see 

Crutzen (2006) and a series of replies in a special issue 

of Climatic Change (vol. 77, no. 3-4, 2006). Such climate engineering schemes, if ever applied, might impact 

stratospheric ozone (Royal Society, 2009). Proposed 

schemes and uncertainties in their implementation are 

discussed in Chapter 5 of this Assessment (Section 

5.4.2.4). Present knowledge strongly suggests that the 

addition of sunlight-reflecting sulfuric acid aerosols into 

the stratosphere would have large impacts on the future 

evolution of stratospheric ozone and as such should be 

considered alongside the other factors discussed in this 

chapter.

As discussed in the previous section, observations 

show a marked decrease in global ozone following the 

explosive volcanic eruptions of El Chichón in 1982 and 

Mt. Pinatubo in 1991. The primary reason for this enhanced ozone loss is the enhanced activation of stratospheric chlorine on volcanic aerosol particles (Kinnison 

et al., 1994; Solomon et al., 1996; Portmann et al., 1996; 

Tilmes et al., 2008b). Similar effects are expected for 

geoengineered stratospheric aerosol enhancements. An 
3.14

Chapter 3

illustration of the impact of an enhanced stratospheric 

aerosol layer on chemical ozone loss is shown in Figure 3-4 (Tilmes et al., 2008a). For large, geoengineered, 

stratospheric aerosol loadings that appear to be required 

for significant climate change mitigation, model estimates suggest that chlorine activation could be more than 

doubled in the Arctic if it is implemented in the next 20 

years, resulting in Arctic ozone depletion of 200–230 DU 

during very cold winters (comparable to the total amount 

of available ozone in the Arctic lower stratosphere). As 

stratospheric halogen levels decline over the 21st century, 

the impact of an enhanced aerosol layer on chlorine activation would be reduced but would still be significant. 

The expected recovery of the Antarctic ozone hole could 

be delayed by between 30 to 70 years depending on the 

assumed geoengineered aerosol size distribution (Tilmes 

et al., 2008a).

Further understanding of the consequences of 

geoengineering for stratospheric ozone and climate has 

been gained from numerical model simulations initially 

using global climate models (Matthews and Caldeira, 

2007; Caldeira and Wood, 2008; Rasch et al., 2008; Robock et al., 2009) and, more recently, CCMs (Tilmes et 

al., 2009; Heckendorn et al., 2009). The CCM investigations confirm the empirical results of Tilmes et al. (2008a) 

that stratospheric sulfate injection could enhance stratospheric ozone depletion in both the Arctic and Antarctic 

while the levels of halogens remain elevated. Tilmes et 

al. (2009) quantified the impact of an enhanced burden of 

stratospheric sulfur on the ozone layer including the impacts on stratospheric dynamics and transport during the 

period when the stratospheric halogen loading is projected 

to slowly decline (Newman et al., 2007). Model calculations with a fixed enhanced sulfate aerosol loading (large 

enough to counteract the forcing generated by a doubling 

of CO2 abundance with respect to preindustrial values) 

predict a one- to two-fold increase in Arctic ozone depletion due to a stronger polar vortex and lower temperatures 

(Rasch et al., 2008). The impact on midlatitude ozone was 

calculated to be smaller for 2050 chlorine levels (Tilmes et 

al., 2009). An additional risk is the possible occurrence of 

a large volcanic eruption, further increasing the amount of 

sulfur in the stratosphere (Tilmes et al., 2009).

Significant uncertainty remains regarding the behavior of artificially injected stratospheric aerosol, since 

the coagulation and settling of aerosol is strongly dependent on the injection scenario. Aerosol heating, in particular at the tropical tropopause, could also result in enhanced 

amounts of water vapor entering the stratosphere (Heckendorn et al., 2009), which would cause stratospheric ozone 

loss. According to CCM simulations of predicted 2050 

conditions (Tilmes et al., 2009; Heckendorn et al., 2009), 

a geoengineered enhanced aerosol loading would lead to 

substantial and coupled changes in stratospheric dynam0

50

100

150

200

Ozone Loss (DU)

Cold Arctic

approx. CO2 doubling

 Volcanic Aerosols 

 Small Aerosols 

 Observations 

 Background Case 

1980 2000 2020 2040 2060 2080 2100

Year

0

50

100

150

200

Ozone Loss (DU)

Moderately Cold Arctic

approx. CO2 doubling

Figure 3-4. An illustration of the possible impact of 

geoengineering by stratospheric injection of sulfate 

aerosols on Arctic chemical ozone loss. The panels 

show estimates of the potential ozone loss (Dobson 

units) from chlorine activation as a function of time for 

different stratospheric aerosol loadings and for two 

types of Arctic winters. The top panel corresponds 

to cold Arctic winters, conditions that are expected 

to lead to the maximum impact of geoengineering 

on ozone depletion. The bottom panel is for moderately cold Arctic winters, representative of about 

half of the past 15 Arctic winters. In each panel, the 

solid and dotted black lines represent the ozone loss 

for a background stratospheric aerosol loading and 

observed stratospheric aerosol loading respectively. 

The red line shows the ozone loss estimated when 

2 teragrams per year (Tg/yr) of stratospheric sulfur 

are added to the stratosphere, starting in 2010 and 

reaching a saturation value of 5.3 Tg of sulfur in the 

form of large volcanic-sized aerosol particles. The 

blue line corresponds to an alternative geoengineering scenario assuming the formation of smaller aerosol particles. In this scenario, a smaller injection of 

1.5 Tg of sulfur per year would achieve the same radiative effect necessary to counteract the impact of 

a doubling of atmospheric carbon dioxide concentrations. Adapted from Tilmes et al. (2008a).
3.15

Future Ozone and UV

ics and ozone chemistry, especially in halogen heterogeneous chemistry. To date there have been no investigations of the impact on ozone caused by a gradual ramp-up 

of the amount of SO2 injected, with the purpose of keeping global average temperature nearly constant (Wigley, 

2006). Overall, we caution that these ozone projections 

are based on simple scenarios, idealized conditions, and 

on only two CCMs, including one without any aerosol 

microphysics. Because the ozone response depends on 

complex and competing interactions between chemical, 

aerosol microphysical, radiative, and dynamical processes 

that are represented differently in different models, more 

reliable ozone projections in a hypothetical geoengineered 

climate state require more realistic simulations with a 

range of models.

3.2.7 Effects of Ozone on Future Surface 

UV

In the 1980s and 1990s, a number of theoretical and 

experimental studies were devoted to quantifying ozone 

effects on UV spectral irradiance reaching the ground (e.g., 

Brühl and Crutzen, 1989; Schwander et al., 1997; Tsay 

and Stamnes, 1992; Madronich et al., 1998; Lapeta et al., 

2000). Figure 3-5 shows the dependence of erythemal irradiance (СIE, 1993) on total ozone and other atmospheric 

parameters (cloud optical thickness, surface albedo, and 

aerosol optical thickness) within the range of values that 

can be observed in real atmospheric conditions. To quantify UV radiation, we use here the “erythemally weighted 

irradiance,” or simply “erythemal irradiance,” which is a 

measure of the biological effectiveness of solar UV radiation incident per unit area of human skin and is commonly 

used for public information and awareness. The effective 

wavelengths of erythemal irradiance lie between 307 and 

325 nm, depending mainly on solar zenith angle and total 

ozone. Other effects on humans and ecosystems are more 

sensitive to shorter or longer wavelengths of UV radiation, hence ozone-induced changes in the corresponding 

weighted irradiances would be either larger (e.g., for DNA 

damage) or smaller (e.g., for phytoplankton damage) compared to erythemal irradiance.

To estimate the response of UV radiation to ozone 

variations, the radiation amplification factor (RAF) approach proposed by Booth and Madronich (1994) is 

widely used. It provides simple but useful estimates of 

different biological weightings of UV irradiance response 

to ozone changes. For erythemal irradiance, the RAF was 

shown to be about −1.1 (WMO, 1999). At the same time, 

the influence of ozone on UV irradiance depends on solar 

elevation. Recently, a simple equation was proposed by 

Madronich (2007) enabling the calculation of erythemal 

irradiance (or the UV Index) under cloud-free skies and 

low surface-albedo conditions as a function of total ozone 

(in the range 200–400 DU), taking into account also the 

effect of solar zenith angle. The uncertainty of these estimates increases for solar zenith angles larger than ~60°. 

However, the application of these approaches can lead to 

random and/or systematic uncertainties in calculating the 

UV response to future total ozone changes because they 

do not account for possible changes in ozone and temperature profiles. The influence of aerosols, clouds, and surface albedo are also neglected.

Figure 3-5. Relative changes (percent) in 

erythemal irradiance due to total ozone, 

cloud optical thickness, surface albedo, and 

aerosol optical thickness (AOT) at 308 nm 

for two values of single scattering albedo 

(SSA), calculated for 50° solar zenith angle 

with the Tropospheric Ultraviolet-Visible 

(TUV) v.2 model (Madronich and Flocke, 

1999). All parameters vary within the range 

observed in real atmospheric conditions. 

Except for the effect of cloud optical thickness, all calculations refer to cloud-free 

skies. For the calculation of erythemal irradiance changes due to ozone, typical values were used for the aerosol optical thickness, AOT = 0.31, (Kinne et al., 2006) and 

the single scattering albedo, SSA = 0.94 

(Chubarova, 2009). For estimating the effects from aerosols, surface albedo, and cloud optical thickness, an 

ozone column of 300 DU was used in the model simulations. Change in Erythemal Irradiance (%)
3.16

Chapter 3

The effects of the vertical distribution of ozone in 

the troposphere and the stratosphere, as well as the temperature profiles (which affect the ozone cross-sections) 

should be taken into account when calculating the UV 

response to ozone changes (Brühl and Crutzen, 1989; 

Schwander et al., 1997; McKenzie et al., 2003). As discussed in Tsay and Stamnes (1992) and Brühl and Crutzen 

(1989), transport of ozone from the stratosphere to the troposphere tends to decrease UV at the surface, but for low 

solar elevation angles an increase may occur. Changes 

in ozone vertical distribution and temperature profile can 

modify erythemal irradiance by as much as 14%, according to Lapeta et al. (2000). In Kazantzidis et al. (2005) it 

was emphasized that the most significant effects of changes in vertical ozone distribution on surface erythemal irradiance are observed at large solar zenith angles (up to 20% 

at 85° solar zenith angle). However, UV levels are usually 

very small for these conditions.

The future evolution of ozone will determine to a 

great extent future levels of surface UV radiation and the 

date of return to the UV levels in the 1980s or 1960s. However, in some regions UV radiation levels in the future may 

not return to historical levels due to influences from factors 

related to climate change (e.g., cloudiness, surface albedo, 

aerosols, UV-absorbing tropospheric gases; see Section 

3.2.8). A detailed analysis of future UV levels due to ozone 

in different geographical areas is given in Section 3.4.

3.2.8 Factors Other Than Stratospheric 

Ozone Affecting Surface UV

In addition to ozone, UV radiation is affected by 

other atmospheric parameters, for example, changes in 

cloudiness, aerosols, surface albedo, and, to some extent, 

by other mineral and organic gas species. Since WMO 

(2007) a number of publications have discussed and quantified the factors other than ozone that affect surface UV 

irradiance (e.g., Kazadzis et al., 2007; Tanskanen and 

Manninen, 2007; Lindfors and Arola, 2008; Staiger et al., 

2008; Badosa et al., 2007; Chubarova, 2008; Rieder et al., 

2008; McKenzie et al., 2008; Chubarova et al., 2009). A 

comparison of the relative effects of changes in ozone and 

in other atmospheric parameters on surface erythemal irradiance is shown in Figure 3-5. Changes in cloud optical 

thickness, typically between 0 and 40, are the most important driver of day-to-day and long-term changes in surface 

UV irradiance, generally dominating the effect of changes 

in total column ozone. Aerosols, and in particular highly absorbing aerosols, also affect surface UV irradiance. 

However, for typical values of aerosol optical thickness 

(~0.3), the effect is small compared to clouds. Increases 

in surface albedo, for example due to changes in snow 

or ice cover, can significantly enhance surface UV. The 

sensitivity of surface UV to the changes in atmospheric 

parameters shown in Figure 3-5 does not account for more 

complicated, nonlinear interactions in their effects (for 

example, the interaction of clouds and inhomogeneous 

surface albedo, or cloud and aerosol interactions) or for 

spatial inhomogeneities in the atmosphere.

Since surface UV is appreciably sensitive to a number of factors other than ozone, projections of these factors, in addition to ozone, are required to make reliable and 

robust projections of surface UV irradiance. However, 

at present, projections of atmospheric factors other than 

ozone have large uncertainties, making accurate projections of surface UV elusive.

3.2.8.1 Clouds

Because climate change is likely to affect future 

cloudiness, and given the sensitivity of surface UV to 

changes in cloudiness (Figure 3-5), understanding potential future changes in cloud cover is essential to quantifying future changes in UV. To this end, projections of 

changes in cloudiness obtained from the Third Coupled 

Model Intercomparison Project (IPCC, 2007) and indirectly from Chapter 10 of SPARC CCMVal (2010) have 

been used. Climate models forced according to the Special Report on Emissions Scenarios (SRES) A1B scenario 

from IPCC (2000) (which is the standard scenario of the 

stratospheric ozone projections discussed here; see Section 3.3.1) predict with some consistency that by the end 

of the 21st century cloud cover would have decreased over 

most of the low and middle latitudes, and substantially increased at high latitudes. Trenberth and Fasullo (2009) 

examined top-of-atmosphere radiation changes in climate 

model projections and also found decreases in cloudiness 

through the 21st century predominantly after 2040. Coverage of optically thick low-level clouds is projected to 

also decrease over low and middle latitudes (Trenberth 

and Fasullo, 2009), in agreement with the tendency for 

low-level clouds to dissipate as the ocean warms (Clement 

et al., 2009). In Chapter 10 of SPARC CCMVal (2010), 

cloud projections were used to estimate future changes 

in surface erythemal irradiance by converting shortwave 

cloud transmittance to erythemal UV cloud transmittance. 

Shortwave cloud transmittances evaluated from the IPCC 

(2007) climate models and from Chapter 10 of SPARC 

CCMVal (2010) CCMs agree reasonably well. However, 

these projections are highly uncertain as the cloud response to climate change appears to be the primary source 

of spread between climate model simulations (Dufresne 

and Bony, 2008).

Projections of surface UV irradiance that include 

the effects of cloud changes result in more complex patterns in projected UV compared to clear-sky UV projections. Increases in erythemal irradiance of 10–15% by the 
3.17

Future Ozone and UV

end of 21st century due to changes in cloudiness are projected for the tropical regions of South-East Asia and Central America, with more moderate increases over Southern Europe in summer (Chapter 10 of SPARC CCMVal, 

2010). A UV reduction of 10–15% is projected by 2100 

due to increases in cloudiness over some northern highlatitude regions and over Antarctica.

3.2.8.2 Aerosols

The effect of atmospheric aerosols on surface UV 

radiation depends on their optical and microphysical properties and total atmospheric loading. While the tropospheric burden of sulfate aerosols is generally projected to 

decrease in the future, projections for black carbon (soot) 

are less certain (IPCC, 2007). Projections of aerosol properties have been aided by recent modeling studies (Kinne et 

al., 2006; Schulz et al., 2006). However, at present, despite 

general agreement on projections of annual mean aerosol 

optical thickness, the disparity of aerosol absorption in 

various models leads to large uncertainties in surface UV, 

precluding robust assessments of future aerosol effects on 

UV irradiance. There are also large uncertainties in the 

methods used to generate emissions scenarios as well as in 

assessing the present-day emissions, especially for black 

carbon and organic carbon (IPCC, 2007). Furthermore, it 

remains unresolved whether emissions of soil dust aerosols 

will increase or decrease in response to changes in the atmospheric state and circulation (Tegen et al., 2004). Differences in emission regulation strategies between countries result in large spatial variability in aerosol trends.

Climate model simulations project increases in 

aerosol optical thickness of ~1.4% by 2030 for the A1B 

scenario but decreases of ~5% for the SRES B1 scenario, 

a lower emission scenario (IPCC, 2000; Jacobson and 

Streets, 2009). Resultant changes in surface UV radiation 

are estimated to be smaller than −0.5% for the A1B scenario, and about 0.1–1.5% for the B1 scenario, depending 

on the aerosol properties and solar elevation.

3.2.8.3 Surface Albedo and Sea Ice Cover

The effects of surface albedo on UV irradiance 

are well documented (WMO, 1999; WMO, 2003; WMO, 

2007). Under clear-sky conditions, the presence of snow/

ice may increase UV by up to 50%, while in overcast or 

close to overcast conditions with optically thick clouds, 

the UV increase due to high snow/ice surface albedo can 

reach several hundred percent because of effective multiple scattering. On the other hand, sea ice cover significantly reduces penetration of harmful UV-B irradiance in 

the underwater environment. As a result, possible climate 

warming-induced changes in snow and sea ice cover may 

affect both terrestrial and aquatic ecosystems.

Reductions in sea ice extent through the 21st century in both the Arctic and Antarctic have been projected 

by a number of models, albeit with a rather large range of 

model responses in the Northern Hemisphere sea ice extent, ranging from very little change to a strong and accelerating reduction over the 21st century (Zhang and Walsh, 

2006). According to Overland and Wang (2007), the projected summer loss in sea ice extent would be greater than 

40% by 2050 for the marginal seas of the Arctic basin. 

Stroeve et al. (2007) reported that sea ice in the Arctic 

is melting faster than projected in IPCC (2007) under all 

SRES emissions scenarios. The accelerated decrease in 

Arctic sea ice extent has also been documented in satellite data (Comiso et al., 2008). This sea ice melting is 

projected to decrease the surface albedo and increase UV 

irradiance in the underwater environment. In the 20th- and 

21st-century simulations, Antarctic sea ice cover is projected to decrease more slowly than in the Arctic (Stroeve 

et al., 2007), particularly in the vicinity of the Ross Sea, 

where most models predict a local minimum in surface 

warming (Meehl et al., 2007).

It is well established that mountain glaciers and 

snow cover have declined on average in both hemispheres 

(IPCC, 2007). They are expected to continue declining 

in the future (Bradley et al., 2004). This would result in 

considerable local reductions in surface albedo and UV in 

the corresponding regions.

3.2.8.4 Tropospheric Gases

Several tropospheric gases, such as SO2, NO2, and 

ozone, can efficiently absorb UV-B radiation. Their influence depends mainly on their total column amount, which 

can be especially high over industrial areas and over areas 

influenced by forest fires. Their impact is also determined 

by the effectiveness of UV absorption discussed in Chapter 

7 of WMO (2007) (Bais and Lubin et al., 2007). The average effect of increasing NO2 on erythemal irradiance has 

been shown to be about −2% in industrial areas (Chubarova, 

2008). However, on some days the influence can be much 

stronger. For example, in Tokyo the amount of NO2 can 

be 20 times higher than the average, decreasing UV-B 

irradiance by about 15% (McKenzie et al., 2008). Similar 

effects from NO2 have been observed during intense forest 

fires, when erythemal irradiance has been attenuated by 

10–15% (Chubarova et al., 2009). SO2 column amounts 

can easily reach 2 DU over areas affected by high volcanic 

activity or over regions close to coal burning industries, 

resulting in ~2% attenuation of erythemal irradiance. At 

some UV wavelengths, the reduction by SO2 can exceed 

20% (McKenzie et al., 2008).

The amounts of these gases in the troposphere 

depend strongly on their emissions. In industrialized 

regions, such as North America and Europe, emissions 
3.18

Chapter 3

of NOx and volatile organic compounds are decreasing, 

while in regions dominated by developing countries, significant growth in emissions is observed (IPCC, 2007). 

Since 1980, SO2 emissions in 25 countries in Europe have 

been reduced by more than a factor of 4, while in the USA 

they have been halved (IPCC, 2007). However, over the 

same period, SO2 emissions from Asia and from developing countries in other regions have been increasing. There 

is a clear positive trend in tropospheric ozone concentrations projected by models under the A2p scenario, with 

increases between 11.4 and 20.5 DU by 2100 (Gauss et 

al., 2003). Taking into account that ozone absorbs UV 

radiation more effectively in the lower troposphere due to 

enhanced scattering, these increases in tropospheric ozone 

may result in approximately a 5–10% reduction in surface 

erythemal irradiance. The modeling study of Jacobson 

and Streets (2009) also projects by 2030 an increase in 

surface ozone of ~14% under the A1B scenario and ~4% 

under the B1 scenario.

3.3 PROJECTIONS OF OZONE THROUGH THE 

21ST CENTURY

Since the 2006 Assessment (WMO, 2007), a new 

suite of 17 chemistry-climate model (CCM) simulations 

coordinated through the SPARC CCMVal activity has become available. These simulations, from the second phase 

of CCMVal (referred to as CCMVal-2), form the basis for 

this section and improve on the CCMVal-1 simulations 

reported in WMO (2007) by:

• Starting in 1960 rather than in 1980. In most regions 

of the atmosphere, ozone depletion occurs prior to 

1980 in nearly all of the CCMs assessed here. By 

starting the simulations in 1960 at a time when ozone 

was not expected to be significantly affected by ODSs, 

and including sensitivity simulations, the CCMs can 

now be used to project the timing of the third stage of 

ozone recovery, i.e., the full recovery of ozone from 

the effects of ODSs. Stage-three ozone recovery was 

not reported on in WMO (2007). Furthermore, the 

availability of pre-1980 model data permits a more 

robust calculation of the 1980 ozone threshold and 

hence a more accurate determination of when ozone 

returns to 1980 levels.

• Having available almost all simulations from 1960 

to 2100. In WMO (2007) only one CCM provided 

simulations beyond 2050. The extension to 2100 now 

permits more robust conclusions to be drawn regarding the expected evolution of ozone through the latter 

half of the 21st century.

• Applying a more rigorous statistical analysis. The 

larger number of models, the availability of continuous simulations from 1960 to 2100 by nearly all 

CCMs, and a new method for calculating and analyzing multi-model time series (described in Section 

3.3.2.2) have allowed a more robust analysis than was 

presented in WMO (2007).

• Having available a number of sensitivity simulations with fixed forcings. In addition to the reference simulations as reported in WMO (2007), simulations where either surface concentrations of ODSs 

or GHGs are held fixed at their 1960 levels (Section 

3.3.1) permit a more in-depth analysis of the factors 

affecting ozone through the 21st century than previously possible and allow the assessment of full ozone 

recovery from ODSs.

• Having available simulations based on different GHG 

emissions scenarios. WMO (2007) reported only on 

simulations based on the SRES A1B scenario.

3.3.1 Model Descriptions and Scenarios

In this chapter the focus is on “future” simulations 

to 2100, whereas the “past” simulations (from 1960 up to 

2006) are discussed in Chapter 2. Ozone projections and 

the attribution of ozone changes to ODSs and GHGs are 

based on the recently completed CCMVal-2 multi-model 

ensemble. These simulations have been extensively analyzed in SPARC CCMVal (2010), as well as in a variety 

of individual model studies. In addition to this large ensemble of future reference simulations from 17 CCMs, 

in which realistic scenarios of ODSs and GHGs are used, 

several more specialized sensitivity simulations by a subset of CCMs (Waugh et al., 2009; Eyring et al., 2010a; 

Eyring et al., 2010b; Charlton-Perez et al., 2010; Oman 

et al., 2010) are assessed here.

The CCM simulations are all transient simulations 

in which ozone responds interactively to the secular 

trends in GHGs, ODSs, and in other boundary conditions. They are commonly separated into “past” (or “historical”) transient reference simulations that are driven 

by observed forcing and “future” transient reference simulations that are forced by trace gas projections and generally use modeled sea surface temperatures (SSTs) and 

sea ice concentrations (SICs). The CCMVal-2 “past” 

reference simulation (REF-B1) is designed to reproduce 

ozone changes from 1960 to the recent past (2006) when 

global ozone observations are available. It includes solar 

and volcanic forcings, and SSTs/SICs from observations. 

It allows a detailed investigation of the role of natural 

variability and other atmospheric changes important for 
3.19

Future Ozone and UV

ozone trends (see Chapter 2). The CCMVal-2 “future” 

reference simulation (REF-B2) is a self-consistent simulation from the past into the future. In this simulation 

the surface time series of halocarbons is based on the 

adjusted A1 halogen scenario (WMO, 2007; see also 

Section 3.2.1). The long-lived GHG concentrations are 

taken from the SRES A1B scenario (IPCC, 2000). SSTs 

and SICs are generally prescribed from coupled ocean 

model simulations. Of the 17 CCMs that provided ozone 

projections analyzed in this chapter, only CMAM was 

coupled to an interactive ocean model (see Table 3-1). 

Some CCMs generated an ensemble of future simulations with the same boundary conditions but different 

initial conditions (see Table 3-1). In general, the ozone 

variability between ensemble members from a single 

model is much smaller than the inter-model differences 

(Chapter 6, WMO, 2007; Austin et al., 2008).

Additional sensitivity simulations based on different emissions scenarios were also performed by some 

CCM groups to attribute the future evolution of stratospheric ozone to ODS and GHG forcings and to study the 

coupled chemistry-climate system under a variety of GHG 

scenarios (see Tables 3-1 and 3-2). These include sensitivity simulations with ODSs fixed at 1960 levels (fODS) 

to assess the milestone of full ozone recovery (Waugh et 

al., 2009; Eyring et al., 2010a) (see Section 3.3.6). In addition, sensitivity simulations with GHGs fixed at 1960 

levels (fGHG) were performed, for example, to address 

the issue of the linear additivity of the effects of GHGs 

and ODSs on ozone. By comparing the sum of the ozone 

responses in the fixed GHG and ODS simulations (each 

relative to the 1960 baseline) with the ozone response in 

the REF-B2 reference simulation, the linear additivity of 

the responses can be assessed (McLandress et al., 2010; 

Eyring et al., 2010a). A subset of four CCMs also provided future projections under GHG scenarios different 

to SRES A1B (Oman et al., 2010; Eyring et al., 2010b). 

These GHG sensitivity simulations (GHG-x) include simulations forced with the SRES A2 and B1 GHG scenario 

from IPCC (2000) and with the new Representative Concentration Pathways (RCPs; Moss et al., 2008) that form 

the basis for the climate simulations of the Coupled Model 

Intercomparison Project, Phase 5 (CMIP5, Taylor et al., 

2009), in support of the Fifth IPCC Assessment Report. 

They are generated by integrated assessment models and 

harmonized with the historical emissions from Lamarque 

et al. (2010) in both amplitude and geographical distribution. The RCP simulations performed by CAM3.5 are 

RCP 8.5 (Riahi et al., 2007), RCP 4.5 (Clarke et al., 2007), 

and RCP 2.6 (van Vuuren et al., 2007), where the number after “RCP” indicates the radiative forcing in W/m2

reached by 2100 in each scenario. By 2100, for example, 

CO2 in the RCP 8.5 and SRES A2 scenarios is ~200 ppm 

and ~100 ppm higher than in SRES A1B, while in the 

SRES B1/RCP 4.5 and RCP 2.6 scenario it is ~150 ppm 

and ~250 ppm lower, respectively. The different levels of 

GHGs lead to differences in ozone projections (see Section 3.2 on factors affecting future ozone).

3.3.2 Model Evaluation and Multi-Model 

Mean Analysis

3.3.2.1 Model Evaluation

Confidence and guidance in interpreting CCM projections of future changes in atmospheric composition 

can be gained by first ensuring that the CCMs are able 

to reproduce key processes for stratospheric ozone (e.g., 

Eyring et al., 2005; SPARC CCMVal, 2010). Limitations 

and deficiencies in the models can be revealed through 

inter-model comparisons and through comparisons with 

observations. An improvement over the approach used in 

CCMVal-1 and WMO (2007) is that for the current Assessment, a more extensive set of ozone-related processes 

was evaluated in the CCMs (SPARC CCMVal, 2010). In 

this chapter the evolution of ozone and inorganic chlorine 

(Cly) is shown for individual CCMs in several figures to 

portray the full distribution of model simulations. However, the discussion and conclusions are based on the time 

series of the multi-model mean and associated statistical 

confidence and prediction intervals (see Section 3.3.2.2).

A detailed summary of the key findings of SPARC 

CCMVal (2010) on the evaluation of CCMs against observations is presented for chemical composition in Chapter 2, and 

for the Brewer-Dobson circulation (BDC) and temperature 

in Chapter 4. This section briefly summarizes relevant conclusions from these chapters, with a focus on processes and 

results that are important for long-term ozone projections.

•	 BDC and Temperature. Both are important drivers 

of the evolution of ozone. Most CCMs are capable 

of reproducing the amplitude and vertical structure 

of the observed trends in global-mean stratospheric 

temperatures, although the model spread is high in 

some regions. Tropical upwelling is well simulated 

in the lower stratosphere compared to meteorological 

analyses. Models consistently predict a strengthening of the BDC and hence a decrease in mean age of 

air as a result of climate change, but they disagree on 

the relative role of resolved and parameterized wave 

drag. This strengthening of the BDC is partly supported by several lines of observational evidence (see 

Section 4.2.2) but not by recent estimates of the age 

of air inferred from tracers observations in the northern midlatitude lower stratosphere, which indicate a 

statistically insignificant trend (Engel et al., 2009). 

However, the small model-simulated trends in the 
3.20

Chapter 3

BDC lie within the large uncertainties of the stratospheric age-of-air measurements.

•	 Tropical and Midlatitude Ozone. In the tropics, climate change and halogen loading influence total column ozone through dynamical processes in the lower 

stratosphere and chemical processes in the upper stratosphere. Over the historical period, the CCMs simulate 

negative trends in tropical upper stratospheric ozone, in 

agreement with observations (see Section 2.4.5.3). The 

Table 3-1. A summary of the CCMs and the simulations used in this chapter. REF-B2 is the future reference simulation, fODS is a simulation with fixed ODSs, fGHG a simulation with fixed GHGs, and GHG-x are 

simulations with a GHG scenario different to SRES A1B; see details in Section 3.3.1 and Table 3-2. N × REF 

means that the group provided N realizations of this simulation. Further details on the models can be found 

in Morgenstern et al. (2010) and in Chapter 2 of SPARC CCMVal (2010) as well as in the references given 

below. EMAC-FUB and NIWA-SOCOL did not contribute a REF-B2 simulation for SPARC CCMVal (2010), but 

provided simulations later (Austin et al., 2010b; Eyring et al., 2010a).

CCM * Group and

Location **

Horizontal

Resolution

Upper

Level

Reference

Simulation fODS fGHG GHG-x References

AMTRAC3 GFDL, USA ~200 km 0.017 hPa REF-B2 --- --- --- Austin and 

Wilson (2010)

CAM3.5 NCAR, USA 1.9° ×

2.5° 3.5 hPa REF-B2 --- ---

RCP2.6

RCP4.5

RCP8.5

Lamarque et al. 

(2008)

CCSRNIES NIES, Tokyo, 

Japan T42 0.012 hPa REF-B2 fODS fGHG SRESB1

SRESA2

Akiyoshi et al. 

(2009)

CMAM

MSC, Univ. of

Toronto, York 

Univ., Canada

T31 0.00081 

hPa

3 ×

REF-B2

3 ×

fODS

3 ×

fGHG ---

Scinocca et 

al. (2008); de 

Grandpré et al. 

(2000)

CNRM-ACM Meteo-France, 

France T63 0.07 hPa REF-B2 --- --- ---

Déqué (2007); 

Teyssèdre et 

al. (2007)

E39CA DLR, Germany T30 10 hPa

REF-B2 

(with solar 

cycle and 

QBO)

--- fGHG ---

Stenke et al. 

(2009); Garny 

et al. (2009)

EMAC-FUB FU Berlin, 

Germany T42 0.01 hPa

REF-B2 

(with solar 

cycle and 

QBO)

--- fGHG ---

Jöckel et al. 

(2006); Nissen 

et al. (2007)

GEOSCCM NASA/GSFC, 

USA 2° × 2.5° 0.015 hPa REF-B2 fODS --- SRESA2 Pawson et al. (2008)

LMDZrepro IPSL, France 2.5° ×

3.75° 0.07 hPa REF-B2 fODS --- --- Jourdain et al. 

(2008)

MRI MRI, Japan T42 0.01 hPa 2 ×

REF-B2 fODS fGHG ---

Shibata and 

Deushi (2008a; 

2008b)

NIWASOCOL

NIWA, New

Zealand T30 0.01 hPa REF-B2 --- --- ---

Schraner et 

al. (2008); 

Egorova et al. 

(2005)
3.21

Future Ozone and UV

SOCOL

PMOD/WRC; 

ETHZ,

Switzerland

T30 0.01 hPa 3 ×

REF-B2

fODS fGHG ---

Schraner et 

al. (2008); 

Egorova et al. 

(2005)

ULAQ Univ. of 

L’Aquila, Italy

R6 / 11.5°

× 22.5° 0.04 hPa REF-B2 fODS fGHG ---

Pitari et al. 

(2002); Eyring 

et al. (2006; 

2007)

UMSLIMCAT Univ. of

Leeds, UK

2.5° ×

3.75° 0.01 hPa REF-B2 fODS ---

Tian and 

Chipperfield 

(2005); Tian et 

al. (2006)

UMUKCAMETO MetOffice, UK 2.5° ×

3.75° 84 km REF-B2 --- --- ---

Morgenstern 

et al. (2008; 

2009)

UMUKCAUCAM

Univ. of 

Cambridge, UK

2.5° ×

3.75° 84 km REF-B2 --- --- ---

Morgenstern 

et al. (2008; 

2009)

WACCM NCAR, USA 1.9° ×

2.5°

5.9603 ×

10−6 hPa 

3 ×

REF-B2 fODS fGHG SRESB1 Garcia et al. (2007)

* CCM acronyms are defined in Appendix B of this Assessment.

** GFDL, Geophysical Fluid Dynamics Laboratory (NOAA); NCAR, National Center for Atmospheric Research; NIES, National Institute for Environmental Studies; MSC, Meteorological Service of Canada; DLR, Deutschen Zentrum für Luft- und Raumfahrt; FU, Freie University; NASA, 

National Aeronautics and Space Administration; GSFC, Goddard Space Flight Center; IPSL, Institut Pierre-Simon Laplace; MRI, Meteorological 

Research Institute; NIWA, National Institute of Water and Atmospheric Research; PMOD, Physical-Meteorological Observatory-Davos; WRC, 

World Radiation Center; ETHZ, Swiss Federal Institute of Technology-Zürich. 

Table 3-1, continued.

Table 3-2. Summary of CCMVal-2 reference and sensitivity simulations used in this chapter.

Simulation

Name Period GHGs ODSs SSTs/SICs Background &Volcanic Aerosol

Solar

Variability QBO

REF-B2

Transient 

simulation 

1960–

2100

SRES A1B 

(medium)

(from IPCC, 

2000)

OBS + 

adjusted 

A1 scenario 

(WMO 2007, 

Table 8-5)

Modeled SSTs 

and SICs

OBS 

Background

surface area 

density from 

2000

No

Only 

internally 

generated

fODS

Fixed ODSs

1960–

2100

Same as in 

REF-B2

ODSs fixed at 

1960 levels

Same as in 

REF-B2

Same as in 

REF-B2

Same as in 

REF-B2

Same as in 

REF-B2

fGHG

Fixed GHGs

1960–

2100

GHG fixed

at 1960

levels

Same as in 

REF-B2

1955–1964 

average of REFB2, repeating 

each year

Same as in 

REF-B2

Same as in 

REF-B2

Same as in 

REF-B2

GHG-x

(SRES A2 and 

B1, RCP 2.6, 

4.5, and 8.5

2000–

2100

GHG scenario 

different from 

SRES A1B

Same as in 

REF-B2

SSTs/SICs distribution consistent with GHG 

scenario

Same as in 

REF-B2

Same as in 

REF-B2

Same as in 

REF-B2
3.22

Chapter 3

multi-model mean over the historical period indicates 

a small negative trend in tropical total column ozone, 

with a rather large model range. This small negative 

modeled trend is difficult to verify against observations 

since their length is limited and natural variability is 

high. Over midlatitudes, the multi-model mean adequately reproduces the negative trends in total column 

ozone although there is significant model spread.

•	 Polar Ozone. Both models and observations indicate 

that Antarctic stratospheric ozone loss, together with 

increasing GHG concentrations, has led to a poleward 

shift and strengthening of the Southern Hemisphere 

westerly tropospheric jet during summer. Most CCMs 

adequately represent lower stratospheric Antarctic vortex isolation, although some have deficiencies with respect to specific chemical or dynamical polar processes. 

Overall, they tend to reproduce well the Antarctic ozone 

losses inferred from observations; however it should be 

noted that many models show a late final warming in the 

Southern Hemisphere. This will extend chemical ozone 

loss later into the Antarctic spring. In the Arctic, while a 

few models represent the chemical ozone loss observed 

over the past three decades, most underestimate the loss, 

mainly because they tend not to capture the low temperatures observed in the Arctic lower stratosphere.

The multi-model mean estimates of past ozone 

changes simulated by CCMs under the REF-B1 scenario 

are generally consistent with the observed changes (see 

Chapter 2 of SPARC CCMVal, 2010). Overall, there 

is sufficient agreement among the CCMs, and between 

CCMs and observations, on the underlying causes of the 

ozone changes so that general conclusions can be drawn 

and some confidence placed in the CCM projections.

3.3.2.2 Analysis Method for Multi-Model Time

Series

Chapter 9 of SPARC CCMVal (2010) introduced 

a time series additive model (TSAM) to analyze the 

CCMVal-2 multi-model projections. This method is used 

to calculate baseline-adjusted anomaly time series relative 

to the values at a particular reference year (here 1960 and 

1980) for ozone and other species for each model simulation. The values at the reference year are obtained from 

a smooth fit to the model time series calculated with the 

TSAM method. This smooth fit is referred to as the individual model trend estimate. Note that the term “trend” 

does not denote the result of a linear regression analysis 

but rather refers to a smooth trajectory passing through the 

data. The multi-model trend estimate is the average of the 

individual model trend estimates. By definition, both the 

individual model trends and the multi-model trends pass 

through zero at the specified reference year. Two types 

of uncertainty intervals are constructed. The first is the 

point-wise 95% confidence interval. This interval has a 

95% chance of overlapping the true trend and represents 

the local uncertainty in the trend at each year. The second 

interval is the 95% prediction interval which, by construction, is larger than the confidence interval. The 95% prediction interval is a combination of the local uncertainty in 

the trend and uncertainty due to natural interannual variability about the trend; it gives a sense of where an ozone 

value for a given year might reasonably lie. Both the confidence and prediction intervals are time varying. 

The credibility of simulated ozone projections is 

linked to a realistic representation of processes that drive 

stratospheric ozone. Process-based performance metrics 

have been used to assess the ability of CCMs to reproduce 

key processes for stratospheric ozone and its impact on 

climate (SPARC CCMVal, 2010). While it could be preferable to use metric-based weightings when calculating 

multi-model means, SPARC CCMVal (2010) concluded 

that more analysis is needed to assess the robustness and 

interpretation of performance metrics, and their possible 

use in assigning relative weights to ozone projections. For 

this reason, and to remain consistent with the multi-model 

mean time series presented in SPARC CCMVal (2010), 

the multi-model mean ozone projections from the 

CCMVal-2 simulations are shown here without applying 

weights. The robustness of the CCMVal-2 multi-model 

mean ozone projections and uncertainties is demonstrated 

by the fact that it is generally insensitive to whether outliers are included or not (Waugh and Eyring, 2008; Chapter 

9 of SPARC CCMVal, 2010). Indeed, the multi-model 

means of total column ozone calculated from simulations 

of the four CCMs selected in Chapter 2 as the highestscoring models are found to be very close to the multimodel means calculated from all 17 CCMs.

3.3.3 Tropical Ozone

3.3.3.1 Long-Term Projections of Tropical

Ozone

The 1960 to 2100 annual mean tropical (25°S–25°N) 

total column ozone time series are shown in Figure 3-6(a) 

for the REF-B2 simulations (Table 3-2). The simulated decrease in annual mean total column ozone of 8 DU from 

1960 to 2000 is larger than observed. Following the simulated minimum in tropical ozone around 2000, by 2025 

about 70% of the ozone lost since 1980 is projected to have 

been replenished and by 2050 ozone levels are projected to 

be at or very slightly above 1980 values (see Table 3-3). In 

the latter half of the 21st century, the annual mean tropical 

column ozone is projected to decline despite the projected 
3.23

Future Ozone and UV

Table 3-3. Summary of the extent to which total column ozone is projected to have returned to 1960 and 1980 levels from its absolute minimum. A value of 0% denotes that ozone has not increased above the minimum, 50% denotes 

that ozone at this date is halfway between the simulated minimum and the 1960 or 1980 level, 100% denotes that ozone 

has returned to the 1960 or 1980 level, and >100% denotes that ozone exceeds the 1960 or 1980 level at this date. The 

reference year and the total column ozone in that year are listed in the second column. The year in which the minimum 

ozone occurred and the total column ozone at the minimum are listed in the 3rd column. In each of the cells in subsequent 

columns, the value obtained from the multi-model trend estimate is listed in boldface on the center line, with the 95% confidence intervals extracted from the TSAM statistics listed on the lines above and below. All ozone values have been rounded 

to the nearest DU and all percentage values to the nearest 5%.

Region

Reference

Year;

Total

Column

Ozone (DU)

Year When

Minimum

Occurs;

Total

Column

Ozone (DU)

Difference to

Reference

Year, in 2025

Difference to

Reference

Year, in 2050

Difference to

Reference

Year, in 2075

Difference to

Reference

Year, in 2100

Unit (DU) (%) (DU) (%) (DU) (%) (DU) (%)

Global 

annual 

mean

1960

(312 DU) 2001 

(296 DU)

−10

−8

−6

40%

50%

65%

−2

−1

1

90%

95%

105%

1

3

5

105%

115%

130%

2

4

6

115%

125%

140%

1980

(306 DU)

−4

−2

0

60%

80%

100%

3

5

7

130%

150%

170%

7

8

10

170%

180%

200%

8

10

11

180%

190%

210%

Tropics 

annual 

mean

1960

(272 DU) 2000

(265 DU)

−5

−4

−3

30%

45%

55%

−3

−2

−1

55%

70%

85%

−4

−3

−2

45%

60%

70%

−7

−6

−4

0%

25%

45%

1980

(270 DU)

−3

−1

0

40%

70%

100%

−1

0

2

80%

110%

140%

−2

−1

1

60%

90%

120%

−4

−3

−2

20%

40%

60%

Northern

midlatitude 

annual 

mean

1960

(357 DU) 2000

(344 DU)

−4

−2

−1

70%

85%

90%

5

7

10

140%

155%

175%

11

14

16

185%

205%

225%

15

18

21

215%

240%

260%

1980

(353 DU)

0

3

5

100%

130%

155%

9

11

14

200%

230%

255%

15

18

21

265%

300%

335%

19

22

25

310%

355%

380%

Southern 

midlatitude 

annual 

mean

1960

(349 DU) 2002

(322 DU)

−18

−15

−12

35%

45%

55%

−5

−2

1

80%

90%

105%

3

6

9

110%

120%

135%

6

9

12

120%

130%

145%

1980

(339 DU)

−8

−5

−2

55%

70%

90%

5

7

11

130%

145%

165%

13

16

19

175%

190%

210%

16

19

22

195%

210%

230%

Antarctic 

October 

mean

1960

(374 DU) 2003

(244 DU)

−100

−95

−85

20%

30%

35%

−61

−53

−44

55%

60%

65%

−35

−26

−18

75%

80%

85%

−18

−9

1

85%

95%

100%

1980

(322 DU)

−50

−42

−34

35%

45%

55%

−10

−1

8

85%

100%

110%

17

25

34

120%

130%

145%

33

43

52

140%

155%

165%

Arctic 

March 

mean

1960

(459 DU) 2002

(422 DU)

−19

−15

−10

50%

60%

75%

2

6

11

105%

115%

130%

18

22

27

150%

160%

175%

27

33

38

175%

190%

205%

1980

(445 DU)

−6

−1

3

75%

95%

115%

15

20

24

165%

185%

205%

31

36

40

235%

255%

275%

41

46

52

290%

300%

325%
3.24

Chapter 3

monotonic decline in 50 hPa tropical Cly through the 21st

century (Figure 3-6(b)), with 1980 Cly values reached 

around 2040. Over the entire 1960 to 2100 period, secular 

variations in tropical column ozone are only ~10 DU.

Austin et al. (2010a; see also Chapter 9 of SPARC 

CCMVal, 2010) compared these future projections with 

those assessed in WMO (2007) and noted that there was 

little change, at least up to 2050. However, they found 

NIWA

(DU) (DU) (DU)

(ppb) (ppb) (ppb)
3.25

Future Ozone and UV

reduced uncertainty in the projections, mainly because of 

a greater number of simulations covering the whole period from 1960 to 2100. Nonetheless, as with the CCM 

projections considered in WMO (2007), there was a wide 

spread among the tropical ozone amounts simulated by 

the individual models and these extended significantly 

above and below the observed values. Apart from two 

models, there were no notable changes in the model total 

column ozone biases in this region (Chapter 9 of SPARC 

CCMVal, 2010). WMO (2007) did not report on the 

projected behavior of tropical total column ozone after 

2050, as only one CCM simulated that period (Bodeker 

and Waugh et al., 2007). On the other hand, the projections from that single model, shown in Figure 6-10 of 

WMO (2007), do not show the late 21st century decline 

in total column ozone found in the latest multi-model 

projections.

Oman et al. (2010) found that using the SRES A2 

scenario rather than the SRES A1B in the GEOSCCM 

significantly increased the abundance of reactive nitrogen 

and hydrogen in the upper stratosphere. However, any 

increased chemical ozone destruction from the additional 

NOy and HOx was largely mitigated by the additional GHG 

cooling in the SRES A2 scenario, which slows gas-phase 

ozone destruction. As a result, the ozone evolution in the 

upper stratosphere was similar for both scenarios. Results 

from the CCSRNIES model confirm this finding (Eyring 

et al., 2010b). In general, Eyring et al. (2010b) found that 

the projected behavior of tropical total column ozone was 

not very sensitive to the range of different GHG scenarios 

(SRES A1B, A2, and B1; RCP 2.6, 4.5 and 8.5; see Section 3.3.1). By 2100 the differences among scenarios was 

only ~4 DU, which is small compared to the differences 

found in the extratropics (see Section 3.3.4).

3.3.3.2 Processes Determining Future Tropical

Ozone

Chapter 9 of SPARC CCMVal (2010) performed 

a multiple linear regression on the time series of tropical 

ozone in the current CCMs and found that between 2000 

and 2100 the change in ozone was quite different above and 

below ~20 hPa (see Figure 3-2(b)). In the upper stratosphere 

ozone is projected to increase, with the largest increase of 

around 1.5 ppm occurring near 3 hPa for the multi-model 

mean. In contrast, lower stratospheric ozone is projected 

to decrease, consistent with earlier studies (WMO, 2007; 

Eyring et al., 2007). In both regions the rate of change in 

ozone is fairly constant through the 21st century (see for example Figure 9.4 of SPARC CCMVal, 2010). The projected 

increase in upper stratospheric ozone was found to be related mainly to a decrease in halogen concentrations (see Figures 3-2(c) and (d)) and GHG-induced cooling (see Figures 

3-2(e) and (f)). These two mechanisms made roughly equal 

contributions to the ozone increase over the 21st century under the SRES A1B GHG and A1 adjusted halogen scenarios 

(Chapter 9 of SPARC CCMVal, 2010). In WMO (2007) 

and in the studies of Shepherd and Jonsson (2008), Jonsson 

et al. (2009), and Waugh et al. (2009), similar conclusions 

were reached based on single models. Figure 3-7 shows the 

results of a regression of global mean ozone concentration 

onto changes in ODS and CO2 heating rates in simulations 

of the CMAM (Jonsson et al., 2009). Note that the change 

in ozone associated with future changes in GHG amounts is 

expected to be somewhat larger than that associated with future temperature changes, since some future GHG-induced 

cooling is balanced by ozone-induced warming (Shepherd 

and Jonsson, 2008). In the tropical lower stratosphere, the 

primary mechanism causing long-term changes in ozone is 

the increase in tropical upwelling through the 21st century, 

which is a robust feature in the CCM simulations (Chapters 

Figure 3-6 (at left). 1980 baseline-adjusted multi-model trend estimates of annually averaged total column ozone 

(DU; left) and Cly at 50 hPa (ppb; right) for the tropics (25°S–25°N, upper row) and midlatitudes (middle row: 35°N–

60°N, lower row: 35°S–60°S) (thick dark gray line) with 95% confidence and 95% prediction intervals appearing as 

light- and dark-gray shaded regions, respectively, about the trend (note the different vertical scale among the panels). 

The baseline-adjusted individual model trends are also plotted (colored lines). The red vertical dashed line indicates 

the year when the multi-model trend in total column ozone (left) and Cly at 50 hPa (right) returns to 1980 values and 

the blue vertical dashed lines indicate the uncertainty in these return dates. The black dotted lines in the left panels 

show observed total column ozone, where a linear least squares regression model was used to remove the effects 

of the quasi-biennial oscillation, solar cycle, El Niño-Southern Oscillation, and volcanoes from four observational data 

sets. 2σ uncertainties on the observations were derived by applying the regression model to 10,000 statistically 

equivalent time series obtained from Monte Carlo resampling of the regression model residuals. The observations 

include ground-based measurements (updated from Fioletov et al., 2002), merged satellite data (Stolarski and Frith, 

2006), the National Institute of Water and Atmospheric Research (NIWA) combined total column ozone database 

(Bodeker et al., 2005), and Solar Backscatter Ultraviolet (SBUV, SBUV/2) retrievals (updated from Miller et al., 2002). 

The observational time series is shifted vertically so that it equals 0 in 1980. See Table 3-1 for model descriptions. 

Redrawn from Figures 9.2, 9.7, 9.8, and 9.9 of SPARC CCMVal (2010) and updated with two new CCM simulations.
3.26

Chapter 3

4 and 9 of SPARC CCMVal 2010; Eyring et al., 2010a). 

This increase in upwelling in the tropics ultimately leads to 

a decrease in ozone levels in the tropical lower stratosphere, 

where ozone levels are mostly determined by a balance between the rate of ozone production and the ascent rate (see 

Section 3.2.4.1). Because GHG increases are projected to 

increase ozone in the tropical upper stratosphere, but decrease ozone in the tropical lower stratosphere, if or when 

ozone returns to historical levels (e.g., to the amounts of 

ozone observed in 1960 or 1980) will vary between these 

altitudes (see Section 3.3.6).

The attribution of ozone changes in the tropical upper and lower stratosphere, as well as in the total column, 

to changes in ODSs and GHGs from four CCMs analyzed 

by Eyring et al. (2010a) is shown in the correlative time 

series plots (i.e., plots showing the temporal evolution of 

the correlation between ozone and ESC) in Figure 3-8. 

The construction and evaluation methods for this figure 

and Figure 3-10 are described in Appendix 3A. The reference simulations show ozone decreasing from 1960 to 

2000 in the upper stratosphere in response to increasing 

ESC. However, as ESC decreases from 2000 to 2100 

ozone does not simply retrace the 1960 to 2000 path but is 

systematically elevated through the 21st century such that 

ozone returns to 1980 values in the late 2020s, well before 

ESC returns to its 1980 value in the mid-2050s. The elevated ozone through the 21st century results from GHGinduced stratospheric cooling (see Section 3.2.3) indicated 

by the red to blue transition from 1960 to 2100 in the reference trace in Figure 3-8(a), with a possible contribution 

from GHG-induced changes in transport. The fixed ODS 

simulation shows ozone in the upper stratosphere slowly 

increasing with time under the influence of GHG-induced 

stratospheric cooling. In contrast to the reference simulation, the fixed GHG simulation shows that the response 

of ozone to ESC through the 21st century is almost identical to that through the 20th century. In this simulation, 

because GHGs are fixed, temperatures show almost no 

25

30

35

40

45

50

Log−pressure height (km)

Ozone trend: 2010−2040

Ozone Trend (%/decade)

Pressure (hPa)

−8 −6 −4 −2 0 2 4 6

100

101

ODS

CO2

Total

Figure 3-7. Attribution of future (2010–2040) global 

mean ozone changes over 50–0.5 hPa using simulations from the Canadian Middle Atmosphere Model 

(CMAM). The simulated ozone mixing ratio trend 

(%/decade) is shown in black while the contributions 

from CO2 and ODSs changes to the ozone trend are 

shown in green and blue, respectively (note the ODS 

contribution is estimated from the upper stratospheric 

Cly). The gray shaded areas indicate the 99% confidence intervals for the linear fits to the ensemble 

average time series. The green and blue shaded 

regions indicate the uncertainty in the CO2 and ODS 

attribution estimates, derived from the 99% confidence intervals for the fitted CO2, ODS, ozone, and 

temperature linear trends. Adapted from Jonsson et 

al. (2009).

Figure 3-8 (at right). Correlative time series plots for 5 hPa, 50 hPa, and column ozone amounts (parts per million or Dobson units, left axis; ppm or DU change with respect to 1960, right axis) averaged between 25°S and 

25°N as extracted from the CCM multi-model trend time series (see Appendix 3A-1). (a) Annual mean tropical 

ozone as a function of ESC = Cly + 5×Bry (in parts per billion) at 5 hPa. (b) As in panel (a) but at 50 hPa and 

with ESC = Cly + 60×Bry. In panels (a) and (b) the REF, fixed ODSs (fODS), and fixed GHGs (fGHG) simulations (see Table 3-2) are shown using traces colored according to the multi-model-mean temperature using the 

scale shown in the bottom left of each panel. The gray traces in these two panels show the additive effects of 

the fODS and fGHG simulations calculated from: GrayESC(t) = fGHGESC(t) + fODSESC(t) – fODSESC(1960) and 

Grayozone(t) = fGHGozone(t) + fODSozone(t) – fODSozone(1960). Differences between the gray and REF traces indicate a lack of linear additivity in the system. Panel (c), as in (b) but for total column ozone and without color 

coding by temperature. In this panel the linear additivity test trace is shown in yellow (yellow = blue + green). 

In all three panels, on each trace, reference years are shown every 10th data point with year labels shown for 

the REF and fODS simulations. The multi-model mean is derived from three CCMs in panel (a) (CCSRNIES, 

CMAM, and MRI) and four CCMs in panel (b) and (c) (CCSRNIES, CMAM, MRI, and WACCM). See Table 3-1 

for model descriptions. From Eyring et al. (2010a), their Figure 4.
3.27

Future Ozone and UV

a)

b)

c)
3.28

Chapter 3

trend from 1960 to 2100. The close agreement between 

the REF and gray traces (the sum of the ozone changes 

due to only the effect of ODSs (fGHG) and due to only the 

effects of GHGs (fODS)) in Figure 3-8(a) indicates that 

the system is close to being linearly additive (i.e., the two 

effects can be considered separately). The system deviates 

most from linear additivity around the turn of the century 

when ODS abundances and ozone depletion are close to 

their maximum. This may result from the fact that the 

fODS simulations are forced by SSTs taken from coupled 

climate model simulations where the radiative forcing effects of the varying ODS levels were included. As a result, 

even though ODS levels are kept fixed at 1960 values in 

the fODS simulations, the radiative effects of the varying 

ODSs could be partly felt through the SSTs.

Eyring et al. (2010a) found that in the tropical lower 

stratosphere (Figure 3-8(b)), ozone shows little response 

to ESC through the 20th and 21st centuries, as seen from 

the fGHG trace. The ~25% decrease in ozone from 1960 

to 2100 in the reference simulation results from GHGinduced changes to stratospheric dynamics as discussed 

above and confirmed by the fixed ODS simulation. Note 

also that in the fixed ODS simulation, ESC decreases with 

time in response to these circulation changes. As in the 

upper stratosphere, the response of ozone to ODSs and 

GHGs is again almost linearly additive as shown by the 

close agreement between the REF and gray traces in Figure 3-8(b). The ozone decreases simulated in the tropical 

lower stratosphere dominate the increases in the tropical 

upper stratosphere such that tropical total column ozone 

(Figure 3-8(c)) remains suppressed below what would be 

expected from changes in ESC from 2000 onwards. Interestingly, unlike ozone at 5 and 50 hPa in the tropics, total 

column ozone shows deviations away from linear additivity demonstrated by the lack of coincidence of the orange 

and yellow traces in Figure 3-8(c).

3.3.4 Midlatitude Ozone

3.3.4.1 Long-Term Projections of Midlatitude

Ozone

Figure 3-6 panels (c) to (f) show projections of 21st

century midlatitude (35°N–60°N and 60°S–35°S) column 

ozone and Cly at 50 hPa (Austin et al., 2010a; Eyring et al., 

2010a; Chapter 9 of SPARC CCMVal, 2010). The 50 hPa 

level was chosen as a representative level of future ozone 

changes in the lower stratosphere in the tropics, midlatitudes, and polar regions (see for example Figure 3-2(b) 

and Figure 3-9(b)). In the multi-model trend, in both hemispheres, minimum total column ozone is reached around 

the turn of the century, followed by a steady and significant increase. By 2025, northern (southern) midlatitude 

total column ozone is projected to have regained 130% 

(70%) of the amount lost between 1980 and 2000 (2002) 

and 230% (145%) of this loss regained by 2050 (see Table 

3-3). In other words, by 2050, midlatitude total column 

ozone in both hemispheres is projected to lie above 1980 

levels. By 2100, the column ozone in the northern (southern) midlatitudes is projected to have increased by 22 DU 

(19 DU) compared to 1980 amounts (Table 3-3).

As noted by Austin et al. (2010a), the midlatitude 

column ozone broadly follows the behavior of Cly at 50 

hPa (compare left and right panels of Figure 3-6) but with 

the ozone returning, on average, to 1980 levels 28+

–

5

4 and 

15 ± 5 years in advance of the chlorine for the northern 

and southern midlatitudes, respectively. For chlorine, 

there is, on average, no significant interhemispheric differences in the timing of the return to 1980 levels. Austin 

et al. (2010a) concluded that the earlier return of ozone 

to 1980 amounts in the Northern Hemisphere was mainly 

due to stronger transport from low latitudes, which was 

also noted in the independent studies of Shepherd (2008) 

and Li et al. (2009).

3.3.4.2 Processes Determining Future

Midlatitude Ozone

Chapter 9 of SPARC CCMVal (2010) found in all 

simulations analyzed that the increase in the midlatitude 

total column ozone resulted from ozone increases throughout the stratosphere. The increase in the upper stratosphere 

peaks in volume mixing ratio at ~3 hPa in both hemispheres 

(see Figures 3-9(a) and (b)). This in turn was the result of 

the combined effects of (i) a decline in Cly (Figures 3-9(c) 

and (d); see also Figures 3-6(d) and (f)) and Bry and, (ii) 

the GHG-induced cooling of the middle and upper stratosphere (Figures 3-9(e) and (f)). The relative importance of 

the factors affecting ozone can be different under a different GHG scenario (see Section 3.3.6). In a separate single 

model study, Waugh et al. (2009) also found that GHGinduced cooling of the middle and upper stratosphere was 

an important process for midlatitude ozone evolution in 

the 21st century under the SRES A1B scenario. For the 

northern midlatitudes, Li et al. (2009) estimated that by 

the 2060s stratospheric cooling, together with increased 

poleward transport by the BDC, increased the extratropical column ozone in the model analyzed by Waugh et al. 

(2009) by up to 6% compared to 1980 amounts. A smaller 

increase of 3% in the southern midlatitudes resulted from 

the smaller increase in transport in that hemisphere, consistent with the findings of Austin et al. (2010a). In both 

hemispheres the largest increase occurred at 60°. Chapter 9 of SPARC CCMVal (2010) found that the projected 

evolution of midlatitude middle and upper stratospheric 

ozone was very similar to that in the tropics in terms of 

the magnitude of the changes (compare Figures 3-2(b) and 
3.29

Future Ozone and UV

(a) (b)

(c) (d)

(e) (f)

O3 (ppm) O3 (ppm)

O3 (ppm)

O3 O (ppm) 3 (ppm)

O3 (ppm)

Southern Midlatitudes Northern Midlatitudes

Figure 3-9. Vertical profiles of differences in midlatitude (30°–50°) ozone (parts per million) over the 21st century based on a Multiple Linear Regression (MLR). Leftmost panels are for southern midlatitudes and rightmost 

panels for northern midlatitudes. (a) and (b) net change in the ozone profile from 2000 to 2100. (c) and (d) 

ozone changes congruent with changes in ESC. (e) and (f) ozone changes congruent with changes in temperature. See Table 3-1 for model descriptions. From Chapter 9 of SPARC CCMVal (2010), their Figure 9.10 

(a)-(f).
3.30

Chapter 3

3-9). In the lower stratosphere the evolution of midlatitude ozone differed from that in the tropics. However, in 

midlatitudes the increase in the residual meridional circulation leads to an increase in ozone rather than a decrease 

(Shepherd, 2008; Li et al., 2009).

An attribution of total column ozone changes in the 

northern and southern midlatitudes to changes in ODSs and 

GHGs from five CCMs analyzed by Eyring et al. (2010a) 

is shown in the correlative time series plot in panels (b) and 

(c) of Figure 3-10. The reference simulations show that 

total column ozone decreases from 1960 to 2000, but at a 

greater rate over southern midlatitudes than over northern 

midlatitudes. Over northern midlatitudes, ozone shows a –7 

DU/ppb sensitivity to ESC over the 1960 to 2000 period, in 

general agreement with the results reported in Guillas et al. 

(2004). Over southern midlatitudes the sensitivity is −16 

DU/ppb. The heightened sensitivity of total column ozone 

to ESC in the Southern Hemisphere most likely does not 

result from differences in the in situ contribution of ESC to 

southern midlatitude ozone destruction, but rather from the 

effects of export of ozone-depleted air from the Antarctic 

ozone hole. Because the multi-model trends shown in Figure 3-10 include fewer models than the multi-model trends 

displayed elsewhere, the quantitative values listed above 

are less certain than would be the case otherwise. A comparison of the reference and fixed GHG simulations for the 

midlatitudes shows that increasing GHGs have elevated 

ozone throughout the period. In both hemispheres, as ESC 

decreases, total column ozone does not simply retrace the 

1960–2000 path, but shows systematically elevated ozone 

through the 21st century. As a result, over northern midlatitudes, total column ozone returns to 1980 values in 

the mid-2020s, well before ESC returns to its 1980 value 

around 2050. Similarly, over southern midlatitudes total 

column ozone returns to 1980 values in the mid-2030s (a 

decade later than in the northern midlatitudes), and well 

before ESC returns to its 1980 value around 2050. From 

the fixed ODS simulation (blue traces in Figure 3-10), it is 

clear that the elevated ozone through the 21st century results from GHG-induced stratospheric cooling and changes in transport, in particular changes in the strength of the 

BDC (see Figure 3-3). These simulations also show ESC 

decreasing with time even though ODSs are fixed at 1960 

values. This likely results from the increasing strength of 

the BDC through the 21st century (Section 3.2.4.1) and a 

resultant decrease in the time available to photolyze ODSs. 

It is also clear from the reference and fixed ODS simulation traces in Figure 3-10 that by 2100, total column ozone 

over midlatitudes is still influenced by ESC. In both the 

northern and southern midlatitudes the effects of ODSs and 

GHGs on ozone are approximately additive (agreement of 

black and yellow traces in Figures 3-10(b) and (c)) in the 

multi-model trend.

3.3.5 Polar Ozone

3.3.5.1 Long-Term Projections of Polar Ozone

The largest ozone depletion seen in CCM simulations 

based on the SRES A1B GHG and A1 adjusted halogen 

scenarios occurs in the polar lower stratosphere, especially 

over Antarctica where large modeled ozone losses give rise 

to a springtime Antarctic ozone hole, consistent with observations. As a result, a major focus is the projected evolution 

of polar lower stratospheric ozone during spring.

Figure 3-11 shows the evolution of the individual 

model and multi-model trend estimates of total column 

ozone (left panels) and lower stratospheric inorganic chlorine (Cly) (right panels), for March in the Arctic (60°N–

90°N) and October in the Antarctic (60°S–90°S) from the 

17 CCMs analyzed by Eyring et al. (2010a). In both polar 

regions, the long-term evolution of total column ozone is 

qualitatively the same as in other regions. There is a broad 

minimum around 2000 followed by a slow increase until the end of the 21st century. There are, however, as in 

the extrapolar regions, significant quantitative differences 

among the models, including a wide spread in minimum 

values over Antarctica around 2000. Austin et al. (2010a) 

and Chapter 9 of SPARC CCMVal (2010) indicate that 

the spread between the individual model simulations of 

Antarctic total column ozone increased from CCMVal-1 

to CCMVal-2, while no change was seen for the Arctic. 

However, when adjusting the raw data to a common 1980 

baseline by removing the individual offset values with the 

TSAM method (Section 3.3.2.2), the model spread is considerably reduced.

In the Antarctic, the strongest ozone depletion in 

the October mean multi-model trend is simulated in 2003 

(see Table 3-3). The minimum in the October multimodel trend total column ozone is about 80 DU lower 

than the 1980 value. For the Arctic, the largest total column ozone depletion simulated in March 2002 is ~30% of 

the Antarctic depletion simulated in October 2003 (Figure 

3-11, left panels). Arctic spring ozone returns earlier to 

historical values than Antarctic spring ozone. Arctic total 

column ozone is projected to regain 95% (75–115%)1 of 

the amount lost between 1980 and 2002 by 2025. Arctic 

column ozone is simulated to increase to 46 DU above 

1980 values by the end of the 21st century, which is equivalent to an increase in column ozone of twice the amount 

lost between 1980 and 2002 (see Table 3-3). In Antarctica, 

45% (35–55%) of the ozone lost since 1980 is projected to 

be replenished by 2025, and 100% (85–110%) by 2050. 

By the end of the 21st century, Antarctic October ozone 

1

 All quoted ranges are 95% confidence intervals.
3.31

Future Ozone and UV

Figure 3-10. Correlative time series plots for total column ozone (Dobson units) averaged over different latitude 

ranges (Appendix 3A). (a) March means for the Arctic, (b) annual means for northern midlatitudes, (c) annual 

means for southern midlatitudes, and (d) October means for the Antarctic. ESC values are defined as ESC = 

Cly + 60×Bry at 50 hPa. The black trace shows the reference simulation with time-varying GHGs and ODSs with 

reference years shown every 10th data point. The blue trace shows results from a simulation where prescribed 

ODSs are fixed at 1960 values. The green trace shows results from a simulation where prescribed GHGs are 

fixed at 1960 values. The yellow traces show the additive effects of the fixed ODS and fixed GHG simulations 

(yellow = blue + green) as in Figure 3-8. Differences between the black and yellow traces indicate a lack of 

linear additivity in the system. The multi-model trend estimate is derived from five CCMs (CCSRNIES, CMAM, 

MRI, ULAQ, and WACCM; see Table 3-1). From Eyring et al. (2010a), their Figure 7.
3.32

Chapter 3

is projected to be 43 DU above 1980 levels (compare to 

78 DU decrease from 1980 to 2003). However, by 2100, 

October mean ozone is still projected to be 9 DU lower 

than in 1960 (95% of the 1960–2003 loss replenished) 

consistent with slightly enhanced ESC in 2100 compared 

to 1960.

Several different indices have been defined and 

applied in previous assessments to quantify variations in 

Antarctic ozone, in terms of the area of the ozone hole, the 

polar cap average ozone, the ozone mass deficit (Bodeker 

et al., 2005), or the daily minimum total column ozone in 

spring (WMO, 2007). Figure 3-12 (upper panel) shows 

the simulated and observed ozone hole areas, based on the 

size of the area with column ozone less than 220 DU, from 

17 CCMVal-2 simulations (Austin et al., 2010b; Chapter 9 

of SPARC CCMVal, 2010). While a few models capture 

the observed size of the Antarctic ozone hole reasonably 

well, the models on average underestimate the observed 

ozone hole area by about 20% (Austin et al., 2010b). This 

value is similar to that found for the CCMVal-1 simulations (Eyring et al., 2006). It reveals that, although some 

of the CCMs have been improved since CCMVal-1, others have become worse, and no fundamental improvement 

was achieved for the majority of the models. The projected date when the Antarctic ozone hole will disappear 

varies from the 2020s in models with small ozone holes to 

(DU) (DU)

(ppb) (ppb) NIWA

Figure 3-11. As in Figure 3-6, but for the latitude range 60°N–90°N in March (upper row) and the latitude 

range 60°S–90°S in October (lower row). The red vertical dashed line indicates the year when multi-model 

trend in total column ozone (DU; left) and Cly at 50 hPa (ppb; right) returns to 1980 values and the blue vertical 

dashed lines indicate the uncertainty in these return dates. Note the different vertical scale among the panels. 

Redrawn from Figures 9.11, 9.12, 9.13, and 9.14 of SPARC CCMVal (2010) and updated with two new CCM 

simulations.
3.33

Future Ozone and UV

the end of the simulation period in 2100 for those models 

that agree best with the observed ozone hole areas in the 

past. Even with some outliers removed, Antarctic spring 

ozone depletion below 220 DU is projected to still occur in a broad range between 2060 to 2100, or later. As 

discussed in Austin et al. (2010b), the representation of 

the ozone hole area in individual models depends on the 

definition used. For example, the use of an ozone isoline 

to define the ozone hole area (e.g., the 220 DU contour) 

in models that have a high or low ozone bias results in 

ozone hole areas that are systematically low or high, respectively, even if the models have realistic interannual 

and long-term variations. Calculating total column ozone 

changes relative to the 1960–1965 ozone mean leads to 

an improvement of the simulated ozone hole area in some 

models. These models have a general high ozone bias. 

Other models perform better when the steepest ozone gradient is used to define the edge of the ozone hole. This 

diagnostic identified some models where the simulated 

Antarctic polar vortices are too small in area compared 

to observations, which would also limit the size of their 

ozone holes (Bodeker et al., 2002). In general, most consistent results among the models are achieved when the 

steepest gradients are used. However, independent of the 

applied index for the ozone hole area, most CCMs have 

ozone holes that are significantly smaller than observed, 

by up to 30%. With the large spread between the individual CCM simulations of the Antarctic ozone hole and its 

likely dependence on the applied definition for the ozone 

hole, quantitative projections of the disappearance of the 

ozone hole remain uncertain.

The minimum spring ozone, calculated from Antarctic daily mean ozone from September to November 

(Figure 3-12, middle panel), shows a wide range of ozone 

values, from one model that does not fall below the 220 

DU ozone hole threshold value to models with lower than 

observed ozone minima. Most models simulate the lowest 

ozone minima for the most recent years and suggest an 

increase of minimum ozone starting around 2010, which 

agrees well with projections in WMO (2007). Only models with a high ozone bias exceed the 220 DU limit by the 

end of the 21st century, while models that represent well 

or underestimate past ozone compared to observations, do 

not predict a return to minimum ozone values above 220 

DU before 2100.

The ozone mass deficit, defined as the mass of 

ozone required to elevate column ozone everywhere over 

Antarctica to 220 DU, averaged over the months of September and October (Bodeker et al., 2005) (Figure 3-12, 

lower panel) is a sensitive diagnostic as it reflects how 

well the CCMs simulate the ozone hole area and the minimum ozone. Austin et al. (2010b) show that models that 

capture the observed size of the ozone hole area typically simulate lower than observed Antarctic ozone spring 

NIWA-SOCOL

40

30

20

10

0

1960 1980 2000 2020 2040 2060 2080 2100

Year

0

1960 1980 2000 2020 2040 2060 2080 2100

Year

0

1960 1980 2000 2020 2040 2060 2080 2100

Year

350

300

250

200

100

50

150

Minimum Ozone Column (DU) Maximum Ozone Hole Area (106 km2) Ozone Mass Deficit (Mt) 60

50

40

30

10

20

Observations

Figure 3-12. Simulated and observed ozone hole 

areas, based on a fixed 220 DU amount (upper 

panel), minimum Antarctic spring total ozone column 

(middle panel), and ozone mass deficit (lower panel) for the period 1960–2100, calculated from the 

CCMVal-2 REF-B2 simulations. The curves indicate 11-year running means of the data for individual 

years. See Table 3-1 for model descriptions. From 

Austin et al. (2010b), their Figures 4(a), 6, and 9.
3.34

Chapter 3

minima, in contrast to models with smaller ozone holes 

that typically show higher than observed Antarctic ozone 

spring minima. Both errors contribute to a large spread 

in the ozone mass deficit of the models for the past and 

future. The models suggest that the ozone mass deficit 

should not increase any longer after about 2010; however, 

due to the sensitivity of this diagnostic, future projection 

of the ozone mass deficit is highly unreliable.

3.3.5.2 Processes Determining Future Polar

Ozone

In both hemispheres, multi-model total column 

ozone in polar spring follows the evolution of Cly (Figure 

3-11). Ozone starts to increase in the first decade of the 

21st century at about the same time as the Cly abundances 

in the polar lower stratosphere have reached their maximum and begin to decline. Compared to CCMVal-1, the 

simulation of Cly, on average, has improved in the 

CCMVal-2 models due to improvements in individual 

CCMs (Austin et al., 2010a; Austin et al., 2010b; Chapter 

9 of SPARC CCMVal, 2010); however, differences in Cly

of up to 1 ppb in the polar lower stratosphere in October 

remain. These Cly differences, together with differences 

in the dynamical processes determining the strength of the 

stratospheric polar vortices and temperatures, cause the 

spread in simulated polar ozone.

The Cly concentrations in Figure 3-11 return to historical levels almost simultaneously in both hemispheres, 

while the ozone return dates show interhemispheric differences of more than two decades. As expected from the 

discussion in Section 3.2, this result implies that processes 

other than halogen chemistry affect the future evolution of 

polar ozone and that the relative impact of these processes 

varies with location. The relative contributions of ESC 

or increasing GHGs to the projected ozone change can 

be estimated by comparing the reference simulation with 

sensitivity simulations that use prescribed fixed historical abundances of GHGs or ODSs (Eyring et al., 2008). 

Waugh et al. (2009) compare a “climate change” simulation (i.e., using fixed ODS abundances for 1960) of the 

GEOSCCM with a reference simulation that accounts for 

GHG and ODS changes from 1960 to 2100. They find 

that in much of the stratosphere, the annual mean GHGinduced ozone increase from 1960–2100 is comparable to 

the ODS-induced ozone decrease for the period with highest ODS concentration (1995–2005). However, ozone 

in the Antarctic lower stratosphere is clearly dominated 

by ODS-induced destruction, and increases in GHG concentrations do not have a significant impact on Antarctic 

polar temperatures or ozone in their CCM. It should be 

noted that most studies of Antarctic ozone recovery focus 

on changes in October means. The effects of an extended 

polar vortex persistence, due to future GHG-induced cooling and an associated seasonal delay of ozone recovery 

to November, on UV exposure has not been thoroughly 

assessed in multi-model studies.

Eyring et al. (2010a) found similar results in a 

multi-model framework using additional CCMVal-2 

sensitivity simulations. As illustrated in Figure 3-10(d), 

reference simulations in the Antarctic show multi-model 

total column ozone decreasing from 1960 to 2000 with an 

average 53 DU/ppb sensitivity to ESC, leading to a 38% 

decrease in total column ozone over this period. Unlike 

the midlatitudes and Arctic, the total column ozone evolution over Antarctica in October shows almost no sensitivity to changes in GHGs, with the return path (21st century) closely tracking the outbound path (20th century). 

This is corroborated by the fixed ODS simulation, which 

shows almost no change in Antarctic ozone in October 

in response to increasing GHGs. If anything, increasing 

GHGs have slightly elevated Antarctic total column ozone 

in the model simulations. This updates earlier studies 

that predicted that increasing radiative cooling generated 

by increasing levels of GHGs or by GHG-induced stratospheric water vapor enhancements would worsen polar 

ozone depletion by increasing the likelihood for PSC formation, and might even create an ozone hole in the Arctic 

(Austin et al., 1992; Shindell et al., 1998; Kirk-Davidoff et 

al., 1999; Waibel et al., 1999). More recent research suggests that in the Arctic, GHG-induced changes in dynamics are expected to dominate the effects of GHG-induced 

radiative cooling on the formation of PSCs (Eyring et al., 

2010a). In comparison, in the Antarctic, where the wintertime polar temperatures typically fall below the threshold 

for PSC formation, stratospheric cooling due to increasing GHGs does not strongly enhance October total column 

ozone depletion.

The multi-model trend of the reference simulations 

in the Arctic (Figure 3-10(a)) shows total column ozone 

decreasing from 1960 to 2000 with a −14 DU/ppb sensitivity to ESC, somewhat smaller than the −20 DU/ppb reported in observations (Dhomse et al., 2006). The increase 

in ESC from 1.1 ppb in 1960 to 3.5 ppb in 2000 leads to 

a 7.4% decrease in total column ozone over this period. 

In contrast to the Antarctic, total column ozone over the 

Arctic is elevated above what would be expected from 

changes in ESC. This most likely does not result from the 

effects of GHG-induced upper stratospheric cooling as in 

the tropics (Figure 3-8(a)) since, as discussed in Butchart 

et al. (2010), in the models the extra radiative cooling from 

increasing GHGs is approximately balanced by a concomitant increase in the adiabatic warming through increased 

polar downwelling. The net effect is a near-zero temperature trend in the Arctic winter lower stratosphere (Butchart 

et al., 2010; Figure 4.4 of SPARC CCMVal, 2010). The 

more likely cause is the strengthening of the BDC (Section 3.2.4.1), which more effectively advects ozone into 
3.35

Future Ozone and UV

the Arctic than the Antarctic stratosphere (Austin and 

Wilson, 2006; Shepherd, 2008). The effects of ODSs and 

GHGs on Arctic total column ozone appear to be largely 

independent and therefore add linearly as evidenced by the 

close agreement of the black and yellow traces in Figure 

3-10(a).

Li et al. (2009) analyzed in more detail the effects 

of climate change on annual mean ozone by comparing 

the post-CFC era (2060–2069) with the period 1975–1984 

in two simulations with the GEOS CCM. They found a 

uniform increase in ozone mixing ratio in the upper stratosphere (Figure 3-13(a)) that they ascribe to a slowing of 

photochemical ozone loss rates due to GHG-induced cooling (Figure 3-13(d)). However, the simulated changes in 

ozone column and their latitudinal structure are controlled 

by changes in the lower stratosphere, where significant 

increases of column ozone are simulated at middle and 

high latitudes, specifically in the Northern Hemisphere 

(Figures 3-13(b) and (c)). They are associated with an acceleration of the Brewer-Dobson circulation in the model 

with increased tropical upwelling and extratropical downwelling (Figure 3-13(e)), leading to enhanced advective 

transport of ozone (Figure 3-13(f)). Although, in the annual mean, these dynamical effects are stronger at midlatitudes than at polar latitudes, the Li et al. (2009) study demonstrates that changes in ozone abundance and the mean 

advective transport have a qualitatively similar pattern in 

the lower stratosphere, and emphasizes the important role 

of changes in the mean advection for lower stratospheric 

ozone changes.

Figure 3-13. Annual mean, decadal differences between 2060–2069 and 1975–1984 from a simulation of the 

GEOS CCM with ODSs fixed at 1960 values. (a) Ozone concentrations in ppm, (b) ozone concentrations in 

DU/km, (c) total column ozone (DU) and contributions from the upper (above 15 hPa) and lower (below 15 hPa) 

stratosphere, (d) temperature (K) (e) Vertical residual velocity w

−* in mm/s, and (f) mean ozone advection (10–6 

DU/km/s). From Li et al. (2009), their Figure 2.

-0.1

0.1

0.3

0.3

0.5

0.7

0.9

(a)

-90 -60 -30 0 30 60 90

Latitude

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100

500 Pressure (hPa)

Ozone (ppmv)

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1.2

(b)

-90 -60 -30 0 30 60 90

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 10

100

500 Pressure (hPa)

10

20

30

40

Height (km)

Latitude

Ozone (DU/km)

Total

Upper Strat

Lower Strat

(c)

-90 -60 -30 0 30 60 90

Latitude

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 5

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Ozone (DU)

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-8

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(d)

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100

500 Pressure (hPa)

T (K)

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0.30.

0.

5

5

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500 Pressure (hPa)

W* (mm/s)

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(f)

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500 Pressure (hPa)

O3 Advection (10-6 DU/km/s)
3.36

Chapter 3

3.3.6 Ozone Return Dates and Ozone 

Recovery

In this section, ozone return dates to levels typical 

of 1960 and 1980, derived from the CCM projections, are 

presented. These are complemented by quantitative statements about the expected date for full recovery of ozone 

from the effects of ODSs based on CCM simulations as 

described in Box 3-2. Specifically, a set of reference simulations and a set of fixed ODS simulations (as described 

in Section 3.3.1) are analyzed for the future date when the 

two sets of ozone projections are no longer statistically 

distinguishable within internal and inter-model variability 

(Eyring et al., 2010a). A Student t-test is applied to quantify the likelihood that, at some specified date, the ozone 

values taken from the two sets of simulations come from 

the same statistical population. The outcome of such tests 

is discussed using the terminology of the IPCC (see Box 

TS.1 of Solomon et al., 2007). For example a t-test result 

of >95% suggests that it is “extremely likely” that, within 

model variability, full ozone recovery from the effects of 

ODSs is projected to have occurred, if >90% it is “very 

likely,” while if >66% it is “likely” that it is projected to 

have occurred. For values between 33% and 66% probability, it is “about as likely as not” and for values <33% 

it is “unlikely” that full recovery of ozone is projected to 

have occurred.

The response to the question of when stratospheric 

ozone will return to undisturbed levels depends on the selected target year. Figure 3-14 and Table 3-4 summarize 

1960 and 1980 return dates of total column ozone in different latitude zones derived from the multi-model trend. 

Return dates generally occur later when referenced to 

1960 and later at higher latitudes (Chapter 9 of SPARC 

CCMVal, 2010). However the increase is not symmetric 

about the equator, i.e., return dates are later in the Antarctic than in the Arctic.

CCM projections suggest that it is as likely as not 

that tropical total column ozone will return to 1980 values, since there is no consensus among the CCMs on a 

return to 1980 values by the end of the century (Austin 

et al., 2010a; Eyring et al., 2010a; Chapter 9 of SPARC 

CCMVal, 2010). Correspondingly, the uncertainty on 

tropical column ozone return dates derived from the 95% 

TSAM confidence interval extends from 2030 to beyond 

the end of the century in Figure 3-14. However, if instead of a return to 1980 values, a return to 1960 values is 

considered, models consistently predict that tropical total 

column ozone remains below values typical of 1960 due 

to the increase in tropical upwelling. In contrast, Cly and 

ESC (Table 3-5) in the tropics return to 1980 values faster 

than in all other regions with only minor differences between them. Full recovery of tropical column ozone from 

the effects of ODSs is not reached at the 95% confidence 

level by the end of the 21st century, while it is likely at 

the 65% level that total column ozone has fully recovered 

by ~2070 in the tropics (see Figure 3-15). In the tropical 

upper stratosphere, although ozone is influenced by ODSs 

throughout the 21st century, ozone returns to 1980 values 

Figure 3-14. Date of return to 1980 total 

column ozone (black triangle and error 

bar), Cly at 50 hPa (red triangle and error 

bar), and ESC at 50 hPa (blue triangle 

and error bar) for the annual average 

(global, tropical, and midlatitude) and 

spring (polar) total ozone column derived 

from the multi-model trend of the 

CCMVal-2 reference simulations (17 

CCMs) in each latitude band. The error 

bars on the multi-model trend estimate of 

return date are derived from the 95% 

confidence intervals. ESC is calculated 

as Cly + 60×Bry except for E39CA where 

Cly instead of ESC was used. While a 

few models project a return of tropical 

total column ozone to 1980 levels, most 

do not with the result that the 95% confidence interval extends from 2030 to 

beyond the end of the century which explains the large error bar in the tropical column ozone return dates. 

Redrawn from Figures 9.20 and 9.21 of SPARC CCMVal (2010) and updated with two new CCM simulations 

(from Eyring et al. (2010a), their Figure 10b).
3.37

Future Ozone and UV

in the late 2020s (Figure 3-8(a)), which is earlier than in 

other regions. Although ozone decreases continuously 

from 1960 to 2100 in the tropical lower stratosphere, together with the Arctic lower stratosphere, they are the only 

regions where ozone has very likely (i.e., more than 90% 

confidence level) fully recovered from the effects of ODSs 

by the end of the 21st century (Eyring et al., 2010a). However, in percentage terms, changes to the tropical ozone 

column over the 21st century are much smaller than those 

in other regions.

Of the five regions considered (see Table 3-4), 

total column ozone returns to 1980 values earliest over 

northern midlatitudes, around 2021 (2017–2026). While 

the total column ozone evolution is qualitatively similar 

Table 3-4. Date of return to 1960 and 1980 total column ozone calculated from the multi-model trend 

estimate of the 17 CCMs’ reference simulations.

Region Date of

Return

Year When MultiModel Mean Total

Column Ozone

Returns

Year When Lower

Bound of Error

Bar Returns

Year When Upper

Bound of Error

Bar Returns

Global annual mean

1960 2053 2046 2064

1980 2032 2027 2038

Tropics annual mean

1960 ---- ---- ----

1980 2042 2028 ----

Northern midlatitude 

annual mean

1960 2029 2024 2036

1980 2021 2017 2026

Southern midlatitude 

annual mean

1960 2055 2049 2064

1980 2035 2030 2040

Antarctic October mean

1960 ---- 2100 ---

1980 2051 2046 2057

Arctic March mean

1960 2041 2036 2048

1980 2026 2023 2031

Table 3-5. Date of return to 1960 and 1980 ESC at 50 hPa calculated from the multi-model trend estimate 

of the 17 CCMs’ reference simulations.

Region Date of

Return

Year When MultiModel Mean ESC

Returns

Year When Lower

Bound of Error Bar

Returns

Year When Upper

Bound of Error Bar

Returns

Global annual mean

1960 --- 2092 ---

1980 2050 2045 2057

Tropics annual mean

1960 2087 2073 ---

1980 2044 2036 2051

Northern midlatitude 

annual mean

1960 --- --- ---

1980 2052 2047 2060

Southern midlatitude

annual mean

1960 --- --- ---

1980 2053 2048 2060

Antarctic October mean

1960 --- --- ---

1980 2060 2053 2069

Arctic March mean

1960 --- --- ---

1980 2060 2053 2069
3.38

Chapter 3

1960 Baseline-Adjusted Total Column Ozone (fODS)

(DU) (DU) (DU)

(DU) (DU) (DU)
3.39

Future Ozone and UV

over the midlatitudes of both hemispheres in the CCMs, 

southern midlatitude total column ozone returns to 1980 

values later than over northern midlatitudes, i.e., around 

2035 (2030–2040) (see Figure 3-14). The difference in 

the date of return to 1980 values appears to be due to interhemispheric differences in changes in transport. The 

increase in stratospheric circulation transports more ozone 

into the northern midlatitude lower stratosphere than into 

the southern midlatitudes (Shepherd, 2008). In addition, 

over southern midlatitudes, ozone is also influenced by 

ozone loss in the Antarctic, where the return to 1980 levels occurs much later. In all CCMs the return of ozone to 

1980 values in the midlatitudes occurs 10 to 30 years earlier than that of Cly and ESC (around 2050 in both hemispheres). Nonetheless, by 2100 total column ozone over 

midlatitudes is still influenced by ODSs. Full recovery 

of total column ozone from ODSs has likely occurred in 

northern midlatitudes but not likely to have occurred over 

southern midlatitudes (Figure 3-15).

The Antarctic spring total column ozone evolution 

is dominated by the evolution of ODSs. Therefore, ozone, 

Cly, and ESC return dates are very similar in most models 

(Figure 3-14). The latest return of total column ozone to 

1980 values is projected to occur over Antarctica which, 

for October means, occurs around 2051 (2046–2057), 

while 1960 values are not reached again before 2100 since 

Cly remains elevated above 1960 levels in 2100 (Figure 

3-11(d)). In the Arctic, the sensitivity of the return date to 

the chosen baseline year (1960 or 1980) is small, with a return to 1980 ozone values around 2026 (2023–2031), and 

to 1960 values around 2041 (2036–2048). In contrast to 

the Antarctic, Arctic spring ozone returns earlier to 1980 

values than Arctic chlorine, indicating that other effects 

such as changes in transport are more important for future 

ozone in the Arctic as is the case for northern midlatitudes 

(Section 3.3.4.2). A return of stratospheric chlorine to the 

low values of 1960 does not happen in the polar regions 

before the end of the 21st century. It is unlikely that full 

ozone recovery from ODSs will be reached by the end of 

the 21st century over Antarctica, while it is likely that it 

will occur in the Arctic by ~2035 (Waugh et al., 2009; 

Eyring et al., 2010a) (see also Figure 3-15).

Global total column ozone is projected to return 

to its 1980 value around 2032 (2027–2038), which is 15 

years earlier than when global Cly at 50 hPa returns to its 

1980 value and 18 years earlier than when global ESC returns to its 1980 value. CCM projections suggest that this 

early return is primarily a result of GHG-induced cooling 

of the upper stratosphere, because the effects of circulation 

changes on tropical and extratropical ozone largely cancel 

in the global mean. Figure 3-16 summarizes return dates 

as a function of latitude for annual mean total column 

ozone to a value appropriate to reference years between 

1965 and 2000 (Austin et al., 2010a; Chapter 9 of SPARC 

CCMVal, 2010). The figure shows results for the partial 

ozone column above 20 hPa (upper panel), the partial column between 500 and 20 hPa (middle panel) and for the 

combined column above 500 hPa. Tropospheric ozone below 500 hPa is excluded. Above 20 hPa, where changes in 

halogens and temperature dominate the ozone evolution, 

the return of ozone to historical levels is projected to occur 

steadily. In the lower stratosphere (Figure 3-16, middle 

panel) a return date could not be determined for the tropics 

due to the strengthening BDC which systematically decreases tropical ozone as the simulations proceed (Waugh 

et al., 2009). The results also show a strong hemispheric 

asymmetry, with Antarctic ozone returning to historical 

levels much later than Arctic ozone. Again, this is largely 

due to the increased BDC, which for the models, on average, has much more influence in the Northern Hemisphere 

Figure 3-15 (at left). 1960 baseline-adjusted total column ozone projections in the reference simulations 

(thick black solid line) compared to the fixed halogen simulation (fODS, thick black dashed line): (a) global 

(90°S–90°N annual mean), (b) tropics (25°S–25°N annual mean), (c) northern midlatitudes (35°N–60°N annual mean), (d) southern midlatitudes (35°S–60°S annual mean), (e) Arctic (60°N–90°N March mean), and (f) 

Antarctic (60°S–90°S October mean). The colored lines show individual models for the subset that submitted 

fODS (see Table 3-1 for model descriptions). The horizontal green line shows the level of the multi-model trend 

estimate in 1980. The red vertical dashed line indicates the year when the multi-model trend estimate in the 

reference simulation returns to 1980 values and the blue vertical dashed lines indicate the uncertainty in these 

return dates. The black solid circles with vertical bars show the observations and their 2σ errors processed as 

in Figures 3-6 and 3-11. Nine CCMs are included in the multi-model mean (CCSRNIES, CMAM, GEOSCCM, 

LMDZrepro, MRI, SOCOL, UMSLIMCAT, ULAQ, and WACCM). In all regions except the tropics, the milestone 

of full ozone recovery occurs significantly later than when ozone returns to its 1980 values. The milestone of 

full ozone recovery is derived from the period when the fixed halogen simulation is not statistically distinguishable from the reference simulation (black dotted line in the bottom of each panel plotted against the percentage 

scale at the bottom right). For clarity, only the 95% confidence intervals are shown. Adapted from Eyring et al. 

(2010a), their Figure SM12.
3.40

Chapter 3

than in the Southern Hemisphere (Austin and Wilson, 

2006; Shepherd, 2008). The combination of the lower and 

upper stratospheric columns (Figure 3-16, lower panel) 

shows that in the tropics the total ozone column increases 

until about 2050 due to decreasing halogen amounts and 

stratospheric cooling, but thereafter ozone decreases due 

to the increasing BDC and does not return to pre-1985 

values before the end of the simulations.

The ozone return dates discussed above apply to a 

single GHG scenario, the SRES A1B scenario. However, 

this scenario represents only one plausible future and it is 

therefore important to also assess the ozone return dates 

under different GHG evolutions. Ideally, all models that 

performed the reference simulations would have also performed all sensitivity simulations. However, the various 

GHG scenarios were only performed by a small subset 

of models (see GHG-x simulations in Table 3-1). Differences in stratospheric column ozone among six GHG 

scenarios (SRES A1B, A2, and B1; RCP 2.6, 4.5, and 

8.5; see also Section 3.3.1) assessed from four CCMs by 

Eyring et al. (2010b) are found to be largest over northern 

midlatitudes and in the Arctic with divergence mainly in 

the second half of the 21st century. In the midlatitudes, the 

return of stratospheric column ozone to 1980 values varies by up to 10 years among the GHG scenarios, while in 

polar regions differences of 15–20 years are found. Overall, the differences of ozone return dates among the GHG 

scenarios simulated by the four individual models is of the 

same order as the uncertainty in ozone return dates derived 

from the multi-model mean of the 17 CCMs in Eyring et 

al. (2010a) in the midlatitudes, while it is smaller in the 

tropics and larger in the Arctic and Antarctic. The results 

suggest that effects of GHG emissions on future stratospheric ozone should be considered in climate change mitigation policy, and ozone projections should be assessed 

under more than a single GHG scenario. However, more 

CCMs will need to perform the GHG sensitivity simulations to arrive at more robust conclusions. Furthermore, to 

assess full recovery of ozone from ODSs under different 

GHG scenarios, simulations with fixed ODSs under each 

scenario would be required.

2035

2025

2015

2045

2035 2025

2015

2005

2065 2045

2015

2055

2035

2025

2025

2015

2055

2065

2005

2035

Figure 3-16. Date of return of the annual mean column ozone to the value appropriate to the reference 

year indicated on the abscissa. The mean model 

result was smoothed with an 11-year running mean 

filter. Contour interval is 5 years; red values indicate 

later dates and blue values indicate earlier dates. 

Data prior to 1965 (which limits the definition of the 

reference year data) or after 2094 (which limits the 

data for the return year) do not exist because of the 

need for an accurate time-smoothed field. The white 

region in the figure indicates where the mean model 

ozone has not recovered by the end of the simulations (nominally 2094). Results are shown for the 

total column above 500 hPa (bottom panel), for the 

range 500–20 hPa (middle panel), and for the column above 20 hPa (top panel). From Chapter 9 of 

SPARC CCMVal (2010), their Figure 9.25.
3.41

Future Ozone and UV

3.3.7 Uncertainties in Model Projections 

and Open Questions

3.3.7.1 Uncertainty in Future Emissions

Scenarios

Uncertainty in projections of stratospheric ozone 

can usefully be broken into four sources (Charlton-Perez 

et al. (2010) using an approach based on Hawkins and 

Sutton (2009)): (a) internal variability of the chemistryclimate system, (b) model uncertainty due to differences in 

the design and parameters of CCMs and missing or poorly 

represented processes in CCMs, (c) uncertainty in future 

emissions scenarios for GHGs, and (d) uncertainty in future emissions scenarios for ODSs, which is likely smaller 

than (a) to (c).

Figure 3-17 shows estimates of uncertainty in each 

of the three uncertainty components (a)–(c) for CCMVal-2 

projections of total column ozone in six geographical regions of the globe. Note that the GHG scenario uncertainty was estimated using only a small subset of models. 

In total, four runs of three CCMs forced with SRES GHG 

scenarios (solid green lines) and three runs of a single 

CCM forced with RCP GHG scenarios (dashed green 

lines) provide the estimates shown (see Charlton-Perez et 

al. (2010) for more details).

In either case, it is clear that, for most geographical 

regions, model uncertainty is the dominant contributor to 

the total uncertainty in projections of future ozone column 

amount at least up to the mean estimate of ozone return 

to 1980 values. In the tropical band, internal variability 

is comparable in size to scenario uncertainty, but in other 

regions internal variability is generally a small contributor to total uncertainty in future ozone column amount. 

Charlton-Perez et al. (2010) therefore suggest that continued investment and development of CCMs could lead 

to a refinement of ozone projections for the period up to 

and including ozone return to 1980 values, assuming that 

CCM developments lead to a reduction in model uncertainty.

There remain other uncertainties associated with 

ozone prediction that are difficult to quantify, particularly 

the mean biases present in the ozone climatology of many 

of the CCMs (see Chapter 9 of SPARC CCMVal, 2010) 

and missing or poorly represented processes in the entire 

ensemble of CCMVal models (for example the quasibiennial oscillation). The effects of solar variability were 

not considered in the CCM simulations assessed in this 

chapter. No CCM studies have yet assessed the effects of 

a long-term sustained change in solar output, such as the 

Maunder minimum, on stratospheric ozone. Other uncertainties, related to major changes in future human behavior 

Global 90 to -90

2000 2020 2040 2060 2080

Year

0

1

2

3

4

Standard Deviation (DU)

Annual 25 to -25

2000 2020 2040 2060 2080

Year

0.0

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Annual 60 to 35

2000 2020 2040 2060 2080

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Standard Deviation (DU)

Annual -35 to -60

2000 2020 2040 2060 2080

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Mar 90 to 60

2000 2020 2040 2060 2080

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10

15

Standard Deviation (DU)

Oct -60 to -90

2000 2020 2040 2060 2080

Year

0

5

10

15

20

25

Standard Deviation (DU)

Figure 3-17. Uncertainty in the predicted totalcolumn decadal-mean ozone amounts relative to 

1980 values estimated for different geographic regions. Uncertainty is expressed as the one standard 

deviation estimate for internal variability (orange), 

model uncertainty (blue), and the total uncertainty 

(black) all shown in solid lines. Solid green lines 

show an estimate of the one standard deviation estimate for scenario uncertainty calculated from integrations of CCSRNIES, GEOSCCM, and WACCM 

forced with alternative SRES greenhouse gas scenarios. Dashed green lines show an alternative estimate of scenario uncertainty calculated using integrations of the CAM3.5 model forced with alternative 

RCP greenhouse gas scenarios. Total uncertainty 

estimates are derived from the SRES scenario integrations only. From Charlton-Perez et al. (2010), 

their Figure 3.
3.42

Chapter 3

that are not considered in the scenarios used in this chapter, are discussed in Chapter 5 (Section 5.4) and include 

geoengineering of the climate system, increases in N2O 

emissions from automotive biofuels, and enhanced emissions from aviation and rockets.

3.3.7.2 Future CCM Development

Following this Assessment, CCMs are likely to 

undergo significant further development that could have 

impacts on their projections of future ozone and could either increase or decrease the spread between models. The 

executive summary of SPARC CCMVal (2010) recommends that “Development should continue toward comprehensive troposphere-stratosphere CCMs, which include 

an interactive ocean, tropospheric chemistry, a naturally 

occurring QBO, spectrally resolved solar irradiance, and a 

fully resolved stratosphere.” The impact of some of these 

processes on ozone projections was partially assessed in 

multiple CCMs by CCMVal.

Other than in CMAM, sea surface temperatures and 

sea ice concentrations were prescribed in all other CCM 

simulations used in this chapter. Important couplings 

between the atmosphere, the oceans, and the cryosphere 

are currently not represented in CCMs. Inclusion of these 

couplings in the CCMs will lead to a more complete representation of the climate system and climate feedbacks, 

which could be important for simulations of stratospheric 

ozone and its impact on tropospheric climate. Several 

recent studies have shown that changes in global mean 

climate will have a significant, regionally dependent influence on future ozone concentrations and the correspondence between the return of ozone to pre-1960 or 1980 

levels and the return of ESC amounts to pre-1960 or 1980 

levels (Waugh et al., 2009; Shepherd and Jonsson, 2008). 

The impact of coupling CCMs to interactive ocean models 

remains to be assessed since the number of CCMs able to 

perform this coupling remains small.

Future assessments should also consider brominated 

very short-lived species (VSLS, see Section 3.2.1), as 

well as the uncertainty in how the organic halogen lower 

boundary condition is prescribed in CCMs. Currently, 

projections of future organic halogen loadings are based 

on projected emission rates and an estimate of the global 

atmospheric lifetime of each organic halogen. These factors are used to create time-dependent volume mixing ratio 

lower boundary conditions that are then used to force the 

CCMs. However, the destruction of each halogen in the 

CCMs is dependent on the tropical upwelling and meridional mixing (both linked to the strength and structure of 

the BDC), and chemical loss rates (e.g., photolysis rates). 

The CCM-derived halogen lifetimes can be very different 

from the lifetimes assumed for the given projection scenario. Some studies (Douglass et al., 2008) suggest that 

the use of fixed mixing ratio lower boundary conditions 

provides an artificial constraint on model-derived ozone 

return dates. Furthermore, multi-model studies are necessary to fully evaluate the impact of ODS flux boundary 

conditions on ozone projections.

3.4 PROJECTIONS OF UV CHANGES RELATED 

TO OZONE CHANGES THROUGH THE 21ST 

CENTURY

Future changes in stratospheric ozone will cause 

changes of opposite sign in solar UV radiation received 

at the surface. However, UV radiation is also affected by 

a number of other factors (see Section 3.2.8), which are 

likely to change in the future. In this section only changes 

due to ozone are addressed. Simulations of future UV are 

based on ozone projections by CCMs, which take into account the effects of climate change on ozone (see Section 

3.3). The effect of changes in ozone on erythemal irradiance can be quantified either with empirical relationships, 

such as the radiation amplification factor (RAF) concept 

(Booth and Madronich, 1994; Madronich, 2007), or, more 

accurately, with radiative transfer models.

3.4.1 Midlatitude and Tropical UV

Calculations with a radiative transfer model were 

used by Tourpali et al. (2009) to simulate the noontime 

erythemal solar irradiance received at the Earth’s surface 

under cloud-free conditions. These simulations were 

based only on total ozone columns, and vertical profiles of 

ozone and temperature, derived from 11 CCMs (CCMVal-1 

models) simulating the evolution of stratospheric ozone 

(Eyring et al., 2006; Eyring et al., 2007). Following a peak 

between the late 1990s and early 2000s, erythemal irradiance was projected to decrease at all latitudes and in all 

seasons during the 21st century. Since erythemal irradiance changes were mostly driven by changes in total 

ozone, they tended to follow the pattern of the stratospheric ozone changes. The weakest changes, of the order 

of a few percent, were found in the tropics. At midlatitudes, the decreases ranged from 5–15% between 2000 

and 2100, while at southern high latitudes the decrease 

was a factor two stronger because of the projected recovery of Antarctic ozone layer there.

A follow-up study (Kazantzidis et al., 2010) provided quantitative estimates for the effect of the projected 

ozone changes on different biological weightings of UV 

irradiance (doses), such as DNA damage and vitamin D 

production, which have different sensitivities to ozone 

changes. For example, in some cases, changes in doses 

relevant to DNA damage can exceed those for erythema 

by a factor of two.
3.43

Future Ozone and UV

Ozone recovery will be affected by climate change 

(see also Section 3.2) and this will in turn affect the levels of UV radiation at the surface. Hegglin and Shepherd 

(2009) reported that the effects of a strengthened BDC 

on stratospheric ozone, which they entirely attributed to 

changes in GHG concentrations (assuming that the effect 

of ODSs on ozone were negligible in 1965 and 2095 in 

their simulations), would result in decreases in erythemal 

irradiance from 1965 to 2095 by about 9% at northern high 

latitudes. This change was found to be about three times 

larger than the change attributed to stratospheric ozone recovery from ODSs only, as estimated from the change in 

erythemal irradiance from 1965 to 2000.

More robust ozone projections can improve the 

accuracy of the estimated changes in surface UV irradiance. A new subset of simulations from 15 CCMs has 

become recently available through the CCMVal-2 activity 

of SPARC (see SPARC CCMVal (2010) and Section 3.3), 

and are used to repeat the erythemal irradiance calculations reported by Tourpali et al. (2009). More realistic and 

spatially varying climatological values for aerosols (Kinne 

et al., 2006) and surface reflectivity (Herman et al., 2001) 

are prescribed in the present calculations. Although cloud 

and aerosol fields are expected to change in the future 

(IPCC, 2007), they are kept constant here. For the period 

covered by the CCM simulations (1965–2100) and over a 

global grid of 10° × 15° respectively, solar UV irradiance 

spectra (280–400 nm) at the surface are calculated with 

the radiative transfer model package libRadtran (Mayer 

and Kylling, 2005) for the 15th of each month.

These new simulations, shown in Figure 3-18, result in lower inter-model spread compared to the results 

of Tourpali et al. (2009), allowing a more robust estimation of the dates when erythemal irradiance is projected 

to return to 1980 values. At midlatitudes, erythemal irradiance is projected to decrease for all models throughout 

the 21st century. The return of the multi-model mean to 

AMTRAC3

CAM3.5

CCSRNIES

CMAM

CNRM-ACM

EMAC

E39CA

GEOSCCM

LMZDrepro

MRI

NIWA-SOCOL

SOCOL

ULAQ

UMSLIMCAT

UMUKCA-UCAM

WACCM

Multi-model mean

1960 1980 2000 2020 2040 2060 2080 2100

Year

-20

-10

0

10

20

Relative change in erythemal irradiance (%)

60°N - 35°N

1960 1980 2000 2020 2040 2060 2080 2100

Year

-10

-5

0

5

10

Relative change in erythemal irradiance (%)

25°N - 25°S

1960 1980 2000 2020 2040 2060 2080 2100

Year

-20

-10

0

10

20

Relative change in erythemal irradiance (%)

35°S - 60°S

Northern Midlatitudes - Annual Mean

Tropics - Annual Mean

Southern Midlatitudes - Annual Mean

Figure 3-18. Annual means of surface erythemal 

irradiance changes (in %, relative to 1975–1985) 

under cloud-free conditions for three latitude belts 

representative for Northern Hemisphere midlatitudes 

(upper panel), tropics (middle panel), and Southern 

Hemisphere midlatitudes (lower panel), calculated 

with a radiative transfer model using projections of 

ozone and temperature from 15 CCMs. See Table 

3-1 for model descriptions. For the models providing multiple runs, the average change of irradiance 

is shown. The black thick line represents the multimodel mean. All lines have been smoothed with a 

1:2:1 filter. Note the different scale for the tropical 

belt. Updated from Tourpali et al. (2009) using projections from 15 CCMVal-2 models.
3.44

Chapter 3

1980 values is projected to occur in 2023 for the Northern 

Hemisphere and in 2035 for the Southern Hemisphere; 

for individual models, the return date ranges from 2016 

to 2055 and from 2038 to 2055, respectively. The large 

range on the erythemal irradiance return date reflects the 

large inter-model differences in total column ozone used 

as inputs in the surface UV calculations. The projected 

increases in column ozone above its levels in 1980 after 

the mid-21st century, especially in the Northern Hemisphere, result in decreases in erythemal irradiance of up 

to 5% below its levels in 1980 (Figure 3-18, upper panel), 

with important implications for ecosystems at high latitudes (UNEP, 2010). For example, such decreases may 

contribute to further deficiency in the vitamin D levels of 

humans (Edvardsen et al., 2007; Kazantzidis et al., 2009; 

Kimlin et al., 2007; Webb and Engelsen, 2006), unless 

other factors, such as a possible decrease in cloudiness, 

would counteract these changes. In the event of increasing cloudiness due to climate change, or increasing aerosol load (likely over inhabited areas), the adverse effects 

concerning vitamin D production would be even larger.

The simulations for the tropics (Figure 3-18, middle panel) show a steady increase in surface erythemal 

irradiance up to the late 1990s and then a slow decrease 

until ~2050, when it reaches a minimum that is still higher than 1980 levels for most models. Thereafter, surface UV is projected to slowly increase toward the end 

of the 21st century, in response to the projected decreases 

in tropical column ozone due to the acceleration of the 

BDC. The tropical changes are small (at most 5% with 

respect to 1980 values) in comparison to the changes 

projected for the higher latitudes. This temporal behavior of tropical erythemal irradiance is common to most 

models and agrees with the simple estimates of Hegglin 

and Shepherd (2009). UV-B radiation in the tropics is 

high due to naturally occurring low total ozone columns 

and high solar elevation angles. Assuming no variations 

from other factors, if UV-B radiation remains above its 

values in 1980, the adverse effects on ecosystems from 

UV exposure might be enhanced in this region during the 

21st century.

The uncertainty of the surface UV projections presented here could be much larger than the model range 

indicates. Indeed, more realistic projections of future UV 

radiation should include not only ozone, but also clouds, 

aerosols, and surface reflectivity. All these will likely be 

affected by climate change. In addition, anthropogenic 

tropospheric ozone and aerosols in the lower troposphere 

are likely to change as well (see Section 3.2.8). Clouds 

are the main modulator of UV radiation and current work 

shows that they may introduce considerable variability 

to the future UV time series and modify, or even reverse, 

the ozone-related trends in UV radiation. The evolution 

of these factors in the future, with patterns exhibiting 

strong spatial variations (e.g., Trenberth and Fasullo, 

2009), may modify accordingly the projected return date 

of UV irradiance to its 1980 levels.

3.4.2 Polar UV

In polar regions, the UV signatures of ozone reductions have been more pronounced compared to midlatitudes (see Chapter 2), resulting in measurable changes and, 

in some cases (e.g., in Antarctica), detectable increases in 

surface UV radiation (Bernhard et al., 2006).

In Tourpali et al. (2009), the ozone recovery projected by CCMs for the southern polar latitudes resulted 

in large reductions in erythemal irradiance under cloudless conditions of more than 30% between 2000 and 2095. 

When, instead of the annual mean, monthly means are 

considered, reductions of ~50% are found in the austral 

spring. The sign of the erythemal irradiance changes is 

reversed for polar regions, when its levels in the 1960s are 

taken as reference. Hegglin and Shepherd (2009) estimated increases in erythemal irradiance of up to 20% between 

1965 and 2095 over the same region in late spring and 

early summer. This increase is nearly half of that generated by the Antarctic ozone hole from the increase in 

ODS levels since the 1960s. The increase is not so large in 

the new calculations, with the multi-model mean in 2100 

being just above the 1960 level, though the range in the 

model estimates is large (see Figure 3-19). Some CCMs 

even report small decreases in erythemal irradiance for the 

same period.

The new UV simulations show that increases in 

clear-sky erythemal irradiance reach a rather broad maximum centered over the 2000s (Figure 3-19). The magnitude of this maximum varies between models, ranging, for 

the spring months, from ~5 to 15% in the Arctic, and from 

~30 to 110% in Antarctica. This spread in the model estimates is a measure of the uncertainty of the simulations, 

which, together with their large year-to-year variability, 

makes it difficult to assess more accurately the magnitude 

and timing of the UV maximum. After the 2000–2010 

decade, erythemal irradiance is projected to decrease at 

different rates for the two regions. In the Arctic, it is projected to return to its 1980 value between 2020 and 2030, 

while in Antarctica the return occurs only in the 2nd half of 

the 21st century. The rather slow decline in UV levels over 

Antarctica may have important biological consequences 

for Antarctic ecosystems, which will continue to be exposed to excess UV for most of the 21st century.

The expected increase in stratospheric ozone levels at southern polar latitudes is projected to accelerate 

surface warming over Antarctica (Son et al., 2009). This 

may accelerate ice and/or snow melting and, ultimately, 

may lead to exposure of organisms under the ice to UV 
3.45

Future Ozone and UV

1960 1980 2000 2020 2040 2060 2080 2100

Year

-30

-20

-10

0

10

20

30

Relative change in erythemal irradiance (%) 

90°N - 60°N

Northern Polar Latitudes - April

1960 1980 2000 2020 2040 2060 2080 2100

Year

-50

0

50

100

150

60°S - 90°S

Southern Polar Latitudes - October

Relative change in erythemal irradiance (%) 

Figure 3-19. As in Figure 3-18 but for April in the Arctic (upper) and October in Antarctica (lower). The legend 

of Figure 3-18 applies also here. Note the different scales in the two panels. Updated from Tourpali et al. (2009) 

using projections from 15 CCMVal-2 models.
3.46

Chapter 3

radiation and to a decrease of surface reflectivity. The reduction in surface reflectivity would reinforce the effect 

of rising stratospheric ozone levels by decreasing further 

surface UV irradiance over this region.

3.4.3 Link to UNEP Environmental Effects 

Panel Assessment

Changes in solar UV radiation at the surface are important due to the biological consequences—both negative 

and positive—for humans and for different ecosystems. 

The adverse effects of UV radiation on human health have 

dominated the public awareness during the last three 

decades because of ozone depletion. In recent years, 

though, much attention has been drawn to the benefits of 

solar UV radiation with respect to its involvement in the 

production of vitamin D, an important agent for human 

health (Edvardsen et al., 2007; Kazantzidis et al., 2009; 

McKenzie et al., 2009; Webb et al., 1988). As the spectral 

characteristics of solar UV radiation at the surface depend 

on the changes in ozone, the recovery of the stratospheric 

ozone layer will reduce the harmful biological doses 

received by humans, but will reduce also the rate of production of vitamin D, especially in the winter months at 

high latitudes. Ozone depletion has also influenced other 

communities, terrestrial and marine, as well as biological 

and chemical processes in the environment (UNEP, 2007). 

Therefore, the recovery of stratospheric ozone and the 

timing of this recovery are important for both. The impacts of ozone depletion and recovery, and of the resulting 

changes in UV radiation on the environment and the ecosystems, will be focal points for discussion in the forthcoming assessment report of the Environmental Effects 

Assessment Panel (EEAP) of UNEP (UNEP, 2010).

Ozone depletion has shown seasonal and latitudinal 

variations. The projections of surface UV irradiance in the 

21st century based on ozone changes projected by CCMs 

(as described in Section 3.4.1) reveal the importance of 

ozone variations in different regions. This can be seen in 

Figure 3-20, which shows the average change in daily erythemal irradiance between 1975–1985 and 2089–2099 for 

different seasons. Evidently there is a strong meridional 

gradient in the UV changes, which are small and positive 

in the tropics (up to 4%) and become negative with increasing magnitude at the middle and high latitudes. Although 

small in percentage, this increase in the tropics may be 

important for ecosystems living in this region, where irradiance levels are already high. As discussed above, the 

changes over Antarctica exceed −30% and are rather zonally symmetric. Such large decreases will have implications for the Antarctic ecosystems. Overall, the changes in 

UV radiation are fairly zonal over the middle and tropical 

latitudes and for all seasons, with longitudinal variations 

of a few percent. In the winter/spring period of the Northern Hemisphere these longitudinal variations are slightly 

larger due to more regional changes in column ozone.

In the present report, the future UV radiation has 

been simulated only in terms of ozone variations. However, climate change may significantly impact surface UV 

levels either directly through changes in clouds, albedo, 

and aerosols, or indirectly through interaction with ozone, 

as discussed in Section 3.2.8. Clear-sky UV simulations, 

calculated from changes in ozone and changes in cloudiness and surface reflectivity taken from climate models, 

will be discussed in more detail in the forthcoming assessment report of the EEAP/UNEP (UNEP, 2010).

Because ozone depletion has been most severe in 

the southern high and polar latitudes, the effects on the 

ecosystems would likely be most pronounced and detectable over these regions. Antarctic communities have been 

exposed to alterations in the surface conditions of temperature, moisture, and ultraviolet radiation resulting from 

climate change and stratospheric ozone depletion (Solomon et al., 2005). After three decades of ozone depletion 

and enhanced UV exposure, communities may already be 

undergoing adaptation, species selection, and changes in 

species assemblages, communities, and distributions, with 

effects on ecosystem function (Wall, 2007). For the Antarctic terrestrial communities, the timing of exposure in 

early spring might be more important than the magnitude 

of UV-B exposure, since, in early spring, organisms may 

be in a physiologically inactive state and unaccustomed 

to these levels (Wall, 2007). Terrestrial organisms are 

also living under the snow cover that may provide protection from UV radiation (Cockell and Cordoba-Jabonero, 

2004). Climate changes may result either in prolongation (through cooling) or in shrinking (through increases 

in sublimation and strengthening of katabatic winds) of 

periods with snow cover (Wall, 2007).

3.5 CONCLUSIONS

The focus of this chapter has been to assess and discuss projections of future ozone and its impact on surface 

UV. Future ozone will be mostly determined by changes 

in stratospheric halogen loading (ESC) and changes in climate parameters. Thus, the chronology of the third and 

last stage of ozone recovery from ODSs as defined in 

WMO (2007), and employed here, depends critically on 

the combination of these two factors. Separately evaluating the role of halogens in ozone depletion is of great 

interest because it is the primary concern of the Parties to 

the Montreal Protocol. Section 3.2 describes the primary 

factors controlling future ozone and surface UV. Stratospheric halogen loading has been the strongest influence 

on ozone in the last decades since the onset of the Antarctic ozone hole and midlatitude depletion. Stratospheric 
3.47

Future Ozone and UV

halogen loading has peaked in all regions of the stratosphere in the last two decades and is now diminishing at 

a rate primarily controlled by the atmospheric lifetimes 

of individual chlorine and bromine gases. Global and regional projections for the removal of halogens have very 

similar time dependences as shown in Figure 3-21 and indicate that the return of halogen amounts to 1960 values 

will not be achieved before 2100, except in the tropical 

lower stratosphere. A prominent feature of the halogen 

time series is that regional peak values are quite different. 

Because of the long average lifetimes of ODSs, the highest 

values are found in polar regions, where the average age of 

stratospheric air is greatest.

Other leading factors discussed in Section 3.2 are 

stratospheric levels of key ozone-destroying hydrogen and 

nitrogen species, stratospheric temperatures, the BrewerDobson circulation, dynamics of the polar vortex, and 

stratospheric aerosols. Many of these factors, which already influence ozone to some degree, are coupled such 

that the effects of an initial perturbation may be amplified to result in a larger effect on ozone than would have 

been the case otherwise. Links between components of 

the chemistry-climate system are indicated in Figure 3-22 

with arrows representing chemistry, radiation, transport 

and other mechanisms. Many of these factors are influenced by anthropogenic emissions of GHGs and ODSs. 

Latitude (deg)

Latitude (deg)

Latitude (deg)

Latitude (deg)

Longitude (deg) Longitude (deg)

Longitude (deg) Longitude (deg)

January

April October

July

-180 -120 -60 0 60 120 180

-90

-60

-30

0

30

60

90

-180 -120 -60 0 60 120 180

-90

-60

-30

0

30

60

90

-180 -120 -60 0 60 120 180

-90

-60

-30

0

30

60

90

-180 -120 -60 0 60 120 180

-90

-60

-30

0

30

60

90

min = -33%

-18-16 -14-12 -10 -8 -6 -4 -2 0 2468 10 12 14 18

Change (%)

Figure 3-20. Multi-model average changes in surface erythemal irradiance between 2089–2099 and 1975–

1985 under cloud-free conditions for four months, calculated with a radiative transfer model using projections 

of ozone and temperature from 15 CCMs. Updated from Tourpali et al. (2009) using projections from 15 

CCMVal-2 models.
3.48

Chapter 3

For example, the increase in stratospheric CO2 abundances is the main cause of the cooling trend in the stratosphere 

(see Section 4.2).

These multiple drivers of ozone changes and their 

interactions will create more diverse regional behavior in 

stratospheric ozone than would be expected from average global behavior or from the evolution of ODSs alone. 

This point is illustrated schematically in the time series 

shown in Figure 3-21, where an anticorrelation between 

the evolutions of ESC and total column ozone can be 

seen for the global average, for the Antarctic lower stratosphere, and for the tropical upper stratosphere. However, 

the influence of other drivers, in particular GHGs, is manifest in the global and tropical upper stratosphere, where 

ozone amounts return to 1960 levels in the middle of 21st

century, although ODSs are still enhanced during this period. It is only in the Antarctic lower stratosphere, where 

the sensitivity of ozone to ESC maximizes, that ESC appears to remain the dominant driver until the end of the 

21st century. In other regions, ozone changes are not well 

correlated with ESC. For example, in the tropical lower 

stratosphere, ozone amounts decrease steadily from 1960 

to 2100 whereas in the northern midlatitude lower stratosphere, ozone shows a small negative response to ESC 

until 2020 before increasing throughout the 21st century. 

In these regions, GHG changes are the dominant driver of 

ozone changes.

The multiple interactions between the components 

of the chemistry-climate system (Figure 3-22) complicate 

a clean attribution of changes in ozone to changes in ODSs 

and other factors as required by the definition of full ozone 

recovery employed in this Assessment. For example, decreases in polar stratospheric ozone, resulting from increases in ESC, cool the stratosphere since heating from 

UV absorption by ozone is also reduced. This cooling in 

turn enhances the effectiveness of ESC by promoting the 

formation of polar stratospheric clouds. Increases in GHG 

concentrations drive additional radiative stratospheric 

cooling, but at the same time strengthen the BDC, which 

adiabatically warms the polar stratosphere and enhances 

mixing with lower latitudes. Sensitivity simulations using CCMs can be used to disentangle the effects of different drivers to ozone changes. Model results indicate that 

changes in ODSs and GHGs appear to affect ozone nearly 

independently.

The primary drivers of ozone changes, and their 

interactions, are largely described in the CCMs that have 

been used to make the ozone and UV projections presented in this chapter. To identify the causes of simulated 

future ozone changes, CCMs were also used to conduct 

sensitivity simulations in which one particular driver of 

ozone changes (e.g., ODSs or GHGs) was held constant. 

There are, however, processes that are not yet fully realized in models, such as potential changes in the quasibiennial oscillation and coupling to the ocean. There are 

additional processes and associated potential feedbacks 

in the ocean-climate-stratosphere system (e.g., changes 

in marine biogenic emissions of halocarbons induced by 

changes in climate and surface UV) and in the terrestrial 

biosphere-climate-stratosphere system (e.g., emissions of 

ozone precursors) that are also currently absent in CCMs. 

Therefore, as progress is made toward the development of 

full Earth Systems Models that include a wider variety of 

processes and feedbacks, more comprehensive and robust 

future projections of ozone will become possible.

Figure 3-21. A schematic contrasting global and 

vertically resolved changes in ozone (blue) and ESC 

(red). In all cases changes with respect to the 1960 

value are shown. Panel (a) shows the most common 

framework in which future projections of ozone are 

considered (compare to Figure 6-1 of WMO, 2007) 

while the other panels show the greater variety of responses of ozone to ESC in different regions of the 

atmosphere resulting from non-ESC drivers of ozone 

changes and their feedbacks and interactions (see 

Figure 3-22). The schematic is generated from multimodel means (solid lines) and 95% prediction intervals (shading) of REF-B2 simulations (Table 3-2) 

from the suite of CCMs used in this chapter.
3.49

Future Ozone and UV

REFERENCES

Akiyoshi, H., L.B. Zhou, Y. Yamashita, K. Sakamoto, 

M. Yoshiki, T. Nagashima, M. Takahashi, J. Kurokawa, M. Takigawa, and T. Imamura, A CCM 

simulation of the breakup of the Antarctic polar 

vortex in the years 1980-2004 under the CCMVal 

scenarios, J. Geophys. Res., 114, D03103, doi: 

10.1029/2007JD009261, 2009.

Austin, J., and R.J. Wilson, Ensemble simulations of 

the decline and recovery of stratospheric ozone, 

J. Geophys. Res., 111, D16314, doi: 10.1029/

2005JD006907, 2006.

Austin, J., and R.J. Wilson, Sensitivity of polar ozone 

to sea surface temperatures and halogen amounts, 

J. Geophys. Res., 115, D18303, doi: 10.1029/ 

2009JD013292, 2010.

Austin, J., N. Butchart, and K.P. Shine, Possibility of an 

Arctic ozone hole in a doubled-CO2 climate, Nature, 360, 221-225, doi: 10.1038/360221a0, 1992.

Austin, J., J. Wilson, F. Li, and H. Vömel, Evolution of 

water vapor concentrations and stratospheric age 

of air in coupled chemistry-climate model simulations, J. Atmos. Sci., 64, 905-921, doi: 10.1175/

JAS3866.1, 2007.

Austin, J., K. Tourpali, E. Rozanov, H. Akiyoshi, S. Bekki, G. Bodeker, C. Brühl, N. Butchart, M. Chipperfield, M. Deushi, V. Fomichev, M. Giorgetta, L. 

Gray, K. Kodera, F. Lott, E. Manzini, D. Marsh, K. 

Matthes, T. Nagashima, K. Shibata, R.S. Stolarski, 

H. Struthers, and W. Tian, Coupled chemistry climate model simulations of the solar cycle in ozone 

and temperature, J. Geophys. Res., 113, D11306, 

doi: 10.1029/2007JD009391, 2008.

Austin, J., J. Scinocca, D. Plummer, L. Oman, D. Waugh, 

H. Akiyoshi, S. Bekki, P. Braesicke, N. Butchart, M. 

Chipperfield, D. Cugnet, M. Dameris, S. Dhomse, 

V. Eyring, S. Frith, R.R. Garcia, H. Garny, A. 

Gettelman, S.C. Hardiman, D. Kinnison, J.F. Lamarque, E. Mancini, M. Marchand, M. Michou, O. 

Morgenstern, T. Nakamura, S. Pawson, G. Pitari, 

J. Pyle, E. Rozanov, T.G. Shepherd, K. Shibata, H. 

Teyssèdre, R.J. Wilson, and Y. Yamashita, Decline 

and recovery of total column ozone using a multimodel time series analysis, J. Geophys. Res., 115, 

D00M10, doi: 10.1029/2010JD013857, 2010a.

Austin, J., H. Struthers, J. Scinocca, D.A. Plummer, H. 

Akiyoshi, A.J.G. Baumgaertner, S. Bekki, G.E. 

Bodeker, P. Braesicke, C. Brühl, N. Butchart, M.P. 

Chipperfield, D. Cugnet, M. Dameris, S. Dhomse, 

S. Frith, H. Garny, A. Gettelman, S.C. Hardiman, 

P. Jöckel, D. Kinnison, A. Kubin, J.F. Lamarque, 

U. Langematz, E. Mancini, M. Marchand, M. Michou, O. Morgenstern, T. Nakamura, J.E. Nielsen, 

G. Pitari, J. Pyle, E. Rozanov, T.G. Shepherd, K. 

Shibata, D. Smale, H. Teyssèdre, and Y.Yamashita, 

Chemistry-climate model simulations of spring 

Antarctic ozone, J. Geophys. Res., 115, D00M11, 

Heterogeneous Chemistry

Chlorine

(Polar Lower Strat.) Bromine

Chemistry

Radiation

Transport

Other

Stratospheric Aerosol

Thermal Wind

Nitrous Oxide

Methane

Heating

Rates

Ozone Gas-Phase Destruction

(Upper Stratosphere)

Relationship

Temperature

Dynamics

Ozone

Water Vapor

Adiabatic Proc.

Advection

ODSs GHGs

Saturation

Figure 3-22. Schematic of ozonefocused stratospheric chemistryclimate interactions. Links between 

elements of the chemistry-climate 

system are indicated with arrows representing chemistry (blue), radiation 

(red), transport (green), and other 

mechanisms (black). Simple and 

more complex feedback cycles can 

be constructed following the linking 

mechanisms. A simple example is 

ozone depletion in the upper stratosphere leading to lower temperatures. 

Lower temperatures slow the gasphase destruction of ozone, thereby 

reducing the amount of ozone depletion. The feedback loops involving 

stratospheric aerosol are currently 

not important, but they might become 

important again after a large volcanic eruption or in the context of some 

geoengineering proposals.
3.50

Chapter 3

doi: 10.1029/2009JD013577, 2010b.

Avallone, L.M., and M.J. Prather, Photochemical evolution of ozone in the lower tropical stratosphere, J. 

Geophys. Res., 101 (D1), 1457-1461, 1996.

Badosa, J., R.L. McKenzie, M. Kotkamp, J. Calbó, J.A. 

González, P.V. Johnston, M. O’Neill, and D.J. Anderson, Towards closure between measured and 

modelled UV under clear skies at four diverse sites, 

Atmos. Chem. Phys., 7, 2817-2837, 2007.

Bais, A.F., and D. Lubin (Lead Authors), 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, and C.S. Zerefos, Surface ultraviolet radiation: Past, Present, and Future, Chapter 7 in Scientificttt Assessment of Ozone Depletion: 2006, Global 

Ozone Research and Monitoring Project–Report 

No. 50, 572 pp., World Meteorological Organization, Geneva, Switzerland, 2007.

Baldwin, M., and M. Dameris (Lead Authors), 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, and D. Thompson, Climateozone connections, Chapter 5 in Scientific Assessment of Ozone Depletion: 2006, Global Ozone Research and Monitoring Project–Report No. 50, 572 

pp., World Meteorological Organization, Geneva, 

Switzerland, 2007.

Bell, C.J., L.J. Gray, and J.E. Kettleborough, Changes in 

Northern Hemisphere stratospheric variability under increased CO2 concentrations, Quart. J. Roy. 

Meteorol. Soc., 136, 1181-1190, doi: 10.1002/

qj.633, 2010.

Bernhard, G., C.R. Booth, J.C. Ehramjian, and S.E. Nichol, UV climatology at McMurdo Station, Antarctica, based on version 2 data of the National Science 

Foundation’s Ultraviolet Radiation Monitoring 

Network, J. Geophys. Res., 111, D11201, doi: 

10.1029/2005JD005857, 2006.

Bodeker, G.E., and D.W. Waugh (Lead Authors), H. 

Akiyoshi, P. Braesicke, V. Eyring, D.W. Fahey, 

E. Manzini, M.J. Newchurch, R.W. Portmann, 

A. Robock, K.P. Shine, W. Steinbrecht, and E.C. 

Weatherhead, The ozone layer in the 21st century, 

Chapter 6 in Scientific Assessment of Ozone Depletion: 2006, Global Ozone Research and Monitoring 

Project–Report No. 50, 572 pp., World Meteorological Organization, Geneva, Switzerland, 2007.

Bodeker, G.E., H. Struthers, and B.J. Connor, Dynamical containment of Antarctic ozone depletion, Geophys. Res. Lett., 29 (7), 1098, doi: 

10.1029/2001GL014206, 2002.

Bodeker, G.E., H. Shiona, and H. Eskes, Indicators of 

Antarctic ozone depletion, Atmos. Chem. Phys., 5, 

2603-2615, 2005.

Booth, C.R., and S. Madronich, Radiation amplification 

factors: Improved formulations accounts for large 

increases in ultraviolet radiation associated with 

Antarctic ozone depletion, in Ultraviolet Radiation in Antarctica: Measurements and Biological 

Effects, AGU Antarctic Research Series, edited by 

C.S. Weiler and P.A. Penhale, 39-42, Washington, 

D.C., 1994.

Bradley, R.S., F.T. Keimig, and H.F. Diaz, Projected temperature changes along the American cordillera and 

the planned GCOS network, Geophys. Res. Lett., 

31, L16210, doi: 10.1029/2004GL020229, 2004.

Braesicke, P., and J.A. Pyle, Changing ozone and changing circulation in northern mid-latitudes: Possible 

feedbacks?, Geophys. Res. Lett. 30 (2), 1059, doi: 

10.1029/2002GL015973, 2003.

Brühl, C., and P.J. Crutzen, On the disproportionate role of 

tropospheric ozone as a filter against solar UV-B radiation, Geophys. Res. Lett., 16 (7), 703-706, 1989.

Butchart, N., and A.A. Scaife, Removal of chlorofluorocarbons by increased mass exchange between the 

stratosphere and troposphere in a changing climate, 

Nature, 410 (6830), 799-802, 2001.

Butchart, N., I. Cionni, V. Eyring, T.G. Shepherd, D.W. 

Waugh, H. Akiyoshi, J. Austin, C. Brühl, M.P. 

Chipperfield, E. Cordero, M. Dameris, R. Deckert, 

S. Dhomse, S.M. Frith, R.R. Garcia, A. Gettelman, 

M.A. Giorgetta, D.E. Kinnison, F. Li, E. Mancini, 

C. McLandress, S. Pawson, G. Pitari, D.A. Plummer, E. Rozanov, F. Sassi, J.F. Scinocca, K. Shibata, B. Steil, and W. Tian, Chemistry-climate model 

simulations of 21st century stratospheric climate 

and circulation changes, J. Clim., 23 (20), 5349-

5374, doi: 10.1175/2010JCLI3404.1, 2010.

Caldeira, K., and L. Wood, Global and Arctic climate engineering: Numerical model studies, Phil. Trans. 

Roy. Soc. A, 366 (1882), 4039-4056, 2008.

Charlton-Perez, A.J., L.M. Polvani, J. Austin, and F. Li, 

The frequency and dynamics of stratospheric sudden warmings in the 21st century, J. Geophys. Res., 

113, D16116, doi: 10.1029/2007JD009571, 2008.

Charlton-Perez, A.J., E. Hawkins, V. Eyring, I. Cionni, 

G.E. Bodeker, D.E. Kinnison, H. Akiyoshi, S.M. 

Frith, R. Garcia, A. Gettelman, J.F. Lamarque, T. 

Nakamura, S. Pawson, Y. Yamashita, S. Bekki, 

P. Braesicke, M.P. Chipperfield, S. Dhomse, M. 

Marchand, E. Mancini, O. Morgenstern, G. Pitari, 

D. Plummer, J.A. Pyle, E. Rozanov, J. Scinocca, K. 

Shibata, T.G. Shepherd, W. Tian, and D.W. Waugh, 

The potential to narrow uncertainty in projections 

of stratospheric ozone over the 21st century, Atmos. 

Chem. Phys., 10, 9473-9486, doi: 10.5194/acp-10-
3.51

Future Ozone and UV

9473-2010, 2010.

Chipperfield, M.P., New version of the TOMCAT/SLIMCAT off-line chemical transport model: Intercomparison of stratospheric tracer experiments, Quart. 

J. Roy. Meteorol. Soc., 132 (617), 1179-1203, doi: 

10.1256/qj.05.51, 2006.

Chipperfield, M.P., and W. Feng, Comment on: Stratospheric ozone depletion at northern mid-latitudes 

in the 21st century: The importance of future concentrations of greenhouse gases nitrous oxide and 

methane, Geophys. Res. Lett., 30 (7), 1389, doi: 

10.1029/2002GL016353, 2003.

Chipperfield, M.P., and V.E. Fioletov (Lead Authors), 

B. Bregman, J. Burrows, B.J. Connor, J.D. Haigh, 

N.R.P. Harris, A. Hauchecorne, L.L. Hood, S.R. 

Kawa, J.W. Krzyścin, J.A. Logan, N.J. Muthama, 

L. Polvani, W.J. Randel, T. Sasaki, J. Stähelin, R.S. 

Stolarski, L.W. Thomason, and J.M. Zawodny, 

Global ozone: Past and Present, Chapter 3 in Scientific Assessment of Ozone Depletion: 2006, Global 

Ozone Research and Monitoring Project–Report 

No. 50, 572 pp., World Meteorological Organization, Geneva, Switzerland, 2007.

Chubarova, N.Y., UV variability in Moscow according 

to long-term UV measurements and reconstruction 

model, Atmos. Chem. Phys., 8, 3025-3031, 2008.

Chubarova, N.Y., Seasonal distribution of aerosol properties over Europe and their impact on UV irradiance, 

Atmos. Meas. Tech., 2, 593-608, doi: 10.5194/amt2-593-2009, 2009.

Chubarova, N.Y., N.G. Prilepsky, A.N. Rublev, and A.R. 

Riebau, A mega-fire event in central Russia: Fire 

weather, radiative, and optical properties of the 

atmosphere, and consequences for subboreal forest plants, in Developments in Environmental Science, Volume 8, edited by A. Bytnerowicz, M. 

Arbaugh, A. Riebau, and C. Andersen, 247-264, 

Elsevier B.V., Amsterdam, The Netherlands, available: http://www.fs.fed.us/psw/publications/4451/

psw_2009_4451-001_247-264.pdf, 2009.

CIE (Commission Internationale de l’Éclairage), Reference Action spectra for ultraviolet induced erythema and pigmentation of different human skin types, 

in Technical Collection 103/3, p.15-21, 1993.

Clarke, L., J. Edmonds, H. Jacoby, H. Pitcher, J. Reilly, 

and R. Richels, Scenarios of Greenhouse Gas 

Emissions and Atmospheric Concentrations. Subreport 2.1A of Synthesis and Assessment Product 

2.1 by the U.S. Climate Change Science Program 

and the Subcommittee on Global Change Research. 

Department of Energy, Office of Biological & Environmental Research, Washington, D.C., U.S.A., 

154 pp., 2007.

Clement, A.C., R. Burgman, and J.R. Norris, Observational and model evidence for positive low-level 

cloud feedback, Science, 325 (5939), 460-464, doi: 

10.1126/science.1171255, 2009.

Cockell, C.S., and C. Córdoba-Jabonero, Coupling of 

climate change and biotic UV exposure through 

changing snow-ice covers in terrestrial habitats, 

Photochem. Photobiol., 79 (1), 26-31, 2004.

Comiso, J.C., C.L. Parkinson, R. Gersten, and L. 

Stock, Accelerated decline in the Arctic sea ice 

cover, Geophys. Res. Lett., 35, L01703, doi: 

10.1029/2007GL0319, 2008.

Crutzen, P.J., Albedo enhancement by stratospheric sulfur injections: A contribution to resolve a policy 

dilemma?, Clim. Change, 77 (3-4), 211-219, doi: 

10.1007/s10584-006-9101-y, 2006.

de Grandpré, J., S.R. Beagley, V.I. Fomichev, E. Griffioen, 

J.C. McConnell, A.S. Medvedev, and T.G. Shepherd, 

Ozone climatology using interactive chemistry: Results from the Canadian Middle Atmosphere Model, 

J. Geophys. Res., 105 (D21), 26475-26491, 2000.

Déqué, M., Frequency of precipitation and temperature 

extremes over France in an anthropogenic scenario: 

Model results and statistical correction according to 

observed values, Global Planet. Change, 57 (1-2), 

16-26, doi: 10.1016/j.gloplacha.2006.11.030, 2007.

Dhomse, S., M. Weber, I. Wohltmann, M. Rex, and J.P. 

Burrows, On the possible causes of recent increases 

in northern hemispheric total ozone from a statistical analysis of satellite data from 1979 to 2003, 

Atmos. Chem. Phys., 6, 1165-1180, 2006.

Douglass, A.R., R.S. Stolarski, M.R. Schoeberl, C.H. 

Jackman, M.L. Gupta, P.A. Newman, J.E. Nielsen, and E.L. Fleming, Relationship of loss, mean 

age of air and the distribution of CFCs to stratospheric circulation and implications for atmospheric lifetimes, J. Geophys. Res., 113, D14309, doi: 

10.1029/2007JD009575, 2008.

Dufresne, J.L., and S. Bony, An assessment of the primary sources of spread of global warming estimates 

from coupled atmosphere–ocean models, J. Clim., 

21 (19), 5135-5144, doi: 10.1175/2008JCLI2239.1, 

2008.

Edvardsen, K., M. Brustad, O. Engelsen, and L. Aksnes, 

The solar UV radiation level needed for cutaneous production of vitamin D3 in the face. A study 

conducted among subjects living at a high latitude 

(68°N), Photochem. Photobiol. Sci., 6, 57-62, doi: 

10.1039/b613263d, 2007.

Egorova, T., E. Rozanov, V. Zubov, E. Manzini, W. 

Schmutz, and T. Peter, Chemistry-climate model 

SOCOL: A validation of the present-day climatology, Atmos. Chem. Phys., 5, 1557-1576, 2005.

Engel, A., T. Möbius, H. Bönisch, U. Schmidt, R. Heinz, 

I. Levin, E. Atlas, S. Aoki, T. Nakazawa, S. Sug-
3.52

Chapter 3

awara, F. Moore, D. Hurst, J. Elkins, S. Schauffler, 

A. Andrews, and K. Boering, Age of stratospheric 

air unchanged within uncertainties over the past 30 

years, Nature Geoscience, 2, 28-31, doi: 10.1038/

ngeo388, 2009.

Eyring, V., N.R.P. Harris, M. Rex, T.G. Shepherd, D.W. 

Fahey, G.T. Amanatidis, J. Austin, M.P. Chipperfield, M. Dameris, P.M. De F. Forster, A. Gettelman, H.F. Graf, T. Nagashima, P.A. Newman, S. 

Pawson, M.J. Prather, J.A. Pyle, R.J. Salawitch, 

B.D. Santer, and D.W. Waugh, A strategy for 

process-oriented validation of coupled chemistryclimate models, Bull. Amer. Meteorol. Soc., 86 (8), 

1117-1133, doi: 10.1175/BAMS-86-8-1117, 2005.

Eyring, V., N. Butchart, D.W. Waugh, H. Akiyoshi, J. 

Austin, S. Bekki, G.E. Bodeker, B.A. Boville, C. 

Brühl, M.P. Chipperfield, E. Cordero, M. Dameris, 

M. Deushi, V.E. Fioletov, S.M. Frith, R.R. Garcia, 

A. Gettelman, M.A. Giorgetta, V. Grewe, L. Jourdain, D.E. Kinnison, E. Mancini, E. Manzini, M. 

Marchand, D.R. Marsh, T. Nagashima, P.A. Newman, J.E. Nielsen, S. Pawson, G. Pitari, D.A. Plummer, E. Rozanov, M. Schraner, T.G. Shepherd, K. 

Shibata, R.S. Stolarski, H. Struthers, W. Tian, and 

M. Yoshiki, Assessment of temperature, trace species, and ozone in chemistry-climate model simulations of the recent past, J. Geophys. Res., 111, 

D22308, doi: 10.1029/2006JD007327, 2006.

Eyring, V., D.W. Waugh, G.E. Bodeker, E. Cordero, H. 

Akiyoshi, J. Austin, S.R. Beagley, B.A. Boville, P. 

Braesicke, C. Brühl, N. Butchart, M.P. Chipperfield, 

M. Dameris, R. Deckert, M. Deushi, S.M. Frith, R.R. 

Garcia, A. Gettelman, M.A. Giorgetta, D.E. Kinnison, E. Mancini, E. Manzini, D.R. Marsh, S. Matthes, T. Nagashima, P.A. Newman, J.E. Nielsen, S. 

Pawson, G. Pitari, D.A. Plummer, E. Rozanov, M. 

Schraner, J.F. Scinocca, K. Semeniuk, T.G. Shepherd, K. Shibata, B. Steil, R.S. Stolarski, W. Tian, 

and M. Yoshiki, Multimodel projections of stratospheric ozone in the 21st century, J. Geophys. Res., 

112, D16303, doi: 10.1029/2006JD008332, 2007.

Eyring, V., M.P. Chipperfield, M.A. Giorgetta, D.E. Kinnison, E. Manzini, K. Matthes, P.A. Newman, S. 

Pawson, T.G. Shepherd, and D.W. Waugh, Overview of the new CCMVal reference and sensitivity 

simulations in support of upcoming ozone and climate assessments and the planned SPARC CCMVal 

Report, SPARC Newsletter No. 30, 20-26, 2008.

Eyring, V., I. Cionni, G.E. Bodeker, A.J. Charlton-Perez, 

D.E. Kinnison, J.F. Scinocca, D.W. Waugh, H. 

Akiyoshi, S. Bekki, M.P. Chipperfield, M. Dameris, 

S. Dhomse, S.M. Frith, H. Garny, A. Gettelman, A. 

Kubin, U. Langematz, E. Mancini, M. Marchand, T. 

Nakamura, L.D. Oman, S. Pawson, G. Pitari, D.A. 

Plummer, E. Rozanov, T.G. Shepherd, K. Shibata, 

W. Tian, P. Braesicke, S.C. Hardiman, J.F. Lamarque, O. Morgenstern, J.A. Pyle, D. Smale, and 

Y. Yamashita, Multi-model assessment of stratospheric ozone return dates and ozone recovery in 

CCMVal-2 models, Atmos. Chem. Phys., 10, 9451-

9472, doi: 10.5194/acp-10-9451-2010, 2010a.

Eyring, V., I. Cionni, J.-F. Lamarque, H. Akiyoshi, 

G.E. Bodeker, A.J. Charlton-Perez, S.M. Frith, 

A. Gettelman, D.E. Kinnison, T. Nakamura, L.D. 

Oman, S. Pawson, and Y. Yamashita, Sensitivity 

of 21st century stratospheric ozone to greenhouse 

gas scenarios, Geophys. Res. Lett., 37, L16807, doi: 

10.1029/2010GL044443, 2010b.

Fleming, E.L., C.H. Jackman, D.K. Weisenstein, and M.K.W. 

Ko, The impact of interannual variability on multidecadal total ozone simulations, J. Geophys. Res.,

112, D10310, doi: 10.1029/2006JD007953, 2007.

Fioletov, V.E., G.E. Bodeker, A.J. Miller, R.D. McPeters, 

and R. Stolarski, Global and zonal total ozone 

variations estimated from ground-based and satellite 

measurements: 1964-2000, J. Geophys. Res., 107

(D22), 4647 doi: 10.1029/2001JD001350, 2002.

Fomichev, V.I., W.E. Ward, S.R. Beagley, C. McLandress, J.C. McConnell, N.A. McFarlane, and T.G. 

Shepherd, Extended Canadian Middle Atmosphere 

Model: Zonal-mean climatology and physical parameterizations, J. Geophys. Res., 107 (D10), 4087, 

doi: 10.1029/2001JD000479, 2002.

Garcia, R.R., D.R. Marsh, D.E. Kinnison, B.A. Boville, 

and F. Sassi, Simulation of secular trends in the 

middle atmosphere, 1950–2003, J. Geophys. Res., 

112, D09301, doi: 10.1029/2006JD007485, 2007.

Garny, H., G.E. Bodeker, and M. Dameris, Trends and 

variability in stratospheric mixing: 1979-2005, Atmos. Chem. Phys., 7, 5611-5624, 2007.

Garny, H., M. Dameris, and A. Stenke, Impact of prescribed SSTs on climatologies and long-term trends 

in CCM simulations, Atmos. Chem. Phys., 9, 6017-

6031, 2009.

Gauss, M., G. Myhre, G. Pitari, M.J. Prather, I.S.A. Isaksen, T.K. Berntsen, G.P. Brasseur, F.J. Dentener, 

R.G. Derwent, D.A. Hauglustaine, L.W. Horowitz, 

D.J. Jacob, M. Johnson, K.S. Law, L.J. Mickley, J.-

F. Müller, P.-H. Plantevin, J.A. Pyle, H.L. Rogers, 

D.S. Stevenson, J.K. Sundet, M. van Weele, and 

O. Wild, Radiative forcing in the 21st century due 

to ozone changes in the troposphere and the lower 

stratosphere, J. Geophys. Res., 108 (D9), 4292, doi: 

10.1029/2002JD002624, 2003.

Gillett, N.P., H. Akiyoshi, S. Bekki, V. Eyring, R. Garcia, 

C.A. McLinden, A. Yu. Karpechko, D.A. Plummer, 

E. Rozanov, J. Scinocca, and K. Shibata, Attribution 

of observed changes in stratospheric ozone and tem-
3.53

Future Ozone and UV

perature, Atmos. Chem. Phys. Discuss., 10, 17341-

17367, doi: 10.5194/acpd-10-17341-2010, 2010.

Guillas, S., M.L. Stein, D.J. Wuebbles, and J. Xia, Using chemistry transport modeling in statistical 

analysis of stratospheric ozone trends from observations, J. Geophys. Res., 109, D22303, doi: 

10.1029/2004JD005049, 2004.

Hadjinicolaou, P., and J.A. Pyle, The impact of Arctic 

ozone depletion on northern middle latitudes: Interannual variability and dynamical control, J. Atmos. 

Chem., 47 (1), 25-43, 2004.

Hadjinicolaou, P., J.A. Pyle, M.P. Chipperfield, and J.A. 

Kettleborough, Effect of interannual meteorological variability on mid-latitude O3, Geophys. Res. 

Lett., 24 (23), 2993-2996, 1997.

Hawkins, E., and R. Sutton, The potential to narrow uncertainty in regional climate predictions, Bull. 

Amer. Meteorol. Soc., 90 (8), 1095-1107, doi: 

10.1175/2009BAMS2607, 2009.

Heckendorn, P., D. Weisenstein, S. Fueglistaler, B.P. Luo, 

E. Rozanov, M. Schraner, L.W. Thomason, and T. 

Peter, Impact of geoengineering aerosols on stratospheric temperature and ozone, Env. Res. Lett., 4 (4), 

045108, doi: 10.1088/1748-9326/4/4/045108, 2009.

Hegglin, M.I., and T.G. Shepherd, Large climate-induced 

changes in ultraviolet index and stratosphere-totroposphere ozone flux, Nature Geoscience, 2, 687-

691 doi: 10.1038/ngeo604, 2009.

Herman, J.R., D. Larko, E. Celarier, and J. Ziemke, Changes in the Earth’s UV reflectivity from the surface, 

clouds and aerosols, J. Geophys. Res., 106 (D6), 

5353-5368, doi: 10.1029/2000JD900435, 2001.

IPCC (Intergovernmental Panel on Climate Change), 

Emissions Scenarios. A Special Report of Working Group III of the Intergovernmental Panel on 

Climate Change, edited by N. Nakicenovic and R. 

Swart, 570 pp., Cambridge University Press, Cambridge, U.K., and New York, NY, U.S.A., 2000.

IPCC (Intergovernmental Panel on Climate Change), Climate Change 2007: The Physical Science Basis: 

Contribution of Working Group I to the Fourth Assessment Report of the Intergovernmental Panel on 

Climate Change, edited by Solomon, S., D. Qin, 

M. Manning, Z. Chen, M. Marquis, K.B. Averyt, 

M. Tignor, and H.L. Miller, 996 pp., Cambridge 

University Press, Cambridge, U.K., and New York, 

NY, U.S.A., 2007.

Jacobson, M.Z., and D.G. Streets, Influence of future anthropogenic emissions on climate, natural emissions, and air quality, J. Geophys. Res., 114, 

D08118, doi: 10.1029/2008JD011476, 2009.

Jöckel, P., H. Tost, A. Pozzer, C. Brühl, J. Buchholz, L. 

Ganzeveld, P. Hoor, A. Kerkweg, M.G. Lawrence, 

R. Sander, B. Steil, G. Stiller, M. Tanarhte, D. 

Taraborrelli, J. van Aardenne, and J. Lelieveld, The 

atmospheric chemistry general circulation model 

ECHAM5/MESSy1: Consistent simulation of 

ozone from the surface to the mesosphere, Atmos. 

Chem. Phys., 6, 5067-5104, doi: 10.5194/acp-6-

5067-2006, 2006.

Jonsson, A.I., V.I. Fomichev, and T.G. Shepherd, The 

effect of nonlinearity in CO2 heating rates on the 

attribution of stratospheric ozone and temperature 

changes. Atmos. Chem. Phys., 9, 8447-8452, doi: 

10.5194/acp-9-8447-2009, 2009.

Jourdain, L., S. Bekki, F. Lott, and F. Lefèvre, The coupled chemistry-climate model LMDz-REPROBUS: 

Description and evaluation of a transient simulation 

of the period 1980–1999, Ann. Geophys., 26, 1391-

1413, doi: 10.5194/angeo-26-1391-2008, 2008.

Kazadzis, S., A. Bais, V. Amiridis, D. Balis, C. Meleti, 

N. Kouremeti, C.S. Zerefos, S. Rapsomanikis, M. 

Petrakakis, A. Kelesis, P. Tzoumaka, and K. Kelektsoglou, Nine years of UV aerosol optical depth 

measurements at Thessaloniki, Greece, Atmos. 

Chem. Phys., 7, 2091-2101, doi: 10.5194/acp-7-

2091-2007, 2007.

Kazantzidis, A., A.F. Bais, D.S. Balis, E. Kosmidis, 

and C.S. Zerefos, Sensitivity of solar UV radiation to ozone and temperature profiles at Thessaloniki (40.5°N, 23°E), Greece, J. Atmos. Sol.-

Terr. Phys., 67 (14), 1321-1330, doi: 10.1016/j.

jastp.2005.05.003, 2005.

Kazantzidis, A., A.F. Bais, M.M. Zempila, S. Kazadzis, 

P.N. den Outer, T. Koskela, and H. Slaper, Calculations of the human vitamin D exposure from UV 

spectral measurements at three European stations, 

Photochem. Photobiol. Sci., 8, 45-51, doi: 10.1039/

b811216a, 2009.

Kazantzidis, A., K. Tourpali, and A.F. Bais, Variability 

of cloud-free ultraviolet dose rates on global scale 

due to modeled scenarios of future ozone recovery, Photochem. Photobiol., 86, 117-122, doi: 

10.1111/j.1751-1097.2009.00645.x, 2010.

Kimlin, M.G, W.J. Olds, and M.R. Moore, Location and 

vitamin D synthesis: Is the hypothesis validated 

by geophysical data?, J. Photochem. Photobiol. 

B: Biol., 86 (3), 234-239, doi: 10.1016/j.jphotobiol.2006.10.004, 2007.

Kinne, S., M. Schulz, C. Textor, S. Guibert, Y. Balkanski, 

S.E. Bauer, T. Berntsen, T.F. Berglen, O. Boucher, 

M. Chin, W. Collins, F. Dentener, T. Diehl, R. Easter, J. Feichter, D. Fillmore, S. Ghan, P. Ginoux, 

S. Gong, A. Grini, J. Hendricks, M. Herzog, L. 

Horowitz, I. Isaksen, T. Iversen, A. Kirkavag, S. 

Kloster, D. Koch, J.E. Kristjansson, M. Krol, A. 

Lauer, J.-F. Lamarque, G. Lesins, X. Liu, U. Lohmann, V. Montanaro, G. Myhre, J. Penner, G. Pitari, 
3.54

Chapter 3

S. Reddy, O. Seland, P. Stier, T. Takemura, and 

X. Tie, An AeroCom initial assessment – optical 

properties in aerosol component modules of global 

models, Atmos. Chem. Phys., 6, 1815-1834, doi: 

10.5194/acp-6-1815-2006, 2006.

Kinnison, D.E., K.E. Grant, P.S. Connell, D.A. Rotman, 

and D.J. Wuebbles, The chemical and radiative effects of the Mount Pinatubo eruption, J. Geophys. 

Res., 99 (D12), 25705-25731, 1994.

Kirk-Davidoff, D.B., E.J. Hintsa, J.G. Anderson, and D.W. 

Keith, The effect of climate change on ozone depletion through changes in stratospheric water vapour, 

Nature, 402, 399-401, doi: 10.1038/46521, 1999.

Konopka, P., J.-U. Grooß, K.W. Hoppel, H.-M. Steinhorst, and R. Müller, Mixing and chemical ozone 

loss during and after the Antarctic polar vortex major warming in September 2002, J. Atmos. Sci., 62 

(3), 848-859, 2005.

Lamarque, J.-F., D.E. Kinnison, P.G. Hess, and F.M. Vitt, 

Simulated lower stratospheric trends between 1970 

and 2005: Identifying the role of climate and composition changes, J. Geophys. Res., 113, D12301, 

doi: 10.1029/2007JD009277, 2008.

Lamarque, J.-F., T.C. Bond, V. Eyring, C. Granier, A. 

Heil, Z. Klimont, D. Lee, C. Liousse, A. Mieville, 

B. Owen, M.G. Schultz, D. Shindell, S.J. Smith, E. 

Stehfest, J. Van Aardenne, O.R. Cooper, M. Kainuma, N. Mahowald, J.R. McConnell, V. Naik, K. Riahi, and D.P. van Vuuren, Historical (1850–2000) 

gridded anthropogenic and biomass burning emissions of reactive gases and aerosols: Methodology 

and application, Atmos. Chem. Phys., 10, 7017-

7039, doi: 10.5194/acp-10-7017-2010, 2010.

Lapeta, B., O. Engelsen, Z. Litynska, B. Kois, and A. Kylling, 

Sensitivity of surface UV radiation and ozone column retrieval to ozone and temperature profiles, J. 

Geophys. Res., 105 (D4), 5001-5007, 2000.

Li, F., R.S. Stolarski, and P.A. Newman, Stratospheric 

ozone in the post-CFC era, Atmos. Chem. Phys., 9, 

2207-2213, doi: 10.5194/acp-9-2207-2009, 2009.

Lindfors, A., and A. Arola, On the wavelengthdependent attenuation of UV radiation by clouds, 

Geophys. Res. Lett., 35, L05806, doi: 10.1029/

2007GL032571, 2008.

Madronich, S., Analytic formula for the clear-sky UV 

index, Photochem. Photobiol., 83 (6), 1537-1538, 

doi: 10.1111/j.1751-1097.2007.00200.x, 2007.

Madronich, S., and S. Flocke: The role of solar radiation 

in atmospheric chemistry, in The Handbook of Environmental Chemistry, edited by P. Boule, 1-26, 

Springer-Verlag, Berlin Heidelberg, 1999.

Madronich S., R.L. McKenzie, L.O. Björn, and M.M. 

Caldwell, Changes in biologically active ultraviolet 

radiation reaching the Earth’s surface, Photochem. 

Photobiol. B: Biology, 46 (1-3), 5-19, doi: 10.1016/

S1011-1344(98)00182-1, 1998.

Matthews, H.D., and K. Caldeira, Transient climatecarbon simulations of planetary geoengineering, 

Proc. Nat. Acad. Sci., 104 (24), 9949-9954, doi: 

10.1073/pnas.0700419104, 2007.

Mayer, B., and A. Kylling, Technical note: The 

libRadtran software package for radiative transfer 

calculations – description and examples of use, 

Atmos. Chem. Phys., 5, 1855-1877, doi: 10.5194/

acp-5-1855-2005, 2005.

McKenzie, R., D. Smale, G. Bodeker, and H. Claude, 

Ozone profile differences between Europe and New 

Zealand: Effects on surface UV irradiance and its estimation from satellite sensors, J. Geophys. Res., 108 

(D6), 4179, doi: 10.1029/2002JD002770, 2003.

McKenzie, R.L., C. Weinreis, P.V. Johnston, B. Liley, H. 

Shiona, M. Kotkamp, D. Smale, N. Takegawa, and 

Y. Kondo, Effects of urban pollution on UV spectral irradiances, Atmos. Chem. Phys., 8, 5683-5697, 

doi: 10.5194/acp-8-5683-2008, 2008.

McKenzie, R.L., J.B. Liley, and L.O. Björn, UV radiation: Balancing risks and benefits, Photochem. 

Photobiol., 85 (1), 88-98, doi: 10.1111/j.1751-

1097.2008.00400.x, 2009.

McLandress, C., and T.G. Shepherd, Simulated anthropogenic changes in the Brewer-Dobson circulation, including its extension to high latitudes, J. Clim., 22 (6), 

1516-1540, doi: 10.1175/2008JCLI2679.1, 2009a.

McLandress, C., and T.G. Shepherd, Impact of climate 

change on stratospheric sudden warmings as simulated by the Canadian Middle Atmosphere Model, 

J. Clim., 22 (20), 5449-5463, 2009b.

McLandress, C., A.I. Jonsson, D.A. Plummer, M.C. 

Reader, J.F. Scinocca, and T.G. Shepherd, Separating the dynamical effects of climate change 

and ozone depletion: Part 1. Southern Hemisphere 

stratosphere, J. Clim., 23 (18), 5002-5020, doi: 

10.1175/2010JCLI3586.1, 2010.

Meehl, G.A., T.F. Stocker, W.D. Collins, P. Friedlingstein, A.T. Gaye, J.M. Gregory, A. Kitoh, R. 

Knutti, J.M. Murphy, A. Noda, S.C.B. Raper, I.G. 

Watterson, A.J. Weaver, and Z.-C. Zhao, Global 

climate projections, in Climate Change 2007: The 

Physical Science Basis. Contribution of Working 

Group I to the Fourth Assessment Report of the Intergovernmental Panel on Climate Change, edited 

by S. Solomon, D. Qin, M. Manning, Z. Chen, M. 

Marquis, K.B. Averyt, M. Tignor and H.L. Miller, 

996 pp., Cambridge University Press, Cambridge, 

U.K., and New York, NY, U.S.A., 2007.

Millard, G.A., T.A. Mather, D.M. Pyle, W.I. Rose, and 

B. Thornton, Halogen emissions from a small volcanic eruption: Modeling the peak concentrations, 
3.55

Future Ozone and UV

dispersion, and volcanically induced ozone loss in 

the stratosphere, Geophys. Res. Lett., 33, L19815, 

doi: 10.1029/2006GL026959, 2006.

Miller, A.J., R.M. Nagatani, L.E. Flynn, S. Kondragunta, 

E. Beach, R. Stolarski, R.D. McPeters, P.K. Bhartia, 

M.T. DeLand, C.H. Jackman, D.J. Wuebbles, K.O. 

Patten, and R.P. Cebula, A cohesive total ozone data 

set from SBUV(/2) satellite system, J. Geophys. Res.,

107(D23), 4701, doi: 10.1029/2001JD000853, 2002.

Morgenstern, O., P. Braesicke, M.M. Hurwitz, F.M. 

O’Connor, A.C. Bushell, C.E. Johnson, and J.A. 

Pyle, The world avoided by the Montreal Protocol, Geophys. Res. Lett., 35, L16811, doi: 

10.1029/2008GL034590, 2008.

Morgenstern, O., P. Braesicke, F.M. O’Connor, A.C. 

Bushell, C.E. Johnson, S.M. Osprey, and J.A. Pyle, 

Evaluation of the new UKCA climate-composition 

model - Part 1: The stratosphere, Geosci. Model 

Dev., 2, 43-57, doi: 10.5194/gmd-2-43-2009, 2009.

Morgenstern, O., M.A. Giorgetta, K. Shibata, V. Eyring, 

D.W. Waugh, T.G. Shepherd, H. Akiyoshi, J. Austin, A.J.G. Baumgaertner, S. Bekki, P. Braesicke, 

C. Brühl, M.P. Chipperfield, D. Cugnet, M. Dameris, S. Dhomse, S.M. Frith, H. Garny, A. Gettelman, S.C. Hardiman, M.I. Hegglin, P. Jöckel, D.E. 

Kinnison, J.-F. Lamarque, E. Mancini, E. Manzini, M. Marchand, M. Michou, T. Nakamura, J.E. 

Nielsen, D. Olivié, G. Pitari, D.A. Plummer, E. 

Rozanov, J.F. Scinocca, D. Smale, H. Teyssèdre, 

M. Toohey, W. Tian, and Y. Yamashita, Review of 

the formulation of present-generation stratospheric 

chemistry-climate models and associated external 

forcings, J. Geophys. Res., 115, D00M02, doi: 

10.1029/2009JD013728, 2010.

Moss, R., M. Babiker, S. Brinkman, E. Calvo, T. Carter, 

J. Edmonds, I. Elgizouli, S. Emori, L. Erda, K. 

Hibbard, R. Jones, M. Kainuma, J. Kelleher, J.-F. 

Lamarque, M. Manning, B. Matthews, J. Meehl, L. 

Meyer, J. Mitchell, N. Nakicenovic, B. O’Neill, R. 

Pichs, K. Riahi, S. Rose, P. Runci, R. Stouffer, D. 

van Vuuren, J. Weyant, T. Wilbanks, J.P. van Ypersele, and M. Zurek, Towards New Scenarios for 

Analysis of Emissions, Climate Change, Impacts, 

and Response Strategies, IPCC Expert Meeting Report, Intergovernmental Panel on Climate Change, 

Geneva, 132 pp., available: http://www.aimes.ucar.

edu/docs/IPCC.meetingreport.final.pdf, 2008.

Newman, P.A., J.S. Daniel, D.W. Waugh, and E.R. Nash, 

A new formulation of equivalent effective stratospheric chlorine (EESC), Atmos. Chem. Phys., 7, 

4537-4552, doi: 10.5194/acp-7-4537-2007, 2007.

Nissen, K.M., K. Matthes, U. Langematz, and B. Mayer, 

Towards a better representation of the solar cycle in 

general circulation models, Atmos. Chem. Phys., 7, 

5391-5400, doi: 10.5194/acp-7-5391-2007, 2007.

Oman, L., D.W. Waugh, S. Pawson, R.S. Stolarski, and 

J.E. Neilsen, Understanding the changes of stratospheric water vapor in coupled chemistry-climate 

model simulations, J. Atmos. Sci., 65 (10), 3278-

3291, doi: 10.1175/2008JAS2696, 2008.

Oman, L.D., D.W. Waugh, S.R. Kawa, R.S. Stolarski, 

A.R. Douglass, and P.A. Newman, Mechanisms 

and feedbacks causing changes in upper stratospheric ozone in the 21st century, J. Geophys. Res., 

115, D05303, doi: 10.1029/2009JD012397, 2010.

Overland, J.E., and M. Wang, Future regional Arctic sea 

ice declines, Geophys. Res. Lett., 34, L17705, doi: 

10.1029/2007GL030808, 2007.

Pawson, S., R.S. Stolarski, A.R. Douglass, P.A. Newman, 

J.E. Nielsen, S.M. Frith, and M.L. Gupta, Goddard 

Earth Observing System chemistry-climate model 

simulations of stratospheric ozone-temperature 

coupling between 1950 and 2005, J. Geophys. Res., 

113, D12103, doi: 10.1029/2007JD009511, 2008.

Pitari, G., E. Mancini, V. Rizi, and D.T. Shindell, Impact 

of future climate and emission changes on stratospheric aerosols and ozone, J. Atmos. Sci., 59 (3), 

414-440, 2002.

Portmann, R.W., and S. Solomon, Indirect radiative forcing of the ozone layer during the 21st 

century, Geophys. Res. Lett., 34, L02813, doi: 

10.1029/2006GL028252, 2007.

Portmann, R.W., S. Solomon, R.R. Garcia, L.W. Thomason, L.R. Poole, and M.P. McCormick, Role of 

aerosol variations in anthropogenic ozone depletion 

in the polar regions, J. Geophys. Res., 101 (D17), 

22991-23006, 1996.

Portmann, R.W., S.S. Brown, T. Gierczak, R.K. Talukdar, 

J.B. Burkholder, and A.R. Ravishankara, Role of 

nitrogen oxides in the stratosphere: A reevaluation 

based on laboratory studies, Geophys. Res. Lett., 26

(15), 2387-2390, 1999.

Randel, W.J., F. Wu, H. Vömel, G.E. Nedoluha, and P. 

Forster, Decreases in stratospheric water vapor after 

2001: Links to changes in the tropical tropopause and 

the Brewer-Dobson circulation, J. Geophys. Res., 

111, D12312, doi: 10.1029/2005JD006744, 2006.

Randel, W.J., K.P. Shine, J. Austin, J. Barnett, C. Claud, 

N.P. Gillett, P. Keckhut, U. Langematz, R. Lin, 

C. Long, C. Mears, A. Miller, J. Nash, D.J. 

Seidel, D.W.J. Thompson, F. Wu, and S. Yoden, 

An update of observed stratospheric temperature trends, J. Geophys. Res., 114, D02107, doi: 

10.1029/2008JD010421, 2009.

Randeniya, L.K., P.F. Vohralik, and I.C. Plumb, Stratospheric ozone depletion at northern mid latitudes 

in the 21st century: The importance of future concentrations of greenhouse gases nitrous oxide and 
3.56

Chapter 3

methane, Geophys. Res. Lett., 29 (4), 1051, doi: 

10.1029/2001GL014295, 2002.

Rasch, P.J., S. Tilmes, R.P. Turco, A. Robock, L. Oman, C.-

C. Chen, G.L. Stenchikov, and R.R. Garcia, An overview of geoengineering of climate using stratospheric 

sulphate aerosols, Phil. Trans. R. Soc. A, 366 (1882), 

4007-4037, doi: 10.1098/rsta.2008.0131, 2008.

Ravishankara, A.R., J.S. Daniel, and R.W. Portmann, Nitrous 

oxide (N2O): The dominant ozone-depleting substance emitted in the 21st century, Science, 326 (5949), 

123-125, doi: 10.1126/science.1176985, 2009.

Riahi, K., A. Grübler, and N. Nakicenovic, Scenarios of 

long-term socio-economic and environmental development under climate stabilization, Technological Forecasting and Social Change, 74 (7), 887-

935, doi: 10.1016/j.techfore.2006.05.026, 2007.

Rieder, H.E., F. Holawe, S. Simic, M. Blumthaler, J.W. 

Krzyścin, J.E. Wagner, A.W. Schmalwieser, and 

P. Weihs, Reconstruction of erythemal UV-doses 

for two stations in Austria: A comparison between 

alpine and urban regions, Atmos. Chem. Phys., 8, 

6309-6323, doi: 10.5194/acp-8-6309-2008, 2008.

Rind, D., J. Lerner, J. Jonas, and C. McLinden, Effects of 

resolution and model physics on tracer transports 

in the NASA Goddard Institute for Space Studies 

general circulation models, J. Geophys Res., 112, 

D09315, doi: 10.1029/2006JD007476, 2007.

Robock, A., A. Marquardt, B. Kravitz, and G. Stenchikov, 

Benefits, risks, and costs of stratospheric 

geoengineering, Geophys. Res. Lett., 36, L19703, 

doi: 10.1029/2009GL039209, 2009.

Rohs, S., C. Schiller, M. Riese, A. Engel, U. Schmidt, 

T. Wetter, I. Levin, T. Nakazawa, and S. Aoki, 

Long-term changes of methane and hydrogen 

in the stratosphere in the period 1978–2003 and 

their impact on the abundance of stratospheric 

water vapor, J. Geophys. Res., 111, D14315, doi: 

10.1029/2005JD006877, 2006.

Rosenfield, J.E., and A.R. Douglass, Doubled CO2 effects 

on NOy in a coupled 2D model, Geophys. Res. Lett.,

25 (23), 4381-4384, 1998.

Royal Society, Geoengineering the Climate: Science,

Governance and Uncertainty, Report 10/09, 98 

pp., London GB, Royal Society, available: http://

royalsociety.org/geoengineering-the-climate/, 2009.

Schraner, M., E. Rozanov, C. Schnadt Poberaj, P. Kenzelmann, A.M. Fischer, V. Zubov, B.P. Luo, C.R. 

Hoyle, T. Egorova, S. Fueglistaler, S. Brönnimann, W. Schmutz, and T. Peter, Technical Note: 

Chemistry-climate model SOCOL: Version 2.0 

with improved transport and chemistry/microphysics schemes, Atmos. Chem. Phys., 8, 5957-5974, 

doi: 10.5194/acp-8-5957-2008, 2008.

Schulz, M., C. Textor, S. Kinne, Y. Balkanski, S. Bauer, 

T. Berntsen, T. Berglen, O. Boucher, F. Dentener, 

S. Guibert, I.S.A. Isaksen, T. Iversen, D. Koch, A. 

Kirkevåg, X. Liu, V. Montanaro, G. Myhre, J.E. 

Penner, G. Pitari, S. Reddy, Ø. Seland, P. Stier, 

and T. Takemura, Radiative forcing by aerosols 

as derived from the AeroCom present-day and 

pre-industrial simulations, Atmos. Chem. Phys., 6, 

5225-5246, doi: 10.5194/acp-6-5225-2006, 2006.

Schwander, H., P. Koepke, and A. Ruggaber, Uncertainties in modeled UV irradiances due to limited accuracy and availability of input data, J. Geophys. Res.,

102 (D8), 9419-9429, doi: 10.1029/97JD00244, 

1997.

Scinocca, J.F., N.A. McFarlane, M. Lazare, J. Li, and D. 

Plummer, Technical Note: The CCCma third generation AGCM and its extension into the middle atmosphere, Atmos. Chem. Phys., 8, 7055-7074, doi: 

10.5194/acp-8-7055-2008, 2008.

Shepherd, T.G., Dynamics, stratospheric ozone, and climate change, Atmos.-Ocean, 46 (1), 117-138, doi: 

10.3137/ao.460106, 2008.

Shepherd, T.G., and A.I. Jonsson, On the attribution of 

stratospheric ozone and temperature changes to 

changes in ozone-depleting substances and wellmixed greenhouse gases, Atmos. Chem. Phys., 8, 

1435-1444, doi: 10.5194/acp-8-1435-2008, 2008.

Shindell, D.T., D. Rind, and P. Lonergan, Increased polar 

stratospheric ozone losses and delayed eventual 

recovery owing to increasing greenhouse-gas concentrations, Nature, 392, 589-592, doi: 10.1038/

33385, 1998.

Shibata, K., and M. Deushi, Long-term variations and trends 

in the simulation of the middle atmosphere 1980-

2004 by the chemistry-climate model of the Meteorological Research Institute, Ann. Geophys., 26, 1299-

1326, doi: 10.5194/angeo-26-1299-2008, 2008a.

Shibata, K., and M. Deushi, Simulations of the stratospheric circulation and ozone during the recent 

past (1980-2004) with the MRI chemistry-climate 

model, CGER’s Supercomputer Monograph Report 

Vol.13, Center for Global Environmental Research, 

National Institute for Environmental Studies, Japan, 154 pp., 2008b.

Sinnhuber, B.-M., and I. Folkins, Estimating the contribution of bromoform to stratospheric bromine and its 

relation to dehydration in the tropical tropopause 

layer, Atmos. Chem. Phys., 6, 4755-4761, doi: 

10.5194/acp-6-4755-2006, 2006.

Sinnhuber, B.-M., N. Sheode, M. Sinnhuber, M.P. Chipperfield, and W. Feng, The contribution of anthropogenic bromine emissions to past stratospheric 

ozone trends: A modeling study, Atmos. Chem. 

Phys., 9, 2863-2871, doi: 10.5194/acp-9-2863-

2009, 2009.
3.57

Future Ozone and UV

Solomon, S., R.W. Portmann, R.R. Garcia, L.W. Thomason, L.R. Poole, and M.P. McCormick, The role of 

aerosol variations in anthropogenic ozone depletion at northern midlatitudes, J. Geophys. Res., 101 

(D3), 6713-6727, doi: 10.1029/95JD03353, 1996.

Solomon, S., R.W. Portmann, T. Sasaki, D.J. Hofmann, 

and D.W.J. Thompson, Four decades of ozonesonde 

measurements over Antarctica, J. Geophys. Res., 

110, D21311, doi: 10.1029/2005JD005917, 2005.

Solomon, S., D. Qin, M. Manning, R.B. Alley, T. Berntsen, N.L. Bindoff, Z. Chen, A. Chidthaisong, J.M. 

Gregory, G.C. Hegerl, M. Heimann, B. Hewitson, 

B.J. Hoskins, F. Joos, J. Jouzel, V. Kattsov, U. 

Lohmann, T. Matsuno, M. Molina, N. Nicholls, J. 

Overpeck, G. Raga, V. Ramaswamy, J. Ren, M. 

Rusticucci, R. Somerville, T.F. Stocker, P. Whetton, R.A. Wood, and D. Wratt, Technical Summary. In: Climate Change 2007: The Physical Science Basis. Contribution of Working Group I to the 

Fourth Assessment Report of the Intergovernmental Panel on Climate Change, edited by S. Solomon, D. Qin, M. Manning, Z. Chen, M. Marquis, 

K.B. Averyt, M. Tignor and H.L. Miller, 996 pp., 

Cambridge University Press, Cambridge, U.K., and 

New York, NY, U.S.A., 2007.

Son, S.-W., N.F. Tandon, L.M. Polvani, and D.W. Waugh, 

Ozone hole and Southern Hemisphere climate 

change, Geophys. Res. Lett., 36, L15705, doi: 

10.1029/2009GL038671, 2009.

SPARC (Stratospheric Processes And their Role in Climate), Assessment of Stratospheric Aerosols 

Properties, edited by L. Thomason and Th. Peter, 

SPARC Report No. 4, WCRP-124, WMO/TD- No. 

1295, 2006.

SPARC CCMVal, SPARC CCMVal Report on the Evaluation of Chemistry-Climate Models, edited by V. Eyring, T.G. Shepherd, D.W. Waugh, SPARC Report 

No. 5, WCRP-132, WMO/TD-No. 1526, available: 

http://www.atmosp.physics.utoronto.ca/SPARC/

ccmval_final/index.php, 2010.

Staiger, H., P.N. den Outer, A.F. Bais, U. Feister, B. 

Johnsen, and L. Vuilleumier, Hourly resolved 

cloud modification factors in the ultraviolet, Atmos. 

Chem. Phys., 8, 2493-2508, doi: 10.5194/acp-8-

2493-2008, 2008.

Stenke, A., V. Grewe, and M. Ponater, Lagrangian 

transport of water vapor and cloud water in the 

ECHAM4 GCM and its impact on the cold bias, 

Clim. Dyn., 31 (5), 491-506, doi: 10.1007/s00382-

007-0347-5, 2008.

Stenke, A., M. Dameris, V. Grewe, and H. Garny, Implications of Lagrangian transport for simulations 

with a coupled chemistry-climate model, Atmos. 

Chem. Phys., 9, 5489-5504, doi: 10.5194/acp-9-

5489-2009, 2009.

Stolarski, R.S., and S. Frith, Search for evidence of trend 

slow-down in the long-term TOMS/SBUV total 

ozone data record: The importance of instrument 

drift uncertainty, Atmos. Chem. Phys., 6, 4057-

4065, doi: 10.5194/acp-6-4057-2006, 2006.

Strahan, S.E., and B.C. Polansky, Meteorological implementation issues in chemistry and transport models, 

Atmos. Chem. Phys., 6, 2895-2910, doi: 10.5194/

acp-6-2895-2006, 2006.

Stroeve, J., M.M. Holland, W. Meier, T. Scambos, and 

M. Serreze, Arctic sea ice decline: Faster than 

forecast, Geophys. Res. Lett., 34, L09501, doi: 

10.1029/2007GL029703, 2007.

Tanskanen, A., and T. Manninen, Effective UV surface 

albedo of seasonally snow-covered lands, Atmos. 

Chem. Phys., 7, 2759-2764, doi: 10.5194/acp-7-

2759-2007, 2007.

Taylor, K.E., R.J. Stouffer, and G.A. Meehl, A Summary 

of the CMIP5 Experiment Design, https://cmip.llnl.

gov/cmip5/docs/Taylor_CMIP5_design.pdf, 2009.

Tegen, I., M. Werner, S.P. Harrison, and K.E. Kohfeld, 

Relative importance of climate and land use in 

determining present and future global soil dust 

emission, Geophys. Res. Lett., 31, L05105, doi: 

10.1029/2003GL019216, 2004.

Telford, P., P. Braesicke, O. Morgenstern, and J. Pyle, Reassessment of causes of ozone column variability 

following the eruption of Mount Pinatubo using a 

nudged CCM, Atmos. Chem. Phys., 9, 4251-4260, 

doi: 10.5194/acp-9-4251-2009, 2009.

Teyssèdre, H., M. Michou, H.L. Clark, B. Josse, F. Karcher, D. Olivié, V.-H. Peuch, D. Saint-Martin, D. Cariolle, J.-L. Attié, P. Nédélec, P. Ricaud, V. Thouret, 

R.J. van der A, A. Volz-Thomas, and F. Chéroux, A 

new tropospheric and stratospheric Chemistry and 

Transport Model MOCAGE-Climat for multi-year 

studies: Evaluation of the present-day climatology 

and sensitivity to surface processes, Atmos. Chem. 

Phys., 7, 5815-5860, doi: 10.5194/acp-7-5815-

2007, 2007.

Tian, W., and M.P. Chipperfield, A new coupled chemistryclimate model for the stratosphere: The importance 

of coupling for future O3-climate predictions, Quart. 

J. Roy. Meteorol. Soc., 131 (605), 281-303, 2005.

Tian, W., M.P. Chipperfield, L.J. Gray, and J.M. Zawodny, Quasi-biennial oscillation and tracer distributions in a coupled chemistry-climate model, 

J. Geophys. Res., 111, D20301, doi: 10.1029/ 

2005JD006871, 2006.

Tian, W., M.P. Chipperfield, and D. Lu, Impact of increasing stratospheric water vapor on ozone depletion and temperature change, Adv. Atm. Sci., 26 (3), 

423-437, doi: 10.1007/s00376-009-0423-3, 2009.
3.58

Chapter 3

Tie, X., and G. Brasseur, The response of stratospheric 

ozone to volcanic eruptions: Sensitivity to atmospheric chlorine loading, Geophys. Res. Lett., 22 

(22), 3035-3038, doi: 10.1029/95GL03057, 1995.

Tilmes, S., R. Müller, and R. Salawitch, The sensitivity 

of polar ozone depletion to proposed geoengineering schemes, Science, 320 (5880), 1201-1204, doi: 

10.1126/science.1153966, 2008a.

Tilmes S., R. Müller, R.J. Salawitch, U. Schmidt, C.R. 

Webster, H. Oelhaf, C.C. Camy-Peyret, and J.M. 

Russell III, Chemical ozone loss in the Arctic winter 1991-1992, Atmos. Chem. Phys., 8, 1897-1910, 

doi: 10.5194/acp-8-1897-2008, 2008b.

Tilmes, S., R.R. Garcia, D.E. Kinnison, A. Gettelman, and 

P.J. Rasch, Impact of geoengineered aerosols on 

the troposphere and stratosphere, J. Geophys. Res.,

114, D12305, doi: 10.1029/2008JD011420, 2009.

Tourpali K., A.F. Bais, A. Kazantzidis, C.S. Zerefos, H. 

Akiyoshi, J. Austin, C. Brühl, N. Butchart, M.P. 

Chipperfield, M. Dameris, M. Deushi, V. Eyring, 

M.A. Giorgetta, D.E. Kinnison, E. Mancini, D.R. 

Marsh, T. Nagashima, G. Pitari, D.A. Plummer, E. 

Rozanov, K. Shibata, and W. Tian, Clear sky UV 

simulations for the 21st century based on ozone and 

temperature projections from Chemistry-Climate 

Models, Atmos. Chem. Phys., 9, 1165-1172, doi: 

10.5194/acp-9-1165-2009, 2009.

Trenberth, K.E., and J.T. Fasullo, Global warming due to increasing absorbed solar radiation, 

Geophys. Res. Lett., 36, L07706, doi: 10.1029/ 

2009GL037527, 2009.

Tsay, S.-C., and K. Stamnes, Ultraviolet radiation in the 

Arctic: The impact of potential ozone depletions 

and cloud effects, J. Geophys. Res., 97 (D8), 7829-

7840, 1992.

UNEP (United Nations Environment Programme), Environmental Effects of Ozone Depletion and Its Interaction with Climate Change: 2006 Assessment, 

Nairobi, 206 pp., 2007.

UNEP (United Nations Environment Programme), Environmental Effects of Ozone Depletion and Its Interactions with Climate Change: 2010 Assessment, 

Nairobi, available: http://www.unep.ch/ozone/

Assessment_Panels/EEAP/index.shtml, 2010.

van Vuuren, D.P., M.G.J. den Elzen, P.L. Lucas, B. Eickhout, B.J. Strengers, B. van Ruijven, S. Wonink, 

and R. van Houdt, Stabilizing greenhouse gas concentrations at low levels: An assessment of reduction strategies and costs, Clim. Change, 81, 119-

159, doi: 10.1007/s/10584-006-9172-9, 2007.

Waibel, A.E., Th. Peter, K.S. Carslaw, H. Oelhaf, G. Wetzel, P.J. Crutzen, U. Pöschl, A. Tsias, E. Reimer, 

and H. Fischer, Arctic ozone loss due to dentrification, Science, 283 (5410), 2064-2069, doi: 10.1126/

science.283.5410.2064, 1999.

Wall, D.H., Global change tipping points: Above- and 

below-ground biotic interactions in a low diversity 

ecosystem, Phil. Trans. Roy. Soc. B, 362 (1488), 

2291-2306, doi: 10.1098/rstb.2006.1950, 2007.

Waugh, D.W., and V. Eyring, Quantitative performance 

metrics for stratospheric-resolving chemistry-climate models, Atmos. Chem. Phys., 8, 5699-5713, 

doi: 10.5194/acp-8-5699-2008, 2008.

Waugh, D.W., L. Oman, S.R. Kawa, R.S. Stolarski, S. 

Pawson, A.R. Douglass, P.A. Newman, and J.E. 

Nielsen, Impacts of climate change on stratospheric 

ozone recovery, Geophys. Res. Lett., 36, L03805, 

doi: 10.1029/2008GL036223, 2009.

Webb, A.R., and O. Engelsen, Calculated ultraviolet exposure levels for a healthy vitamin D status, Photochem. Photobiol., 82, 1697-1703, 2006.

Webb, A.R., L. Kline, and M.F. Holick, Influence of season and latitude on the cutaneous synthesis of vitamin D3: Exposure to winter sunlight in Boston and 

Edmonton will not promote vitamin D3 synthesis in 

human skin, J. Clinic. Endocrin. Metabol., 67 (2), 

373-378, doi: 10.1210/jcem-67-2-373, 1988.

Wigley, T.M.L., A combined mitigation/geoengineering approach to climate stabilization, Science, 314 (5798), 

452- 454, doi: 10.1126/science.1131728, 2006.

Wohltmann, I., R. Lehmann, M. Rex, D. Brunner, and 

J.A. Mäder, A process-oriented regression model 

for column ozone, J. Geophys. Res., 112, D12304,

1-18, doi: 10.1029/2006JD007573, 2007.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 1998, Global 

Ozone Research and Monitoring Project–Report 

No. 44, Geneva, Switzerland, 1999.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 2002, Global 

Ozone Research and Monitoring Project–Report 

No. 47, 498 pp., Geneva, Switzerland, 2003.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 2006, Global 

Ozone Research Monitoring Project–Report No. 

50, 572 pp., Geneva, Switzerland, 2007.

Zhang, X., and J.E. Walsh, Toward a seasonally ice-covered Arctic Ocean: Scenarios from the IPCC AR4 

model simulations, J. Clim., 19 (9), 1730-1747, doi: 

10.1175/JCLI3767.1, 2006.
3.59

Future Ozone and UV

APPENDIX 3A

Constructing Correlative Time Series Plots

The correlative time series plots shown in Figures 3-8 and 3-10 were constructed to illustrate how the rise and fall of 

ESC correlates with stratospheric ozone changes over multi-decadal time periods in scenarios that include and exclude the 

effects of GHGs and ODSs. This appendix describes how these plots are constructed and evaluated.

Calculated global or regional stratospheric ozone values typically are shown as time series and qualified by the corresponding time series of temperature and ESC because of the strong influence these parameters have in controlling ozone 

amounts. As an example, Figure 3A-1(a) shows multi-model area weighted ozone mixing ratios from 25°S to 25°N at 5 

hPa from 1960 to 2100 for the reference (black), ODSs fixed at 1960 values (blue), and GHGs fixed at 1960 values (green) 

simulations (Table 3-2). The combination of the fixed ODS and fixed GHG simulations (detailed below), used to test for 

linear additivity, is shown in yellow. The CCSRNIES, MRI, and WACCM chemistry-climate models were used to form 

the multi-model trend and all time series were subjected to the TSAM smoothing described in Section 3.3.2.2. The results 

for temperature and ESC in the same model runs are shown in Figure 3A-1(b) and (c).

To test the linear additivity of the ozone responses to changes in ODSs and GHGs, the sum of the individual ozone 

responses in the fixed ODS and fixed GHG simulations was evaluated as a function of time (t) as

Ozonesum(t) = fGHGozone(t) + fODSozone(t) − fODSozone(1960)

where fODSozone(1960) is the ozone 

value in 1960 in all simulations (dashed 

line in Figure 3A-1(a)). This Ozonesum

time series is shown as a yellow trace in 

Figure 3A-1(a) and (d). As for ozone, 

the linear additivity of the effects of 

changes in ODSs and GHGs on ESC is 

tested by calculating 

ESCsum(t) = fGHGESC(t) + fODSESC(t) 

− fODSESC(1960)

where the ESC terms correspond to 

those defined above for ozone and 

fODSESC(1960) is the ESC value in 

1960 in all simulations. The ESCsum

time series is shown as a yellow trace in 

Figure 3A-1(c) and (d). In the interests 

of clarity, the linear additivity of the effects of changes in ODSs and GHGs on 

temperature were not investigated and 

are therefore not shown in panel (b). 

A correlative time series plot corresponding to the ozone, temperature, 

Figure 3A-1. Time series for (a) 

ozone (ppm), (b) temperature (K), 

and (c) ESC (ppb) from CCM simulation results for 5 hPa between 

25°S and 25°N (area-weighted) 

and corresponding correlative time 

series plot (d). The traces in (d) are 

color-coded by temperature, and 

decade year markers are shown 

with symbols and labels.
3.60

Chapter 3

and ESC time series is shown in Figure 3A-1(d) with a separate line for each of the scenarios discussed above. Ozone is 

plotted on the ordinate and ESC on the abscissa because ESC is the expected main driver in the long-term ozone evolution. 

Symbols are used to mark decade years along each time series line and temperature is indicated with a color scale. A correlative time-series plot is of value because it highlights in a single panel the sensitivity and response of ozone to ESC and other 

variables in the CCMs over a multi-decadal period in a specific region or globally.

There are several important features of this plot type. For example, in this case, the fODS simulation appears in 

Figure 3A-1(d) as a nearly vertical line, the length of which indicates the overall effect of non-ODS factors on ozone. The 

fODS line may also deviate from the vertical indicating that non-ODS factors affect the conversion of ODSs to ESC. If 

the non-ODS factors had no influence on ozone and ESC, the fODS result would appear as a point on the plot. The temperature gradient along the line indicates a systematic change in this controlling factor. Apart from the fODS simulation, 

all simulations trace a broad path along the ESC axis from very low concentrations of ESC in 1960, through the peak in 

ESC around 2000 on the far right of the plot, and back to lower ESC concentrations by the end of the 21st century. In these 

cases, if ozone was controlled solely by ESC, the increasing and decreasing legs of this time series would trace the same 

path. Any deviation between the increasing and decreasing legs is indicative of a systematic influence of non-ODS factors. 

Finally, the separation at each point between the reference and linear additivity test traces indicates the degree to which the 

influence of ODS and GHG changes on ozone do not act independently, assuming that the uncertainties in these projections are small compared to this separation (McLandress et al., 2010). Thus, the time series displays and the correlative 

time series plot offer complementary views and insights into the evaluation of the CCM results.
Chapter 4

Stratospheric Changes and Climate

Coordinating Lead Authors:

P.M. Forster

D.W.J. Thompson

Lead Authors:

M.P. Baldwin

M.P. Chipperfield

M. Dameris

J.D. Haigh

D.J. Karoly

P.J. Kushner

W.J. Randel

K.H. Rosenlof

D.J. Seidel

S. Solomon

Coauthors:

G. Beig

P. Braesicke

N. Butchart

N.P. Gillett

K.M. Grise

D.R. Marsh

C. McLandress

T.N. Rao

S.-W. Son

G.L. Stenchikov

S. Yoden

Contributors:

E.C. Cordero

M.P. Free

A.I. Jonsson

J. Logan

D. Stevenson

Chapter 4

STRATOSPHERIC CHANGES AND CLIMATE

Contents

SCIENTIFIC SUMMARY .............................................................................................................................................1

4.0 INTRODUCTION AND SCOPE............................................................................................................................3

4.1 OBSERVED VARIATIONS IN STRATOSPHERIC CONSTITUENTS THAT RELATE TO CLIMATE.........4

4.1.1 Long-Lived Greenhouse Gases and Ozone-Depleting Substances...........................................................4

4.1.2 Ozone.........................................................................................................................................................4

4.1.3 Stratospheric Water Vapor........................................................................................................................4

Box 4-1. How Do Stratospheric Composition Changes Affect Stratospheric Climate?..........................5

4.1.4 Stratospheric Aerosols...............................................................................................................................8

4.2 OBSERVED VARIATIONS IN STRATOSPHERIC CLIMATE........................................................................10

4.2.1 Observations of Long-Term Changes in Stratospheric Temperature......................................................10

4.2.2 Observations of Long-Term Changes in the Stratospheric Circulation ..................................................14

4.2.2.1 Stratospheric Zonal Flow..........................................................................................................14

4.2.2.2 Brewer-Dobson Circulation......................................................................................................14

4.3 SIMULATIONS OF STRATOSPHERIC CLIMATE CHANGE.........................................................................16

4.3.1 Simulation of Stratospheric Temperature Trends from Chemistry-Climate

Models and Climate Models....................................................................................................................16

4.3.2 Simulation of Brewer-Dobson Circulation Trends in Chemistry-Climate Models.................................22

4.4 EFFECTS OF VARIATIONS IN STRATOSPHERIC CLIMATE ON THE 

TROPOSPHERE AND SURFACE.......................................................................................................................25

4.4.1 Effects of Stratospheric Composition Changes on Global-Mean Surface 

Temperature and Tropospheric Temperature..........................................................................................25

4.4.2 Surface Climate Impacts of the Antarctic Ozone Hole ...........................................................................28

4.4.2.1 Effects on Winds, Storm Tracks, and Precipitation..................................................................30

4.4.2.2 Effects on Surface Temperatures and Sea Ice...........................................................................31

4.4.2.3 Effects on Southern Ocean Temperatures and Circulation.......................................................32

4.4.2.4 Stratospheric Links to Southern Ocean Carbon........................................................................35

4.4.3 Stratospheric Variations and the Width of the Tropical Belt..................................................................37

4.4.4 Effects of Stratospheric Variations on Tropospheric Chemistry.............................................................38

4.4.5 Influence of the Stratosphere on the Impact of Solar Variability on Surface Climate............................39

4.5 WHAT TO EXPECT IN THE FUTURE ..............................................................................................................41

4.5.1 Stratospheric Temperatures.....................................................................................................................41

4.5.2 Brewer-Dobson Circulation ....................................................................................................................42

4.5.3 Stratospheric Water Vapor......................................................................................................................42

4.5.4 Tropopause Height and Width of the Tropical Belt................................................................................43

4.5.5 Radiative Effects and Surface Temperature............................................................................................43

4.5.6 Tropospheric Annular Modes and Stratosphere-Troposphere Coupling ................................................43

4.5.7 Tropospheric Chemistry..........................................................................................................................45

4.5.8 Solar and Volcanic Influences.................................................................................................................45

REFERENCES .............................................................................................................................................................46

4.1

Stratospheric Changes and Climate

SCIENTIFIC SUMMARY

•	 Stratospheric climate trends since 1980 are better understood and characterized than in previous Assessments

and continue to show the clear influence of both human and natural factors.

•		 New analyses of both satellite and radiosonde data give increased confidence relative to previous

Assessments of the complex time/space evolution of stratospheric temperatures between 1980 and 2009.

The global-mean lower stratosphere cooled by 1–2 K and the upper stratosphere cooled by 4–6 K from 1980 to 

about 1995. There have been no significant long-term trends in global-mean lower-stratospheric temperatures 

since about 1995. The global-mean lower-stratospheric cooling did not occur linearly but was manifested as 

downward steps in temperature in the early 1980s and the early 1990s. The cooling of the lower stratosphere 

included the tropics and was not limited to extratropical regions as previously thought.

•		 The complex evolution of lower-stratospheric temperature is influenced by a combination of natural and

human factors that has varied over time. Ozone decreases dominate the lower-stratospheric cooling over the 

long term (since 1980). Major volcanic eruptions and solar activity have clear shorter-term effects. Since the 

mid-1990s, slowing ozone loss has contributed to the lack of temperature trend. Models that consider all of these 

factors are able to reproduce this complex temperature time history.

•		 The largest lower-stratospheric cooling continues to be found in the Antarctic ozone hole region during

austral spring and early summer. The cooling due to the ozone hole strengthened the Southern Hemisphere 

polar stratospheric vortex compared with the pre-ozone hole period during these seasons.

•		 Tropical lower-stratospheric water vapor amounts decreased by roughly 0.5 parts per million by volume

(ppmv) around 2000 and remained low through 2009. This followed an apparent but uncertain increase in 

stratospheric water vapor amounts from 1980–2000. The mechanisms driving long-term changes in stratospheric 

water vapor are not well understood.

•		 Stratospheric aerosol concentrations increased by between 4 to 7% per year, depending on location, from

the late 1990s to 2009. The reasons for the increases in aerosol are not yet clear, but small volcanic eruptions 

and increased coal burning are possible contributing factors.

•	 There is new and stronger evidence for radiative and dynamical linkages between stratospheric change and

specific changes in surface climate.

•		 Changes in stratospheric ozone, water vapor, and aerosols all radiatively affect surface temperature. The 

radiative forcing of climate in 2008 due to stratospheric ozone depletion (−0.05 ± 0.1 Watts per square meter 

(W/m2)) is much smaller than the positive radiative forcing due to the chlorofluorocarbons (CFCs) and hydrochlorofluorocarbons (HCFCs) largely responsible for that depletion (+0.31 ± 0.03 W/m2

). Radiative calculations 

and climate modeling studies suggest that the radiative effects of variability in stratospheric water vapor (roughly 

±0.1 W/m2 per decade) can contribute to decadal variability in globally averaged surface temperature. Climate 

models and observations show that the negative radiative forcing from a major volcanic eruption such as Mt. 

Pinatubo in 1991 (roughly −3 W/m2) can lead to a surface cooling that persists for about two years.

•		 Observations and model simulations show that the Antarctic ozone hole caused much of the observed

southward shift of the Southern Hemisphere middle latitude jet in the troposphere during summer since

1980. The horizontal structure, seasonality, and amplitude of the observed trends in the Southern Hemisphere 

tropospheric jet are only reproducible in climate models forced with Antarctic ozone depletion. The southward 

shift in the tropospheric jet extends to the surface of the Earth and is linked dynamically to the ozone holeinduced strengthening of the Southern Hemisphere stratospheric polar vortex.

•		 The southward shift of the Southern Hemisphere tropospheric jet due to the ozone hole has been linked

to a range of observed climate trends over Southern Hemisphere mid and high latitudes during summer.
4.2

Chapter 4

Because of this shift, the ozone hole has contributed to robust summertime trends in surface winds, warming over 

the Antarctic Peninsula, and cooling over the high plateau. Other impacts of the ozone hole on surface climate 

have been investigated but have yet to be fully quantified. These include observed increases in sea ice area averaged around Antarctica; a southward shift of the Southern Hemisphere storm track and associated precipitation; 

warming of the subsurface Southern Ocean at depths up to several hundred meters; and decreases of carbon 

uptake over the Southern Ocean.

•		 In the Northern Hemisphere, robust linkages between Arctic stratospheric ozone depletion and the tropospheric and surface circulation have not been established, consistent with the comparatively small ozone

losses there.

•	 The influence of stratospheric changes on climate will continue during and after stratospheric ozone recovery.

•		 The global middle and upper stratosphere are expected to cool in the coming century, mainly due to

carbon dioxide (CO2) increases. The cooling due to CO2 will cause ozone levels to increase in the middle 

and upper stratosphere, which will slightly reduce the cooling. Stratospheric ozone recovery will also reduce 

the cooling. These ozone changes will contribute a positive radiative forcing of climate (roughly +0.1 W/m2) 

compared to 2009 levels, adding slightly to the positive forcing from continued increases in atmospheric CO2

abundances. Future hydrofluorocarbon (HFC) abundances in the atmosphere are expected to warm the tropical 

lower stratosphere and tropopause region by roughly 0.3 K per part per billion (ppb) and provide a positive radiative forcing of climate.

•		 Chemistry-climate models predict increases of stratospheric water vapor, but confidence in these predictions is low. Confidence is low since these same models (1) have a poor representation of the seasonal cycle 

in tropical tropopause temperatures (which control global stratospheric water vapor abundances) and (2) cannot 

reproduce past changes in stratospheric water vapor abundances.

•		 Future recovery of the Antarctic ozone hole and increases in greenhouse gases are expected to have opposite effects on the Southern Hemisphere tropospheric middle latitude jet. Over the next 50 years, the recovery of the ozone hole is expected to reverse the recent southward shift of the Southern Hemisphere tropospheric 

jet during summer. However, future increases in greenhouse gases are expected to drive a southward shift in the 

Southern Hemisphere tropospheric jet during all seasons. The net effect of these two forcings on the jet during 

summer is uncertain.

•		 Climate simulations forced with increasing greenhouse gases suggest a future acceleration of the stratospheric Brewer-Dobson circulation. Such an acceleration would lead to decreases in column ozone in the 

tropics and increases in column ozone elsewhere by redistributing ozone within the stratosphere. The causal 

linkages between increasing greenhouse gases and the acceleration of the Brewer-Dobson circulation remain 

unclear.

•		 Future stratospheric climate change will affect tropospheric ozone abundances. In chemistry-climate models, the projected acceleration of the Brewer-Dobson circulation and ozone recovery act together to increase the 

transport of stratospheric ozone into the troposphere. Stratospheric ozone redistribution will also affect tropospheric ozone by changing the penetration of ultraviolet radiation into the troposphere, thus affecting photolysis 

rates.
4.3

Stratospheric Changes and Climate

4.0 INTRODUCTION AND SCOPE

Climate is changing at all levels in the atmosphere. 

This chapter considers changes in the stratosphere and 

the related aspects of troposphere and surface climate. 

While covering some of the same aspects presented in 

Chapter 5 (Baldwin and Dameris et al., 2007) of the 

previous Ozone Assessment (WMO, 2007), the current 

chapter is broader in scope and also addresses aspects 

of climate change beyond those associated with stratospheric ozone.

It is evident that the 1987 Montreal Protocol and its 

subsequent Amendments and Adjustments have led to reduced emissions of ozone-destroying halocarbons, many of 

which are greenhouse gases. The current chapter helps to 

place the Protocol’s climate impact within a wider context 

by critically assessing the effect of stratospheric climate 

changes on the troposphere and surface climate, following a formal request for this information by the Parties to 

the Montreal Protocol. As requested, the current chapter 

also considers the effects on stratospheric climate of some 

emissions that are not addressed by the Montreal Protocol, 

but are included in the 1997 Kyoto Protocol. Hence, the 

chapter covers some of the issues assessed in past Intergovernmental Panel on Climate Change (IPCC) reports (IPCC, 

2007; IPCC/TEAP, 2005). The current chapter is designed 

to provide useful input to future IPCC assessments.

The troposphere and surface climate are affected by 

many types of stratospheric change. Ozone plays a key 

role in such stratospheric climate change, but other physical factors play important roles as well. For this reason, 

we consider here the effects on the stratosphere of not 

only emissions of ozone-depleting substances (ODSs), but 

also of emissions of greenhouse gases, natural phenomena 

(e.g., solar variability and volcanic eruptions), and chemical, radiative, and dynamical stratosphere/troposphere 

coupling (Figure 4-1).

First, the chapter combines information about past 

trace gas emissions (from Chapter 1) and past ozone concentrations (from Chapter 2) with a new assessment of 

other relevant emissions (Section 4.1). It then draws on the 

assessed changes in emissions (Section 4.1) and observed 

stratospheric change (Section 4.2) to assess the nature and 

drivers of stratospheric climate change (Section 4.3). The 

chapter subsequently assesses the physical linkages between stratospheric climate change and climate change at 

Earth’s surface (Section 4.4). The chapter closes with a 

discussion of future stratospheric climate change and its 

influence on the troposphere and surface climate (Section 4.5), which links to the Chapter 3 discussion of future 

Land Ice

Sea Ice

Ozone 

Hole

Carbon Fluxes

Solar Variability

Stratospheric Composition 

and 

Temperature Changes

Tropospheric 

Composition and 

Temperature 

Changes

Explosive Volcano

Radiative, Dynamical,

and Chemical Coupling

Tropospheric Temperature Changes

Surface 

Temperatures

Ocean Circulation 

and Temperatures

Winds

Human Emissions of 

Ozone-Depleting 

Substances

Figure 4-1. Schematic of the drivers and mechanisms considered in this chapter.
4.4

Chapter 4

ozone trends. Section 4.5 also provides input into discussions of future scenarios in Chapter 5.

4.1 OBSERVED VARIATIONS IN 

STRATOSPHERIC CONSTITUENTS 

THAT RELATE TO CLIMATE

In this section we assess our current understanding 

of stratospheric composition changes. The mechanisms 

whereby such composition changes affect climate are 

reviewed briefly in Box 4-1. 

4.1.1 Long-Lived Greenhouse Gases and 

Ozone-Depleting Substances

ODSs, carbon dioxide (CO2), nitrous oxide (N2O), 

and methane (CH4) are all gases of tropospheric origin that 

impact climate and stratospheric ozone amounts. ODSs 

and N2O directly impact ozone chemistry. Changes in atmospheric concentrations of CH4 will lead to changes in 

stratospheric water vapor that in turn impact ozone chemistry and climate. Such gases also affect ozone indirectly 

via their effects on climate. Recent measurements and 

growth rates for these gases are covered in Chapter 1 of 

this report, with a summary of recent growth rates shown 

in Table 1-1 for the ODSs and Table 1-15 for other greenhouse gases. Atmospheric concentrations of CO2, N2O, 

CH4, and ODS replacements are projected to increase in 

the future, as discussed in more detail in Section 4.5.

4.1.2 Ozone

Variations and trends in stratospheric ozone influence climate via direct radiative effects and the resulting 

temperature and circulation changes. Past changes in 

stratospheric ozone are reviewed in Chapter 2, and the key 

points are summarized here as they relate to climate.

Column ozone for the recent past (2006–2008) is 

approximately 2.5% and 3.5% lower than pre-1980 values 

for 60°N–60°S and the globe, respectively. Time series 

of global ozone anomalies show a relative minimum 

during the middle 1990s, followed by an increase and 

relatively constant values since 1999 (Chapter 2). Locally, 

the largest losses have occurred in the Antarctic ozone 

hole, which is associated with a near-total loss of ozone 

in the lower stratosphere (~15–22 km) during Southern 

Hemisphere spring. The Antarctic ozone hole has led to 

large changes in temperature and circulation in the Southern Hemisphere polar stratosphere, as assessed in WMO 

(2007) and Section 4.2 of this chapter. The ozone hole has 

also led to changes at the Southern Hemisphere surface, 

as assessed here in Section 4.4.2. In contrast to the Antarctic, the Arctic is marked by smaller long-term trends 

and larger year-to-year variance in ozone during winter 

and spring (the variance is linked to meteorological variability). Observed changes in profile ozone (see Section 

2.1.4.2 of Chapter 2) show relatively large percentage decreases across the globe in the upper stratosphere (~35–47 

km), with net changes of over 10% between 1980 and 

2009. For the same period, relatively small decreases of 

ozone concentrations have been observed for the altitude 

range ~24–32 km.

Significant long-term changes in ozone concentrations are found in the lower stratosphere (below 24 km), 

although the observational record in this region of the 

stratosphere is more uncertain due to the dearth of highvertical resolution ozone measurements, including a lack 

of global long-term sampling from ozonesondes and continuous observations from high-vertical resolution satellite 

instruments. Despite this uncertainty, high-vertical resolution Stratospheric Aerosol and Gas Experiment (SAGE)-

based ozone trends will be used in the climate simulations 

run for the IPCC Fifth Assessment Report. The longterm global observations from the SAGE I and II satellite 

data (covering 1979–2005) show net lower-stratospheric 

ozone concentration decreases of ~5–10% near 20 km, 

with the largest percentage decreases in the tropics (over 

30°N-30°S) (Randel and Wu, 2007; see also Figure 2-27 

in Chapter 2 of this Assessment). The decreases in lower 

tropical stratospheric ozone are uncertain (see discussion 

in Chapter 2), but are consistent with decreases in tropical stratospheric temperatures (Thompson and Solomon, 

2005; Randel et al., 2009), and if robust provide possible observational evidence of increased upwelling in the 

lower tropical stratosphere (see Sections 4.2.2 and 4.3.2). 

Updated estimates of lower-stratospheric variations and 

trends are a topic of current research.

Changes in the amount of ozone since 1980 have 

caused a cooling of the lower and upper stratosphere (Section 4.3) and have likely contributed a negative radiative 

forcing of the surface climate (Section 4.4). They have 

also affected stratospheric circulation (Section 4.3) and 

caused significant changes to the surface-troposphere climate of the Southern Hemisphere (Section 4.4). Stratospheric ozone concentrations will continue to change in 

response to changes in ODSs and chemical feedbacks associated with stratospheric temperatures and composition, 

and these changes will continue to affect the climate of the 

stratosphere, troposphere, and surface (Section 4.5).

4.1.3 Stratospheric Water Vapor

Water vapor is the principal greenhouse gas and 

plays a key role in tropospheric and stratospheric chemistry. Throughout the atmosphere, water vapor plays both a 

radiative and chemical role. An increase in stratospheric 
4.5

Stratospheric Changes and Climate

Box 4-1. How Do Stratospheric Composition Changes Affect Stratospheric Climate?

The vertical temperature structure of Earth’s stratosphere is primarily driven by radiative processes (Figure 1). Solar radiation is principally absorbed by ozone, stratospheric aerosols, and molecular oxygen, and this absorption warms the stratosphere. 

Outgoing longwave (thermal infrared) radiation from the surface/troposphere is absorbed and re-emitted by greenhouse gases. 

Prime greenhouse gases in the atmosphere are water vapor (H2O), ozone (O3), carbon dioxide (CO2), methane (CH4), and nitrous 

oxide (N2O). Additionally, halocarbons (i.e., chlorofluorocarbons (CFCs), hydrochlorofluorocarbons (HCFCs), and brominated 

chlorofluorocarbons (halons)) not only affect atmospheric chemistry, but are also significant greenhouse gases. Aerosols can 

also emit in the longwave range.

Whether a change of greenhouse gas concentration warms or cools the stratosphere depends on the strength of the absorption bands of the gas and the opacity of the troposphere at the wavelength of absorption (e.g., Clough and Iacono, 1995). In high 

opacity regions of the spectrum, where powerful greenhouse gases such as CO2 and H2O absorb, there is little transmission from 

the surface to the stratosphere. These gases in the stratosphere therefore receive upwelling radiation from the generally cooler 

tropopause region below and emit longwave radiation at the stratosphere’s generally higher temperature. As this layer in the 

stratosphere emits more than it absorbed, the stratosphere cools if greenhouse gas concentrations are increasing. This is in contrast to their warming effect on the troposphere-surface region. CFCs, hydrofluorocarbons (HFCs), and other gases that absorb 

weakly in the “window” regions of the spectrum warm both the lower stratosphere and troposphere (e.g., Forster and Joshi, 2005) 

as they receive upwelling radiation from near the generally warmer surface and emit their radiation at a lower temperature. For 

strong absorbers the cooling effect in the stratosphere increases with altitude as these gases cool to space, maximizing at the 

stratopause near 50 km altitude where temperature is highest. For weaker absorbers their warming effect is maximized at the 

tropopause region, where the temperature contrast with the surface is largest.

The solar forcing of the Earth’s atmosphere is modulated by the 11-year activity cycle of the sun that is reflected in fluctuations in the intensity of solar radiation at different wavelengths. 11-year solar UV irradiance variations have a direct impact on the 

radiation and ozone budget of the stratosphere (e.g., Haigh, 1994). During years with high solar activity the solar UV irradiance 

is clearly enhanced, leading to additional ozone production and heating in the stratosphere and above (e.g., Lee and Smith, 2003).

The feedbacks operating between temperature and ozone are determined not only by radiative processes but also by chemical processes (Figure 1). Stratospheric composition is intimately related to the absorption of incoming solar (shortwave) radiation. 

Solar ultraviolet (UV) radiation is involved in both the 

creation and destruction of ozone, resulting in a maximum ozone concentration near 25 km. The dominant 

ozone loss cycles in the middle and upper stratosphere 

(via the catalytic cycles of nitrogen oxides (NOx), chlorine radicals (ClOx), and odd hydrogen (HOx)) slow 

with decreasing temperatures (e.g., Haigh and Pyle, 

1982), leading to higher ozone, concentrations. The 

situation is even more complicated in the polar lower 

stratosphere in late winter and spring. In addition to 

the gas-phase ozone loss cycles, as described above, 

there is an offset by chlorine- and bromine-containing 

reservoir species (Zeng and Pyle, 2003). These chemical substances are activated via heterogeneous processes on surfaces of polar stratospheric clouds. The 

rate of chlorine and bromine activation that determines 

the rate of ozone depletion is strongly dependent on 

stratospheric temperatures, increasing significantly 

below approximately 195 K due to enhanced particle 

formation. The amount of stratospheric ozone is also 

affected by heterogeneous chemical reactions acting 

on the surfaces of stratospheric aerosol particles. The 

injection of sulfate aerosols into the stratosphere leads 

to transformation of inactive chlorine compounds to 

active forms that destroy ozone.

Box 4-1, Figure 1. Schematic of ozone-temperature feedbacks due 

to changes in stratospheric chemical composition. Stratospheric temperature is determined directly by concentrations of radiatively active 

gases (e.g., long-lived greenhouse gases, LLGHGs) and aerosols (both 

emitted by explosive volcanic eruptions and human activities) via the 

absorption and emission of short- and longwave radiation. Moreover, 

the amount of stratospheric ozone is defined by transport via winds 

(i.e., dynamics) and chemical processes, which on its own part depends 

on concentrations of other greenhouse gases and aerosols. The picture is even more complex since stratospheric temperature influences 

net ozone production due to temperature-dependent reactions rates. 

The sun drives radiative, dynamical, and chemical processes affecting 

ozone and stratospheric temperature.

CFCs, HCFCs,

halons

volcanic and

nonvolcanic

aerosols

stratospheric

ozone

water vapor LLGHGs:

CO2, CH4, N2O

stratospheric

temperature

CH4

N2O
4.6

Chapter 4

water vapor will radiatively cool the lower stratosphere 

and also affect the frequency of occurrence of polar stratospheric clouds, thereby impacting stratospheric ozone 

chemistry (Kirk-Davidoff et al., 1999; Feck et al., 2008). 

Enhanced levels of stratospheric water vapor strengthen 

ozone loss in the presence of ODSs. Hence, a climate 

with increased stratospheric water vapor will have a delayed ozone recovery even while ODSs are reduced (Shindell, 2001; Shindell and Grewe, 2002; Tian et al., 2009). 

Changes in stratospheric water vapor also can be a significant radiative forcing for surface climate (see Section 

4.4.1).

The principal sources of stratospheric water vapor 

are entry through the tropical tropopause (Brewer, 1949) 

and oxidation of methane within the stratosphere (Jones 

et al., 1986; le Texier et al., 1988). Oxidation of molecular hydrogen (H2) is another source of stratospheric water 

vapor, albeit currently small, but with the potential to grow 

in the future if hydrogen fuel cells come into common use 

(Tromp et al., 2003; Schultz et al., 2003).

Each source of stratospheric water vapor is associated with a distinct timescale. The input of water vapor 

into the stratosphere by an individual air parcel in the tropics is largely a function of the lowest temperature a parcel 

encounters on its transit into the stratosphere, as originally 

noted by Brewer (1949) and more recently discussed in 

Schiller et al. (2009). The actual trajectory a parcel takes 

does need to be considered (Fueglistaler et al., 2009), but 

a simple model of horizontal processing of air by passage through the Western Pacific cold point tropopause 

reasonably reproduces observed stratospheric humidity 

(Holton and Gettelman, 2001; Geller et al., 2002; Scaife 

et al., 2003). Variability in stratospheric water vapor on 

seasonal and interannual timescales has been well reproduced by climatological trajectory studies using saturation 

mixing ratios calculated from global temperature analyses 

(Jensen and Pfister, 2004; Fueglistaler and Haynes, 2005), 

demonstrating that to first order, variability in the entry 

of water vapor into the stratosphere is controlled by variability in tropical cold point temperatures. Convection 

overshooting into the stratosphere has been observed on 

limited occasions, and when it occurs likely hydrates the 

stratosphere locally (Khaykin et al., 2009; Nielsen et al., 

2007). However, evidence for a global impact of this phenomenon is lacking at this time (Schiller et al., 2009). Because there is a relatively short turnover time for air in the 

lowermost stratosphere, on the order of months (Rosenlof 

and Holton, 1993), changes in the entry value of water vapor to the stratosphere due to tropical cold point temperature changes will be seen almost immediately throughout 

the lowermost stratosphere. However, there will be a 

time lag before the signal reaches the middle and upper 

stratosphere on the order of years. As noted in Engel et al. 

(2009) (and references therein), the mean age of air in the 

middle stratosphere at northern midlatitudes (between 20 

and 40 km) is approximately 4 years, hence the multiyear 

time lag before seeing a signal in the middle stratosphere.

Measurement of stratospheric water vapor concentrations is highly challenging and uncertain. There are 

significant discrepancies noted between coincident measurements of stratospheric water vapor concentrations 

using different in situ and satellite techniques (Vömel et 

al., 2007; Kley et al., 2000; Lambert et al., 2007; Weinstock et al., 2009). These discrepancies range from 10% 

to 50% or even greater in some cases and preclude combining data sets for trend analysis without extreme care. 

However, data quality is sufficient to examine annual and 

interannual variations of water vapor in the tropical lower 

stratosphere. Such variations have been noted to be in 

quantitative agreement with the idea that observed variations in tropical tropopause temperatures control the entry 

value of stratospheric water vapor (Randel et al., 2004; 

Fueglistaler and Haynes, 2005). Chapter 7 in SPARC 

CCMVal (2010) notes that most chemistry-climate models are able to reproduce the general sense of the annual 

cycle of water vapor in the tropical lower stratosphere 

with a minimum in Northern Hemisphere spring and a 

maximum in Northern Hemisphere fall and winter. There 

is a wide spread in the stratospheric entry value of water 

vapor in the models ranging from 2–6 parts per million by 

volume (ppmv). Kley et al. (2000) presented observationally based estimates ranging from 2.0–4.1 ppmv for the 

stratospheric entry value, and noted differences between 

measurement systems larger than the stated uncertainties 

for those instruments. Assessing the exact mechanism for 

the amount of dehydration of air entering the stratosphere 

requires better accuracy than currently exists. As concluded in Weinstock et al. (2009), the differences using coincident measurement noted between independent in situ instruments are sufficiently large that different conclusions 

can be reached with regard to the impact of convective 

processes in the tropical tropopause layer, and the degree 

of supersaturation that is plausible (Peter et al., 2006).

Global water vapor trend determination from the 

historical record is difficult. There are differences in trends 

noted between measurement systems covering the same 

time period (for example the Northern Hemisphere frost 

point balloon as compared with the Upper Atmosphere Research Satellite (UARS) Halogen Occultation Experiment 

(HALOE) satellite instrument, as noted in Randel et al., 

2004). The multidecadal stratospheric water vapor record 

is limited to Northern Hemisphere midlatitudes. The longest continuous record of stratospheric water vapor data 

is from frost point balloon measurements taken at 40°N 

from Boulder, Colorado. At present, the longest satellite 

records are from SAGE II and HALOE instruments. Both 

of these instruments ceased operation in 2005, and there 

are a number of newer satellite instruments currently mea-
4.7

Stratospheric Changes and Climate

suring stratospheric water vapor concentrations, including 

Aura Microwave Limb Sounder (MLS), which has extensive spatial coverage. Using overlap periods between 

instruments, it may be possible to continue estimation of 

global trends; however, this is a current research endeavor 

and there is no relevant literature to assess at this time.

From the historic record of the amount of stratospheric water vapor, an increase based on midlatitude frost 

point balloon measurements below 30 km in Washington, 

D.C., and Boulder, Colorado, on the order of 0.05 ± 0.01 

ppmv/yr for the period from 1964–2000 was reported by 

Oltmans et al. (2000). Corrections to the Boulder frost 

point data were reported by Scherer et al. (2008), which 

reduced the trend for the period from 1980–2000 to 0.03 ± 

0.005 ppmv/yr. The time series of the revised Boulder data 

is shown in Figure 4-2 as well as the comparable time series from HALOE, SAGE II, and Aura MLS. Independent 

data from a variety of remote sounding and in situ sources 

show an average trend for the period from 1960–2000 of 

0.045 ppmv/yr at Northern Hemisphere midlatitudes at 

levels below 30 kilometers (km) (Rosenlof et al., 2001). 

Analyses by Rohs et al. (2006) using in situ balloon measurements show that changes in methane mixing ratio can 

account for a midlatitude trend below 30 km of .0132 ± 

0.002 ppmv/yr of the Boulder increase. The exact mechanism for the remainder of the observed increase of water 

vapor concentration for the period ending in 2000 is in 

question, with circulation changes postulated related to the 

width of the tropics (Zhou et al., 2001; Rosenlof, 2002), as 

well as changes in aerosol processes near the tropical tropopause (Notholt et al., 2005; Sherwood, 2002). As shown 

in Figure 4-2 (lower stratosphere midlatitudes) and Figure 

4-3 (tropics), since the end of 2000, a decrease in the mixing ratio of water vapor entering the tropical stratosphere 

occurred (Randel et al., 2006; Rosenlof and Reid, 2008), 

coincident with a drop in tropical tropopause temperatures 

that has occurred during a period without an increase of 

methane concentration (Dlugokencky et al., 2009). The 

drop in tropical water vapor entry values estimated using 

HALOE data at 82 hectoPascals (hPa) is ~0.5 ppmv, or 

approximately 10% of average stratospheric water vapor 

values (Solomon et al., 2010), and 25% of the nominal 

maximum-to-minimum difference in the annual cycle of 

82 hPa tropical water vapor as estimated from HALOE 

measurements. The drops in tropical tropopause temperature and entry of water vapor into the stratosphere at the 

end of 2000 appear to be associated with an increase in 

the rate of tropical upwelling (Randel et al., 2006; Rosenlof and Reid, 2008) and associated changes in eddy wave 

driving (Dhomse et al., 2008). There is not agreement in 

the literature as to the reason for the strengthening of tropical upwelling near the tropical tropopause; both tropical 

Figure 4-2. Observed changes in stratospheric water vapor. 

Time series of stratospheric water vapor 

mixing ratio (ppmv) 

averaged from 70 to 

100 hPa near Boulder Colorado (40°N, 

105.25°W) from a balloonborne frost point 

hygrometer covering the period 1981 

through 2009; satellite measurements 

are monthly averages, balloon data 

plotted are from individual flights. Also plotted are zonally averaged satellite measurements in the 35°N–45°N 

latitude range at 82 hPa from the Aura MLS (turquoise squares), UARS HALOE (blue diamonds), and SAGE II 

instruments (red diamonds). The SAGE II and HALOE data have been adjusted to match MLS during the overlap period from mid-2004 to the end of 2005, as there are known biases (Lambert et al., 2007). Representative 

uncertainties are given by the colored bars; for the satellite data sets these show the uncertainty as indicated by 

the monthly standard deviations, while for the balloon dataset this is the estimated uncertainty provided in the 

Boulder data files. Figure adapted from Solomon et al. (2010).

Water Vapor

1980 1990 2000 2010

Time

2

3

4

5

6

Mixing Ratio (ppmv)

Satellite measurements, monthly zonal averages, 82 hPa, 35-45 N

Aura MLS, UARS HALOE (shifted to match MLS)

SAGE II, aerosol filtered and shifted to match MLS

Boulder water 70-100 hPa average

 
4.8

Chapter 4

sea surface temperature changes (Deckert and Dameris, 

2008) and changes in high-latitude wave forcing (Ueyama 

and Wallace, 2010) have been suggested.

There is a good understanding of the annual cycle 

of water vapor entering the stratosphere (Figure 4-3). The 

amplitude of the annual cycle is 50% to 60% of the mean 

and well explained by the known annual cycle in tropical tropopause temperatures (Reed and Vleck, 1969). In 

contrast, the trend in stratospheric water vapor is not well 

understood. Over the period 1950–2000 there was an increase in entry-level stratospheric water vapor on the order 

of 1%/yr (Rosenlof et al., 2001) during a period of increasing tropospheric methane and decreasing tropopause temperatures (Zhou et al., 2001). At the end of 2000 there was 

a decrease in stratospheric entry-level water vapor coincident with a step-like drop in tropical tropopause temperatures (Randel et al., 2006; Rosenlof and Reid, 2008). The 

observed long-term increase in stratospheric water vapor 

over the 1950–2000 period cannot be explained through 

tropical tropopause temperature trends, although some 

aspects of interannual variability can be. The more recent 

decrease in stratospheric water vapor can be explained 

by tropical tropopause temperature changes, although the 

mechanism driving that temperature change is not well 

understood. Given the uncertainties in our understanding 

and modeling of past water vapor changes, it is difficult to 

predict changes expected in a future climate.

4.1.4 Stratospheric Aerosols

The stratospheric aerosol layer has often been characterized as a “background” punctuated by volcanic enhancements. The composition of these aerosols is largely 

sulfuric acid/water solutions, and hence is strongly dependent on sources of stratospheric sulfur. Carbonyl sulfide 

(OCS) is an important source of sulfur to the stratosphere 

(Crutzen, 1976). However, the observed abundance of 

background stratospheric aerosol is many times larger 

10°N-10°S Water Vapor

1993 1994 1995 1996 1997 1998 1999 2000 2001 2002 2003 2004 2005 2006 2007 2008 2009 2010

100

10

Pressure (hPa)

20

30

40

50

125 2.3

2.6

2.9

3.2

3.5

3.8

4.1

4.4

4.7

5.0

ppm

v

Year

Figure 4-3. Tropical water vapor (ppmv, 10°N–10°S, monthly averages) plotted versus time, showing upward 

propagation of the water vapor tape recorder (Mote et al., 1996). This is a combination of UARS HALOE and 

Aura MLS measurements. During the period of data overlap (from mid-2004 through the end of 2005), differences were computed for matching profiles at each pressure level. That average shift was applied to the 

HALOE measurements at each level; for 82 hPa it is on the order of 0.5 ppmv. The key feature to note here is 

the change to lower values of the water vapor minimum (hygropause) at the end of 2000, and upward propagation of those lower values in subsequent years. Update of Figure 10 from Rosenlof and Reid (2008).
4.9

Stratospheric Changes and Climate

than can be explained using OCS alone (Chin and Davis, 

1995; Weisenstein et al., 1997; Pitari et al., 2002), which 

suggests an important role for other sources—such as 

sulfur dioxide (SO2)—in pollution. Explosive volcanic 

eruptions that occurred in the past several decades include 

Mt. Pinatubo in 1991 and El Chichón in 1982, and these 

eruptions increased the integrated stratospheric aerosol 

abundance by more than a factor of ten, as shown in Figure 4-4. Volcanic aerosols directly affect stratospheric 

temperatures (as discussed in this chapter), as well as midlatitude and polar surface chemistry and thus stratospheric 

ozone depletion (as discussed in Chapters 2 and 3). Large 

explosive eruptions also cause episodic cooling of global 

average surface temperatures for a few years and other 

climate effects (see Section 4.4.1). There is presently no 

systematic global monitoring system to document longterm future changes in stratospheric aerosol that could 

affect ozone and climate.

Stratospheric aerosols have been measured at a few 

key sites using balloonborne optical counters and laser ranging (lidar) methods, beginning at some stations in the early 

to mid-1970s (e.g., Hofmann, 1990; Jäger, 2005; Deshler, 

2008). Systematic global satellite measurements using visible spectroscopy (SAGE) began in the mid-1980s (Thomason et al., 1997) but were terminated in 2005. The data 

sets, methods used, their intercomparison, and the range of 

available records were recently reviewed under the auspices 

of SPARC (SPARC, 2006) and by Deshler (2008).

10-5

10-4

10-3

10-2 10-5

10-4

10-3

10-2

104

105

106

1970 1975 1980 1985 1990 1995 2000 2005 2010

104

105

106

Hampton (37oN, 76oW), Boulder (40°N, 105°W) Garmisch (47

o

N, 11oE)

Integrated Backscatter (sr-1

)

Latitude

< 30o

> 30o

São José dos Campos (23oS, 46oW) Mauna Loa (20

o

N, 156oW)

Laramie (41 20-25 km oN, 105oW)

Laramie N(r > 0.15 m)

 N(r > 0.25 m)

Latitude

< 30o

> 30o

15-20 km

N(r > 0.15 m)

 N(r > 0.25 m)

5 km Aerosol column (cm-2

)

 Year

10-4

10-4

104

105

106

1995 2000 2005 2010

104

105

106

Hampton (37

o

N, 76oW) Boulder (40°N, 105°W) Garmisch (47

o

N, 11oE)

Integrated Backscatter (sr-1

)

São José dos Campos (23

o

S, 46oW) Mauna Loa (20

o

N, 156oW)

20-25 km, Laramie (41

o

N, 105oW) N(r > 0.15 m)

 N(r > 0.25 m)

15-20 km, Laramie (41

o

N, 105oW)

5 km Aerosol column (cm-2

)

 Year

Figure 4-4. History of stratospheric integrated optical backscatter (sr–1) at 694 nm from lidar measurements at 

five locations (top two panels) and 5 km aerosol column concentration (cm–2) from in situ measurements over 

two altitude intervals above the tropopause at Laramie, Wyoming, USA (bottom two panels). Top panel shows 

measurements from São José dos Campos, Brazil, integration 17–35 km, and Mauna Loa, Hawaii, USA, integration 15.8–33 km. Second panel show measurements from Hampton, Virginia, USA, integration tropopause 

to 30 km, Boulder, Colorado, USA, integration 20–33 km, and Garmisch-Partenkirchen, Germany, integration 

tropopause +1 km - layer top. The measurements from São José dos Campos (589 nm), Boulder (532 nm), 

and Mauna Loa (532 nm since 1999) are scaled to 694 nm using a wavelength exponent of −1.4. The times of 

volcanic eruptions are indicated in the top and bottom panels with triangles, separated into eruptions at latitudes 

less (green upper symbols) and greater (blue lower symbols) than 30 degrees. Eruptions with volcanic explosivity indices of 5 (large closed symbols), 4 (small open symbols), and 4 with some uncertainty (tiny open symbols) 

are shown. This figure extends that presented by Deshler (2008) and Hofmann et al. (2009). The Hampton 

measurements have been discontinued. Right multipanel plot is an expansion of the data since 1994.
4.10

Chapter 4

Hofmann (1990) noted an apparent increase in 

stratospheric aerosols between the late 1970s and the 

late 1980s, which are two periods with little volcanic 

influence (Figure 4-4); he suggested a positive trend 

in the nonvolcanic aerosol background of about 5%/yr 

over that decade. However, following the major eruption of Mt. Pinatubo in June 1991, stratospheric aerosol declined to the lowest values observed in at least 

two decades. Taken over the period 1970–2005, there 

has been no significant trend in background aerosols 

(Deshler, 2008), raising questions about the origin of 

the positive trend in the earlier data. However, while 

recent data remain close to that of the 1970s, they also 

reveal trends over limited time intervals. A closer look 

at the most recent data from numerous sites reveals 

increases of about 4%/yr to 7%/yr in backscatter from 

20–30 km since the late 1990s (see Hofmann et al., 

2009 and insets in Figure 4-4).

Hofmann et al. (2009) suggested that these recent 

increases could be linked to sulfur emissions from coal 

burning in China, which has dramatically increased in the 

past decade. Notholt et al. (2005) noted the importance 

of the Asian summer monsoon for transport of SO2 to 

the tropical upper troposphere and for cross-tropopause 

transport, which they suggest could affect stratospheric 

water vapor transport. Such transport has the potential to 

influence the source of sulfur to the stratosphere as well. 

However, input from volcanoes, including from some 

less explosive eruptions previously thought to be small, 

may be more important than previously thought. For 

example, spaceborne laser ranging (lidar) observations 

at high resolution show evidence for substantial volcanic 

inputs to stratospheric aerosol associated with the eruption of Soufriere on Montserrat in mid-2006 (Vernier 

et al., 2009). The conclusion is that decadal variability 

in stratospheric aerosol is larger than previously anticipated. The relative contribution of recent anthropogenic 

versus natural emissions to changes in stratospheric 

aerosol loading remains an area of active research.

4.2 OBSERVED VARIATIONS IN 

STRATOSPHERIC CLIMATE

4.2.1 Observations of Long-Term Changes 

in Stratospheric Temperature

Substantial progress has been made since the 

2006 Assessment (WMO, 2007) in the evaluation of past 

stratospheric temperature changes. As a result of this 

recent work, we have increased confidence relative to 

previous assessments of the magnitude and meridional 

structure of temperature trends in the lower stratosphere. 

We also have a better understanding of the errors inherent in measurements of temperature changes in the middle and upper stratosphere. Three factors contribute to 

the changes in our understanding: 

1. Improved knowledge of the inherent uncertainties in 

stratospheric data derived from spaceborne instruments (e.g., CCSP, 2006; Mears and Wentz, 2009; 

Shine et al., 2008) and radiosondes (e.g., Lanzante 

et al., 2003; Free et al., 2005; Sherwood et al., 2005; 

Thorne et al., 2005; CCSP, 2006; Free and Seidel, 

2007; Randel et al., 2009), and in stratospheric products available from reanalysis products (e.g., Randel et 

al., 2009; Section 2.4 in Chapter 2 of this Assessment); 

2. The emergence of several independent analyses 

of satellite and radiosonde data sets, with distinct 

approaches to homogeneity adjustments (e.g., Free et 

al., 2005; Haimberger, 2007; Haimberger et al., 2008; 

Thorne et al., 2005; Randel and Wu, 2006; Sherwood 

et al., 2008); and 

3. The lengthening of data records with the passing 

of time.

There are now six global lower-stratospheric 

temperature data sets specifically developed for climate 

studies based on radiosonde data: RATPAC (Free et al., 

2005); HadAT (Thorne et al., 2005); RATPAC-lite (an 

abridged version of the RATPAC data set; Randel and 

Wu, 2006); RAOBCORE (Haimberger, 2007); RICH 

(Haimberger et al., 2008); and IUK (Sherwood et al., 

2008) (see Appendix B for definitions of these acronyms). 

The radiosonde data sets are not fully independent, but 

their different approaches to identifying and adjusting 

temporal inhomogeneities that can affect trends (particularly in the stratosphere) help us to characterize the 

overall uncertainty in estimates of long-term stratospheric 

temperature change since the late 1950s. There are now 

three lower-stratospheric temperature data sets derived 

from Microwave Sounding Unit (MSU) and Advanced 

MSU (AMSU) observations from polar-orbiting satellites 

since 1979 (University of Alabama-Huntsville (UAH), 

Christy et al., 2003; Remote Sensing Systems (RSS), 

Mears and Wentz, 2009; Center for Satellite Applications 

and Research (STAR), Zou et al., 2009). The lowerstratospheric MSU/AMSU temperature data are derived 

by blending MSU Channel 4 with AMSU Channel 9 data 

(see the discussion in Randel et al., 2009). The blended 

data are hereafter referred to simply as MSU4.

Figure 4-5 (from Thorne, 2009) presents time 

series of global-mean lower-stratospheric temperatures 

from five radiosonde and three MSU4 data sets. The 

radiosonde data have been vertically weighted as per the 
4.11

Stratospheric Changes and Climate

MSU4 weighting function (the general time evolution of 

lower-stratospheric temperatures shown in Figure 4-5 is 

mirrored in time series based on global-mean radiosonde 

data at individual standard pressure levels between 100 

and 30 hPa; not shown). The figure is an updated and 

extended version of the global-mean time series shown 

in, for example, Ramaswamy et al. (2001, 2006), Seidel 

and Lanzante (2004), CCSP (2006), and Thompson and 

Solomon (2009). Three key aspects of global-mean lowerstratospheric temperature changes are evidenced in all 

data sets:

1. In the global mean, the lower stratosphere has cooled 

by ~0.5 K/decade since 1980 (~0.35 K/decade in 

RAwinsonde OBservation (RAOB) data extended 

back to 1958). The robustness of the global-mean 

lower-stratospheric cooling has been documented 

in numerous recent studies (e.g., see the recent review by Randel et al., 2009), but varies slightly from 

data set to data set. For example, the cooling during 1980–2008 is 0.33 to 0.42 K/decade for the three 

MSU4 data sets but 0.50 ± 0.16 K/decade for the 

(vertically weighted) radiosonde data sets (Thorne, 

2009; Figure 4-5).

2. The global-mean lower-stratospheric cooling has not 

occurred linearly but rather appears to be manifested as 

two downward steps in temperature coincident with the 

end of the transient warming associated with explosive 

volcanic eruptions. The steps are most pronounced 

after the eruptions of El Chichón (1982) and Mt. 

Pinatubo (1991) and have been emphasized in numerous studies (Pawson et al., 1998; Seidel and Lanzante, 

2004; Ramaswamy et al., 2006; Eyring et al., 2006; 

Free and Lanzante, 2009; Thompson and Solomon, 

2009). Thompson and Solomon (2009) argue that the 

steps are consistent with the superposition of (i) longterm stratospheric cooling; (ii) transient warming due 

to volcanic aerosols loading; and (iii) transient cooling 

due to volcanically induced ozone depletion.

3. In the global mean, the lower stratosphere has not 

cooled noticeably since 1995. Global-mean lowerstratospheric temperatures during the period following 1995 are significantly lower than they were during the decades prior to 1980, but have not dropped 

further since 1995.

Another key aspect of recent stratospheric temperature trends is the near uniformity of the cooling 

at all latitudes outside of the polar regions since 1980. 

Trends based on lower-stratospheric data from multiple radiosonde data sets show cooling of ~0.4 K/decade 

(from RATPAC-lite data; Randel and Wu, 2006) to ~0.8 

1960 1970 1980 1990 2000 2010

Year

Lower Stratospheric Temperature Anomaly (°C)

-1.0

0.0

1.0

Temperature Anomaly Difference (°C)

1.0

0.0

-1.0

HadAT - Mean

RATPAC - Mean

IUK - Mean

RAOBCORE - Mean

RICH - Mean

RSS - Mean

UAH - Mean

Radiosonde Mean STAR - Mean

Figure 4-5. Global-mean lower-stratospheric temperature anomalies (1958–2008) from multiple data sets, 

including five radiosonde data sets (HadAT, IUK, RAOBCORE, RATPAC, and RICH) and three satellite MSU 

data sets (RSS, UAH, and STAR). Acronyms are defined in Appendix B of this Assessment. All time series are 

for the layer sampled by MSU channel 4, spanning 10–25 km in altitude, with a peak near 18 km. Black curve 

is the average of all available radiosonde data sets, and the colored curves show differences between individual 

data sets and this average. (Based on Thorne, 2009.)
4.12

Chapter 4

K/decade (from RATPAC data; Free et al., 2005) for 

1980–2008 for zonal bands within about 45 degrees north 

and south of the equator (Figure 4-6 top; Thompson and 

Solomon, 2005; Free et al., 2005). The presence of significant lower-stratospheric cooling at tropical latitudes 

(Thompson and Solomon, 2005) has implications for 

the attribution of changes in stratospheric circulation, 

as discussed in Sections 4.2.2 and 4.3. The structure 

and amplitude of the lower-stratospheric cooling also 

has implications for the interpretation of tropospheric 

temperature trends estimated from the MSU2 satellite, 

since the MSU2 weighting function samples the lowermost stratosphere (e.g., Fu et al., 2004).

In the annual mean, the cooling of the tropical and 

middle latitudes since 1980 occurred primarily before 1995 

(compare the top and bottom panels of Figure 4-6). Since 

1995, annual-mean lower-stratospheric temperatures have 

remained steady over much of globe, albeit with significant 

rises over the polar regions in one but not all data sets (the 

IUK data; Figure 4-6 bottom). The drop in tropical tropopause temperatures circa 2001 highlighted in Section 4.1.3 

is centered on a very narrow layer about the tropical tropopause (Randel et al., 2006) and is not apparent in trends at 

50 hPa (which are shown in Figure 4-6 bottom).

Lower-stratospheric temperature trends also exhibit 

considerable seasonal variability. The top panel in Figure 

4-7 (from Fu et al., 2010; see also Figure 11 in Randel et 

al., 2009) shows updated trends in the RSS MSU4 data as 

a function of latitude and calendar month. Regions of 

90% significance are denoted by hatching. The bottom 

panel in Figure 4-7 shows time series of polar stratospheric 

temperatures averaged over the seasons indicated based on 

radiosonde data (from Randel et al., 2009). The tropical 

cooling evident in the annual-mean is largest between 

June and January (Figure 4-7, top). Between 1979 and 

2007, the Southern Hemisphere polar regions are marked 

by significant (at the 90% level) cooling between November and March. During the same period, the Northern 

Hemisphere polar regions are marked by significant cooling during March/April and June–September. The polar 

warming in August–September in the Southern Hemisphere and December–January in the Northern Hemisphere are not statistically significant at the 90% level in 

the zonal mean (Figure 4-7, top).

Time series of 100 hPa polar temperature anomalies from radiosonde data confirm visually the following 

aspects of lower-stratospheric temperature trends (Figure 

4-7, bottom): (1) the robust cooling of the polar regions 

during the spring and summer seasons in both hemispheres 

(with notably larger cooling observed in the Antarctic), 

and (2) the absence of significant temperature trends 

during the winter season in both hemispheres.

In contrast to the lower stratosphere, temperature 

changes in the middle and upper stratosphere are relatively 

uncertain due to the limited availability of long-term temperature data there. Radiosonde data are generally available only up to about 20 hPa, and the quality of radiosonde 

data diminishes with height. There is currently only one 

satellite data record (from the Stratospheric Sounding Unit) 

and one corresponding analysis (the analysis combines 

the available SSU zonal temperature anomaly data from 

ten separate satellites and is available through 2005; see 

Randel et al., 2009). Temperature trends based on lidar 

measurements have large sampling limitations and are only 

available at limited locations throughout the globe (see discussions in Randel et al., 2009, and Funatsu et al., 2008).

The Stratospheric Sounding Unit senses emissions 

from carbon dioxide and thus is sensitive to the increases 

in CO2 over the past few decades (Shine et al., 2008). For 

this reason, trends derived from Stratospheric Sounding 

3

−1

0

1

2

1995–2008

K / dec

RAOBCORE

60°S 30°S EQ 30°N 60°N

RLite

RICH

IUK

RATPAC HadAT2

60°S 30°S EQ 30°N 60°N

1979–2008

3

−1

0

1

2

K / dec

Figure 4-6. Zonal-mean temperature trends (K/

decade) at 50 hPa from six adjusted radiosonde data 

sets for the periods 1979–2008 (top) and 1995–2008 

(bottom). Gray shading indicates the 2-sigma trend 

confidence interval from the RATPAC data set 

(others are comparable). Trends are computed from 

temperature anomaly time series, omitting data for 

two years after the El Chichón and Mt. Pinatubo volcanic eruptions.
4.13

Stratospheric Changes and Climate

Unit data that are not treated for the influence of increasing 

CO2 (i.e., all Stratospheric Sounding Unit trends published 

prior to 2008) are affected by uncorrected changes in the 

retrieval weighting function. The principal effect of correcting the Stratospheric Sounding Unit weighting function for increasing CO2 is to increase the cooling trends 

by as much as ~0.2–0.4 K/decade throughout much of the 

stratosphere (see Figure 4 in Shine et al., 2008). Recent 

analyses of Stratospheric Sounding Unit temperature data 

corrected for the increases in atmospheric CO2 are summarized in Randel et al. (2009; compare Figures 18 and 19), 

and the updated figures of 60°S-60°N mean Stratospheric 

Sounding Unit temperatures are shown in Figure 4-8. The 

Stratospheric Sounding Unit temperature data suggest that 

(1) the middle and upper stratosphere cooled more rapidly 

than the lower stratosphere (~1.5 K/decade for 1980–2005 

for channels centered ~40–50 km) and (2) stratospheric 

temperatures remained steady from ~1995–2005 from the 

lower stratosphere up to ~1 hPa.

The outlook for evaluation of future changes in 

stratospheric temperature is mixed. It appears likely that 

multiple radiosonde and MSU/AMSU lower-stratospheric 

temperature analyses will continue to be available from 

several research teams. The recent initiation of a reference upper-air observing network (Seidel et al., 2009) 

bodes well for the eventual availability of high quality 

temperature (and water vapor and other) observations to 

calibrate and evaluate satellite and radiosonde data. Other 

data sets that will likely prove useful for future analyses of 

stratospheric temperatures include lidar deployed within 

the Network for the Detection of Atmospheric Composi- Network for the Detection of Atmospheric ComposiTemperature Anomaly (K) Temperature Anomaly (K)

Month

Latitude

K/Decade

1 2 3 4 5 6 7 8 9 10 11 12

−80

−60

−40

−20

0

20

40

60

80

−2.1

−1.4

−0.7

0

0.7

1.4

2.1

Figure 4-7. (top) Zonal-mean lower-stratospheric 

temperature trends (K/decade) for 1979–2007 for 

each calendar month from MSU4 observations 

(MSU4 spans roughly 10–25 km in altitude, with a 

peak near 18 km). The color contour interval is 0.35 

K/decade. Warm colors indicate warming; cool colors indicate cooling. Hatching indicates where the 

trends are significant at the 90% confidence level. 

Results reproduced from Fu et al. (2010). (bottom) 

Time series of 100 hPa temperature anomalies (K) 

averaged over the polar regions based on radiosonde 

measurements (adapted from Randel et al., 2009).

27

26

25

36x

26x

° °

Temperature Anomaly (K) Figure 4-8. Time series of SSU temperature anomalies (K) for channels indicated (Figure 18 from Randel et al., 2009). Data for channels 26x and 36x 

are shifted for clarity. Exact weighting functions for 

the SSU satellite instrument can be found in Figure 

1 of Randel et al. (2009). Channel 27 corresponds 

to ~34–52 km altitude, channel 36x to ~38–52 km, 

channel 26 to ~26–46 km, channel 25 to ~20–38 km, 

and channel 26x to ~21–39 km.
4.14

Chapter 4

tion Change and the Global Positioning System radio oc and the Global Positioning System radio occultation temperature profiles (the latter are available continuously starting in 2001). For the upper stratosphere, 

Stratospheric Sounding Unit data could be merged with 

AMSU observations to extend the record past 2005, but 

the lack of multiple analyses will continue to limit our 

understanding of temperature changes in the middle and 

upper stratosphere.

4.2.2 Observations of Long-Term Changes 

in the Stratospheric Circulation

4.2.2.1 Stratospheric Zonal Flow

Assessment of long-term trends in the stratospheric circulation is more difficult than it is for temperature 

because 

1. Observations of horizontal stratospheric winds from 

rawinsondes are sparse and, in general, have not been 

homogenized for changes in instrumentation. 

2. Estimates of the meridional overturning (i.e., the 

Brewer-Dobson) circulation based on chemical and 

age-of-air measurements are very noisy and have error bars that exceed the amplitude of the observed 

trends (Baldwin and Dameris et al., 2007; Engel et 

al., 2009). 

Observed trends in the stratospheric horizontal wind field 

must be inferred indirectly from measurements of atmospheric temperature (derived from both radiosonde and 

spaceborne instruments) and atmospheric geopotential 

height (derived from radiosonde measurements through 

application of the hypsometric equation).

The trends in the Southern Hemisphere stratospheric vortex during austral spring are well established and 

were assessed in the 2006 Ozone Assessment (Baldwin 

and Dameris et al., 2007). They are revisited briefly here 

since they provide important context for the tropospheric 

trends assessed in Section 4.4.2. As assessed in the 2006 

Assessment, temperatures and geopotential heights both 

dropped throughout the Southern Hemisphere polar stratosphere from ~1970 to the late 1990s (Thompson and Solomon, 2002). Figure 4-9 shows such trends extended to 

2003 (from Thompson and Solomon, 2005). More recent 

updates of trends in polar geopotential height are not available. Since the trends in geopotential height are relatively 

small at middle latitudes (not shown), the polar trends in 

geopotential height shown in Figure 4-9 are consistent 

with an anomalous eastward acceleration of the Southern 

Hemisphere stratospheric polar vortex. The eastward acceleration of the Southern Hemisphere stratospheric vortex between November and January is consistent with 

diabatic cooling associated with the Antarctic ozone hole 

(Waugh et al., 1999; Thompson and Solomon, 2002) and 

is reflected in a delay in the dynamical breakdown of the 

Southern Hemisphere stratospheric vortex (Waugh et al., 

1999; Zhou et al., 2000; Karpetchko et al., 2005; Haigh 

and Roscoe, 2009).

The Northern Hemisphere polar stratosphere cooled 

markedly during winter and spring between the 1960s and 

the late 1990s (e.g., WMO, 2003), but exhibited a string of 

warm winters during the 2000s (Section 4.2.1; Randel et 

al., 2009). Hence, winter and springtime trends in Northern Hemisphere stratospheric zonal-flow are less significant and more difficult to interpret than those in the Southern Hemisphere. As evidenced in the time series in Figure 

4-7 (bottom), the Northern Hemisphere polar stratospheric 

temperature trends are not significant during December–

February. The statistically significant cooling and geopotential height decreases in the Northern Hemisphere polar 

stratosphere during summer (Figure 4-7; Figure 4-9) are 

not associated with marked meridional gradients, and thus 

do not imply changes in the stratospheric thermal wind.

4.2.2.2 Brewer-Dobson Circulation

The Brewer-Dobson circulation model is a simple circulation suggested by Brewer (1949) and Dobson 

(1956), and consists of three basic parts. The first part 

is the rising tropical motion of air from the troposphere 

into the stratosphere. The second part is poleward transport of air in the stratosphere. The third part is descending 

motion of air in both the stratospheric middle and polar 

latitudes. This seemingly simple picture leaves a number 

of possible ambiguities. For instance, the distribution of 

the tropical air rising through the tropical tropopause is 

important in that air rising in the inner tropics (near the 

equator) encounters lower tropopause temperatures than 

air rising in the outer tropics. The latitudinal distribution of the descending air is also important since middle 

latitude descending air is transported back into the troposphere (as a result of isentropic mixing), while the polar 

latitude descending air is transported into the polar lower 

stratosphere (via diabatic descent). The Brewer-Dobson 

circulation is driven primarily by the dissipation of Rossby 

and gravity waves that have propagated upward from the 

troposphere. Convective overshooting may also play a 

role in the vertical transport of trace species in the tropical 

lower stratosphere, but the role of convective overshooting is less understood than the role of stratospheric wave 

drag (e.g., Fueglistaler et al., 2009).

Measures of the Brewer-Dobson circulation include 

the upward flux of tropical air, but if this air descends elsewhere in the tropics, it should not be considered a part 

of the Brewer-Dobson circulation. The net upward mass 
4.15

Stratospheric Changes and Climate

flux of air into the tropical stratosphere is often taken as a 

measure of the Brewer-Dobson circulation, but this does 

not distinguish the latitudinal distribution of the rising 

tropical air, which can affect the ozone distribution differently. Theory (e.g., Plumb and Eluszkiewicz, 1999; 

Semeniuk and Shepherd, 2001; Zhou et al., 2006) indicates that for annually averaged upwelling at the equator to exist, wave Eliassen-Palm flux convergences must 

extend into the inner tropics (to about 12–15° latitude), 

and the distribution of the descending motions at middle 

and high latitudes depends on the distribution of Eliassen 

and Palm flux convergences there. However, the importance of tropical wave drag in driving the Brewer-Dobson 

circulation has recently been questioned by Ueyama and 

Wallace (2010), on the basis of their observational analysis. Thus interpreting trends in the modeled and observed 

Brewer-Dobson circulation will require further research.

As discussed later in Section 4.3.2, climate model 

simulations consistently predict an acceleration of the 

Brewer-Dobson circulation in response to increasing 

greenhouse gases (e.g., Rind et al., 1998; Butchart and 

Scaife, 2001), amounting to an average increase of about 

2%/decade in the annual mean net upward mass flux at 

70 hPa through the 21st century (Butchart et al., 2006; 

McLandress and Shepherd, 2009; Butchart et al., 2010; 

Chapter 4 of SPARC CCMVal, 2010). Since the BrewerDobson circulation partially determines the distribution 

of stratospheric ozone, the simulated trends in the 

J A S O N D J F M A M J J M A M J J A S O N D J F M

30

50

100

200

500

1000

Pressure (hPa) 30

50

100

200

500

1000

Pressure (hPa)

60°N-90°N 60°S-90°S

Temperature trends from radiosonde data

Geopotential height trends from radiosonde data

-3.75

-88 -200

-24

-2.25

-0.75

J A S O N D J F M A M J J M A M J J A S O N D J F M

Month Month

Figure 4-9. Trends in (top) temperature (K/decade) and (bottom) geopotential height (m/decade) averaged 

over (left) 60°N–90°N and (right) 60°S–90°S for 1979–2003. Trends are shown as a function of month and 

pressure level. Shading denotes trends that exceed the 95% confidence level. Tick marks on the abscissa denote the center of the respective month. Note that the calendar months on the abscissa are shifted between the 

Northern and Southern Hemisphere. Trends based on radiosonde data from the Integrated Global Radiosonde 

Archive (IGRA). Units: K/decade (−0.25, 0.25, 0.75...) and m/decade (−8, 8, 24...). Based on Thompson and 

Solomon (2005).
4.16

Chapter 4

Brewer-Dobson circulation link human emissions of CO2

with the distribution of stratospheric ozone. The simulated trends have implications for stratospheric ozone 

and water vapor trends, and thus for the radiative forcing 

of the troposphere and surface climate (as discussed in 

Section 4.3.2).

The detection of trends in the Brewer-Dobson circulation in observations is complicated by two factors: 

1. The trends in the Brewer-Dobson circulation are 

small through 2010, and only in the next few decades 

are predicted to depart from the natural variability 

(Section 4.3.2).

2. The Brewer-Dobson circulation is not a measurable 

quantity and hence trends in the Brewer-Dobson circulation cannot be directly observed, but rather are 

inferred from changes in the horizontal structure of 

lower-stratospheric temperature, constituent trends, 

and trends in the estimated age of air.

The three primary lines of observational evidence 

for increases in the strength of the lower-stratospheric 

meridional circulation include (1) cooling in the lower 

stratosphere that exceeds the cooling predicted by ozone 

depletion in the tropics but is less than the cooling predicted as a direct radiative response to ozone depletion 

(Thompson and Solomon, 2009); (2) out-of-phase temperature trends between polar and tropical latitudes during the Northern Hemisphere and Southern Hemisphere 

cold seasons, with largest cooling in the tropics (Thompson and Solomon, 2009; Fu et al., 2010); and (3) localized decreases in ozone in the lower tropical stratosphere 

in trends derived from the SAGE instruments (Chapter 2 

and Figure 2-27). In general, the evidence for changes 

in the Brewer-Dobson circulation based on the structure 

of lower-stratospheric ozone and temperature trends is 

most robust in the tropics and less clear at middle and 

high latitudes.

There are two primary caveats associated with the 

above lines of evidence: 

1. Stratospheric age-of-air estimates based on balloonborne in situ measurement of sulfur hexafluoride 

(SF6) and CO2 do not indicate statistically significant 

trends in the Brewer-Dobson circulation (Engel et al., 

2009), albeit age-of-air estimates do not uniquely reflect the Brewer-Dobson circulation, and the uncertainty bars on such estimates are very large (i.e., the 

modeled trends in the Brewer-Dobson circulation lie 

within the uncertainty estimates of the stratospheric 

age-of-air measurements in numerous climate change 

simulations).

2. Trends in tropical lower-stratospheric ozone from the 

SAGE data are subject to considerable uncertainty 

(Chapter 2).

4.3 SIMULATIONS OF STRATOSPHERIC 

CLIMATE CHANGE

4.3.1 Simulation of Stratospheric Temperature Trends from Chemistry-Climate 

Models and Climate Models

Multiyear observations of stratospheric temperature clearly indicate both large variability on multiple 

timescales and long-term changes (Section 4.2.1). The 

thermal structure of the stratosphere is influenced by natural as well as anthropogenic factors. An important task is 

to understand the impacts of natural forcing on the stratosphere to enable the identification and quantification of 

implications of human activities. Several studies have 

been performed to describe the evolution of stratospheric 

temperature, in particular the observed cooling (e.g., 

Ramaswamy et al., 2006; Dall’Amico et al., 2010a; Gillett 

et al., 2010; Randel et al., 2009; see Figures 4-5 and 4-9 

for example time series of stratospheric temperature).

Attribution of stratospheric temperature variations 

and long-term changes can be undertaken through comparison of observed changes with simulations by a hierarchy of numerical models considering radiative, dynamical, and chemical processes in the atmosphere. Such 

numerical models can be binned into two general classes: 

(1) climate models, i.e., atmospheric general circulation 

models that are coupled to an ocean model, but mostly do 

not resolve the whole stratosphere; and (2) chemistryclimate models, i.e., atmospheric general circulation models that are interactively coupled to a detailed chemistry 

module, almost fully consider the stratosphere, but so far 

are generally not coupled to an ocean model. Both types 

of models can be used to assess both natural and anthropogenic forcings affecting the behavior of the atmosphere.

The prime natural forcings of stratospheric temperature fluctuations on interannual and decadal timescales 

are the solar activity cycle (with a timescale of 11 years), 

the El Niño-Southern Oscillation (with a timescale of 

~3–5 years), and explosive volcanic eruptions (with sporadic timescales). The quasi-biennial oscillation in tropcal stratospheric zonal winds is internally generated but 

can be considered as a forcing for extratropical stratospheric variability. Most climate models and chemistryclimate models consider the ~11-year activity cycle of the 

sun (i.e., solar variability is explicitly enforced by variation in radiative heating and photolysis rates) and the radiative and chemical effects of sporadic large volcanic 
4.17

Stratospheric Changes and Climate

eruptions (i.e., changes in heating rates and heterogeneous 

chemistry due to an enhanced aerosol loading). Climate 

models predict ocean temperatures and to varying degrees 

are able to capture coupled air-sea variability associated 

with the El Niño-Southern Oscillation (e.g., AchutaRao 

and Sperber, 2006). In most chemistry-climate models, 

ocean surface temperatures are prescribed, using either 

historical measurements or values derived from climate 

model simulations. In the latter case, the realism of the El 

Niño-Southern Oscillation and its related variability is 

constrained by the realism of the prescribed (climate 

model-calculated) sea surface temperatures.

Atmospheric models reaching up into the mesosphere and with sufficient vertical resolution in the whole 

model domain (i.e., less than 1 km) are able to generate 

stratospheric quasi-biennial-like oscillations (e.g., Giorgetta et al., 2006; Punge and Giorgetta, 2008). Additionally, an adequate description of stratospheric quasibiennial-like oscillations in stratosphere-resolving climate 

models requires small-scale gravity wave forcing (e.g., 

Takahashi, 1999; Scaife et al., 2000a) which is consistent 

with recent observational estimates (e.g., Ern and Preusse, 

2009). Chemistry-climate models that do not generate a 

stratospheric quasi-biennial oscillation internally mostly 

have the observed behavior prescribed, i.e., by assimilating observed wind fields in the tropical lower stratosphere 

(see Morgenstern et al., 2010a).

In climate models and chemistry-climate models, 

the impact of human activities is generally considered as 

follows: long-lived greenhouse gases (i.e., CO2, CH4, and 

N2O), anthropogenic aerosol loading, and ODSs (CFCs, 

HCFCs, and halons) are prescribed as either emissions or 

atmospheric concentrations. In chemistry-climate models, 

changes in the amount and distribution of water vapor and 

ozone are calculated in a self-consistent manner considering interactions of chemical, radiative, and dynamical 

(transport and mixing) processes, whereas in climate 

models, ozone fields are mostly prescribed without considering feedbacks; and effects of chemical processes 

on water vapor are neglected. Natural forcings such as 

changes in solar activity and explosive volcanic eruptions 

are also taken into account. Detailed summaries of suggested and used boundary conditions for recent chemistryclimate model simulations (see Chapter 3 of this report) 

have been given in Eyring et al. (2008), in Chapter 2 of 

the SPARC CCMVal report (2010), and in Morgenstern 

et al. (2010a)1.

1 The CCMVal project uses chemistry-climate models to simulate 

the atmosphere from 1960 to about 2100. CCMVal-2 is the second 

CCMVal project (SPARC CCMVal, 2010). The scenarios used in 

CCMVal-2 are outlined in Chapter 2 of the SPARC CCMVal (2010) 

report. REF-B1 (1960-2006) is defined as a transient run from 

1960 (with a 10-year spin-up period) to the present (see Table 3-2 

of this Assessment).

Figure 4-10 shows the time series of global-mean 

stratospheric temperature anomalies as derived from these 

chemistry-climate model simulations (colored lines) compared to satellite data (Microwave Sounding Unit/Stratospheric Sounding Unit) weighted over specific vertical 

levels (black lines). There is strong evidence for a large 

and significant cooling in most of the stratosphere during 

the last decades (Section 4.2.1; e.g., Randel et al., 2009; 

Randel, 2010; Thompson and Solomon, 2009). The observed stratospheric cooling has not evolved uniformly 

since the 1960s and the overall development is well reproduced by the majority of chemistry-climate models (see 

Figure 4-10).

Ramaswamy et al. (2006) noted that the observed 

cooling of the global lower stratosphere between 1979 and 

2003 occurred in two distinct step-like transitions (see also 

Section 4.2.1). After a 2–3 year period of stratospheric 

warming due to enhanced aerosol loading of the stratosphere, these appeared subsequent to the two explosive 

volcanic eruptions, El Chichón (April 1982) and Mt. Pinatubo (June 1991), with each cooling transition being followed by a period of relatively steady temperatures. 

Ramaswamy et al. (2006) concluded that the anthropogenic factorsforced the overall cooling and the naturalfactors modulated the temporal evolution of the cooling. 

Simulations performed with chemistry-climate models 

support these findings: the space-time structure of the observed cooling is largely attributable to the combined effect of stratospheric ozone depletion and increases in 

greenhouse gas concentrations, superimposed on effects 

due to solar irradiance variation and explosive volcanic 

eruptions (see Chapter 3 of SPARC CCMVal, 2010). 

Moreover, Thompson and Solomon (2009) provided observational analyses that indicate the step-like behavior of 

global mean stratospheric temperatures depends also on 

the temporal variability in global mean ozone following 

large volcanic eruptions. They argued that the warming 

and cooling pattern in global mean temperatures following 

major volcanic eruptions is consistent with the competing 

radiative and chemical effects of volcanic eruptions on 

stratospheric temperature and ozone. This conclusion was 

supported by a modeling exercise (Dall’Amico et al., 

2010a) showing that if observed ozone values are prescribed in their model instead of the assumption of a “simple” linear ozone trend, the stepwise nature of the stratospheric cooling is better reproduced.

Thompson and Solomon (2009) also demonstrated 

that the contrasting latitudinal structures of recent stratospheric temperature (i.e., stronger cooling in the tropical 

lower stratosphere than in the extratropics) and ozone 

trends (i.e., enhanced ozone reduction in the tropical lower 

stratosphere) are consistent with the assumption of increases in the stratospheric overturning Brewer-Dobson circulation. The seasonality of tropical lower-stratospheric tem-
4.18

Chapter 4

1960 1970 1980 1990 2000

0

2

0

2

0

2

0

2

0

2

0

2

MSU4

SSU25

SSU26x

SSU26

SSU27

SSU36x

CNRM-ACM

E39CA

EMAC

GEOSCCM

AMTRAC3

CAM3.5

CCSRNIES

CMAM

LMDZrepro UMSLIMCAT

UMUKCA-METO

UMUKCA-UCAM

WACCM

Year

MRI

NIWA_SOCOL

SOCOL

ULAQ MSU/SSU

Temperature Anomaly (K)

Figure 4-10. Global mean time series of satellite observed (Randel et al., 2009) and chemistry-climate model 

temperature anomalies (K) weighted for MSU/SSU weighting functions. The anomalies are calculated with 

respect to the period 1980–1994, as in the provided MSU/SSU anomalies. Exact weighting functions for the 

MSU/SSU satellite instruments can be found in Figure 1 of Randel et al. (2009). Channel 27 corresponds to 

~34–52 km altitude, channel 36x to ~38–52 km, channel 26 to ~26–46 km, channel 25 to ~20–38 km, channel 

26x to ~21–39 km, and channel MSU4 to ~13–22 km. See Table 3-1 in Chapter 3 of this Assessment for a 

description of the chemistry-climate models shown in the figure.
4.19

Stratospheric Changes and Climate

perature trends using the MSU4 (~13–22 km altitude) data 

(1979–2007) was examined by Fu et al. (2010). Evidence 

was given that the seasonality is a response to changes in 

the Brewer-Dobson circulation, as discussed further in Sections 4.2.2 and 4.3.2. Based on the same data set but focusing on the Southern Hemisphere, Lin et al. (2009) showed 

that the observed temperature patterns are characterized by 

cooling and warming regions of similar magnitudes, with 

strongest local trends occurring in austral spring. Canziani 

et al. (2008) found that the spring warming observed in the 

Southern Hemisphere stratosphere could be linked to an enhanced stationary planetary wavenumber 1 and increased 

total ozone associated with this wave at midlatitudes.

The results presented in Figure 4-10 show a better 

overall agreement between observations and chemistryclimate model results than in previous comparison studies 

(Eyring et al., 2006; WMO, 2007, Chapter 5). Global 

mean temperature anomalies, particularly the longer-term 

trend, derived from observations are reproduced well by 

most chemistry-climate model simulation results (Chapter 

3 of SPARC CCMVal, 2010). On the one hand, this may 

be due to improved chemistry-climate models; on the other hand, this may be due to improvements in the retrievals 

of the observed temperature data. Nevertheless, there are 

still some differences in detail between observations and 

chemistry-climate model results in specific regions of the 

atmosphere, e.g., in the Arctic and Antarctic atmosphere 

(Austin et al., 2009). Since regional trends could be affected by various processes, possible reasons for such 

discrepancies can be manifold and need more specific 

investigations.

Figure 4-10 illustrates that many chemistry-climate 

models capture the leveling of the stratospheric temperature since the late 1990s. The impact of the prescribed 

sea surface temperatures is apparent as MSU4 temperatures (centered ~13–22 km) and chemistry-climate model 

temperatures are particularly well correlated compared to 

other levels. A disagreement between the models and observations is clearly seen in Stratospheric Sounding Unit 

channel 26 (SSU26; centered near ~26–46 km) over the 

last decade. SSU26 has a maximum weight at about 5 

hPa and a considerable contribution from the lower stratosphere. Although the agreement is better in Stratospheric 

Sounding Unit channel 27 (SSU27; centered near ~34–52 

km), which peaks at 2 hPa with less contribution from the 

lower stratosphere, the timing and duration of the flattening is quite different between models and observations. 

So far the reasons for these findings are not understood.

The corresponding vertical profiles of global temperature trends derived from MSU/Stratospheric Sounding Unit data and chemistry-climate models are shown in 

Figure 4-11. The observed stratospheric trend between 

1980 and 2005 is connected with emissions of CO2 and 

ODSs (Shepherd and Jonsson 2008; Gillett et al., 2010) 

and is controlled radiatively by changes in CO2, O3, and 

H2O concentrations (Shine et al., 2003). Jonsson et al. 

(2009) found that CO2-induced cooling reduced upperstratospheric ozone depletion by a factor of approximately 

20% over the period from 1975–1995 (when the amount 

of stratospheric ODSs was increasing most rapidly) when 

compared with what would have been expected based on 

the ODS increases alone. This may not be true in the future as climate-ozone interaction becomes more important 

(see Section 4.5).

Nearly all chemistry-climate models successfully 

model the broad features of the observed temperature 

trend, with warming in the troposphere and cooling in 

Model Temperature Trend (K/decade)

Annual Mean Trend

1980-2005, 70°N-70°S

Figure 4-11. Near global (70°S–70°N) and annual 

mean trends for 1980-2005 (K/decade) for temperature of CCMVal REF-B1 model simulations. The figure includes satellite observed MSU/SSU trends and 

95% confidence intervals. MSU/SSU data points include channels: MSU-4 (at 70 hPa), SSU25 (15 hPa), 

SSU26 (5 hPa), SSU27 (2 hPa), SSU15X (45 hPa), 

SSU26X (15 hPa), and SSU36X (1 hPa), where the 

specified pressure levels represent the approximate 

weighted mean heights derived from the MSU/SSU 

vertical weighting functions for each channel (see 

Randel et al., 2009). The solid black lines indicate 

the multi-model mean results. See Table 3-1 in 

Chapter 3 of this Assessment for a description of the 

chemistry-climate models shown in the figure. From 

Figure 3.2a from SPARC CCMVal (2010).
4.20

Chapter 4

20°S-20°N

ERAI

JRA25

NCEP2

NCEP

ERA40

WACCM

UMUKCA-METO

UMUKCA-UCAM

UMSLIMCAT

UMETRAC

ULAQ

SOCOL

NIWA_SOCOL

MRI

LMDZrepro

GEOSCCM

E39CA

EMAC

CNRM-ACM

CMAM

CCSRNIES

CAM3.5

AMTRAC3

MEAN

Figure 4-12. Annual cycle of tropical (20°S–20°N) cold point tropopause temperature (K) from chemistry-climate models and observations. Model output and observations are from the period 1980–1999. Gray shaded region is 3σ variability from ERA40 analyses. Reanalysis systems in brown with different line styles: European Centre for Medium-Range 

Weather Forecasts (ECMWF) 40-year reanalysis (ERA40; solid), ECMWF Interim Re-Analysis (ERAI; short dash), 

Japanese Re-Analysis (JRA25; dot dash), National Centers for Environmental Prediction (NCEP; dotted), and NCEP2 

(long dashed). The multi-model mean (MEAN) is the thick black line. See Table 3-1 in Chapter 3 of this Assessment 

for a description of the chemistry-climate models shown in the figure. From Gettelman et al. (2010).

HALOE

WACCM

UMETRAC

UMSLIMCAT

ULAQ

SOCOL

NIWA_SOCOL

MRI

LMDZrepro

GEOSCCM

E39CA

EMAC

CNRM-ACM

CMAM

CCSRNIES

CAM3.5

AMTRAC3

MEAN

Water Vapor, 20°S-20°N, 80hPa

Figure 4-13. Annual cycle of tropical (20°S–20°N) water vapor mixing ratio at 80 hPa from chemistry-climate models and observations. Model output from the period 1992–2004. Gray shaded region is 3σ variability from HALOE 

observations over 1992–2004 (thick brown dashed line). The multi-model mean (MEAN) is the thick black line. See 

Table 3-1 in Chapter 3 of this Assessment for a description of the chemistry-climate models shown in the figure. From 

Gettelman et al. (2010).
4.21

Stratospheric Changes and Climate

the stratosphere over this period. It should be noted that 

the tropospheric warming is by construction constrained 

towards observations because the historically observed 

sea surface temperatures are prescribed in most of the 

chemistry-climate models. In the stratosphere, most models capture the overall behavior of the vertical variations, 

with cooling maxima in the upper and lower stratosphere 

(Chapter 3 of SPARC CCMVal, 2010). This indicates that 

the natural and anthropogenic forcings considered in the 

chemistry-climate models seem to be sufficient to explain 

the global-mean temperature changes in the stratosphere 

over the last 50 years. Nevertheless, specific model results indicate that there are still several uncertainties.

For example, Gettelman et al. (2010) analyzed the 

annual cycle of the tropical cold point tropopause temperature, i.e., the region of lowest temperatures, comparing the 

results of chemistry-climate models (i.e., using the REF-B1 

CCMVal-2 model fields; see footnote 1 in this section) 

with corresponding data derived from several reanalysis 

systems (Figure 4-12). Although almost all chemistryclimate models are able to reproduce the annual cycle, 

there are significant offsets between some models and observations. Interestingly, the multi-model mean is very 

close to ERA-40, and the model results presented in Figure 

4-12 are generally in better agreement with observations 

than results from the chemistry-climate models used in the 

first set of CCMVal (CCMVal-1) experiments (Gettelman 

et al., 2009). The uncertainties in the chemistry-climate 

models’ tropical cold point tropopause temperature affect 

the water vapor content of air entering the stratosphere 

(Figure 4-13), which is critical for the chemistry and the 

climate of the stratosphere (see also Section 4.1.3).

Other examples of model uncertainties were presented by Austin et al. (2008), who analyzed multidecadal 

simulations of chemistry-climate models concerning the 

presence of the 11-year solar activity cycle in ozone and 

temperature data, and compared it with respective satellite 

measurements. They showed that the currently available 

model systems have improved significantly compared 

to studies that were based on the previous generation of 

chemistry-climate models (e.g., Soukharev and Hood, 

2006); model results and observations mostly agree, 

within stated uncertainties, with regard to the vertical 

structure of the ozone solar response, particularly in the 

tropics where a double vertical peak structure is now reproduced by chemistry-climate models with a minimum 

near 20 hPa, which is slightly below the minimum identified in observations (~10 hPa). The tropical temperature 

solar response calculated by the chemistry-climate models 

(Figure 4-14) also shows a “double peak,” whereas temperature data derived from satellite based measurements 

(Stratospheric Sounding Unit and MSU) do not indicate 

such a vertical response pattern. However, it is likely that 

the Stratospheric Sounding Unit has insufficient vertical 

resolution to identify such a double peak structure (Gray 

et al., 2009).

Another uncertainty with regard to the impact of 

solar variability arises from recent measurements (April 

2004–March 2008) of solar spectral irradiance acquired 

by the Solar Radiation and Climate Experiment satellite 

(Harder et al., 2009). While these observations show 

variations in total irradiance largely in line with previous solar activity cycles, they cast doubt on currently accepted models of the spectral composition of the irradiance variability. In particular they suggest much larger 

(by approximately a factor of six) differences at ultraviolet 

wavelengths, and negative changes (i.e., greater emissions 

at solar minimum relative to solar maximum) in the visible and near infrared. These measurements are currently 

available only over a four-year period during the final declining stages of the most recent solar cycle and therefore 

remain to be validated; however, if correct, they imply a 

much larger solar signal in stratospheric ozone (Haigh et 

al., 2010). The magnitudes of experimental uncertainties 

in both the irradiance and contemporaneous ozone data do 

not yet allow this to be tested.

Trends in stratospheric temperature could also be 

affected by changes in transport of radiatively active trace 

Observations

WACCM

MRIV2

AMTRAC

0.3

1.0

3.0

10.0

30.0

100.0

0.0 0.5 1.0 1.5 2.0

Pressure (hPa)

K/100 units F10.7

Figure 4-14. Temperature solar response averaged 

over the latitude range 25°S to 25°N. Units on x-axis 

in Kelvin per 100 units (in 10−22 W m−2 Hz−1) of the 

solar radio flux at 10.7 cm (F10.7). The figure shows 

the results from ensemble simulations for models 

that consider the 11-year solar activity cycle. The 

analyzed periods are: 1960–2004 (AMTRAC), 1980–

2004 (MRI), and 1950–2003 (WACCM). The solar 

cycle derived from SSU and MSU data is indicated 

by the dotted black line. Data derived from observations (1979–1997) were reprocessed from Scaife at 

al. (2000b). (From Figure 12 of Austin et al., 2008.)
4.22

Chapter 4

gases via changes in stratospheric circulation that might 

be induced by climate change (e.g., Cook and Roscoe, 

2009). As noted in Section 4.3.2, the stratospheric BrewerDobson circulation is likely to be modified in the presence of enhanced greenhouse gas concentrations. Certainly any change in the strength of the Brewer-Dobson 

circulation would alter the thermal structure of the stratosphere. Moreover, it must be considered that an intensified Brewer-Dobson circulation would affect stratospheretroposphere mass transport and lead to a quicker turnover 

time of stratospheric air, i.e., lower age of stratospheric air.

The simulated mean temperature changes (i.e., 

multi-model mean values) from 1979 to 2000 derived 

from a subset of the CMIP32 coupled ocean-atmosphere 

climate model simulations are dominated by zonally symmetric radiative cooling due to ozone depletion. Based 

on this result Lin et al. (2009) concluded that the climate 

models fail to simulate the warming in the southern polar stratosphere (see Figure 4-6), indicating a lack of the 

Brewer-Dobson circulation strengthening in these models. This interpretation contradicts the statements given 

above referring to chemistry-climate model results. A 

reason could be that most of the climate models used do 

not adequately resolve the stratosphere and therefore inadequately reproduce the Brewer-Dobson circulation and its 

changes. On the other hand it is still questionable if the recent (1995–2008) warming of the polar lower stratosphere 

is a robust and statistically significant feature. Gravity 

wave parameterizations within the climate models may 

also contribute to these differences.

The stratospheric temperature trends simulated 

by the large ensemble of climate models available from 

the CMIP3 experiments have been evaluated in the IPCC 

Fourth Assessment Report, although the main emphasis 

there was on attribution of climate variations in the troposphere (Hegerl et al., 2007 in IPCC, 2007). Those climate 

models do not generally include stratospheric chemical 

processes but some include specified changes in stratospheric ozone concentrations. For example, Schwarzkopf 

and Ramaswamy (2008) analyzed simulations of a CMIP3 

climate model (GFDL-CM2.1) for the period from 1861 

to 2003. They showed that the global-mean cooling in 

the stratosphere (mainly due to enhanced CO2 concentrations) became significant as early as the first quarter of 

the 20th century. Thereafter the magnitude of stratospheric 

cooling by mid-20th century at ~20–50 hPa exceeded the 

magnitude of the warming near the surface. By the late 

1970s, significant stratospheric ozone depletion had begun, enhancing the cooling.

2 CMIP3 refers to simulations of 20th and 21st century climate 

organized by the Coupled Model Intercomparison Project Phase 3 

for use in IPCC experiments (Meehl et al., 2007).

The observed global mean temperature trends in the 

troposphere and lower stratosphere for recent decades are 

simulated well by most chemistry-climate models (e.g., 

Figure 4-10), but are not generally as well simulated by 

some CMIP3 climate models (Figure 4-15; from Cordero 

and Forster, 2006). The observed global mean cooling 

trend in the lower stratosphere is grossly underestimated by 

climate models that do not include specified decreases in 

stratospheric ozone over the last three decades. Even climate models that include specified ozone decreases slightly 

underestimate the observed cooling trend (for details see 

Cordero and Forster, 2006). In addition, climate models 

that include specified increases in stratospheric volcanic 

aerosols appear to overestimate the warming of the lower 

stratosphere after major volcanic eruptions (Figure 4-15), 

possibly because they neglect the ozone changes associated 

with the eruption and do not include the influence of the 

stratospheric quasi-biennial oscillation (Dall’Amico et al., 

2010a). Moreover, Waugh et al. (2009) showed that the 

calculated cooling simulated in response to ozone depletion is larger in a model with interactive chemistry (i.e., a 

chemistry-climate model) than that determined by the same 

underlying atmospheric general circulation model that prescribes the identical zonal mean ozone forcing.

Climate models have also been used to show that 

tropopause height variations are a sensitive indicator of 

temperature changes in the upper troposphere and lower 

stratosphere. They have enabled attribution of observed 

trends in mean tropopause height over the last three decades of the last century to the combined influence of increasing concentrations of greenhouse gases and aerosols 

and decreasing amounts of stratospheric ozone (Santer et 

al., 2003). More recently, Son et al. (2009a) investigated 

results derived from six chemistry-climate models and 

found that historical changes (since 1960) in tropopause 

height are linked to ozone depletion.

4.3.2 Simulation of Brewer-Dobson 

Circulation Trends in ChemistryClimate Models

The Brewer-Dobson circulation is the stratospheric 

wave-driven circulation that transports mass and constituents upward in the tropics and poleward and downward 

at higher latitudes (see Section 4.2.2 for a more detailed 

description). In the 2006 Assessment (WMO, 2007), 

chemistry-climate models were found to consistently predict an acceleration of the Brewer-Dobson circulation in 

response to increasing greenhouse gas concentrations, 

amounting to an average increase of about 2%/decade in 

the annual mean net upward mass flux at 70 hPa through 

the 21st century. Butchart et al. (2006) also found that over 

half of the annual mean trend was explained by changes 
4.23

Stratospheric Changes and Climate

in resolved wave drag, with the remainder assumed to be 

due to parameterized wave drag. They were unable to 

conclude whether changes in stratospheric ozone had any 

effect on the predicted trend in upward mass flux.

Subsequent to the previous Assessment (WMO, 

2007) our understanding of the mechanisms driving the 

Brewer-Dobson circulation in simulations and its predicted future increase has advanced considerably. Chemistryclimate model intercomparison studies (Butchart et al., 

2010; Chapter 4 of SPARC CCMVal, 2010) have shown 

that the ~2%/decade increase is a robust value for the lower stratosphere (Figure 4-16) in model simulations driven 

by comparable scenarios. Using a better set of model simulations (i.e., only transient simulations with common 

forcings using more mature models) and a more accurate 

calculation of the mass fluxes, Butchart et al. (2010) found 

that ~70% of the upward mass flux trend was due to parameterized orographic gravity wave drag, a value that was 

computed from only the subset of models that provided 

the gravity wave drag data. The importance of these 

waves in driving the acceleration of the Brewer-Dobson 

circulation has also been elucidated by a number of studies 

(Li et al., 2008; McLandress and Shepherd, 2009; Butchart 

et al., 2010). These studies have demonstrated that the 

predicted future increase in lower-stratospheric parameterized orographic gravity wave drag is a robust model 

response to climate change, resulting from the eastward 

acceleration of the subtropical jets which increases the 

gravity wave momentum flux reaching the lower stratosphere. While the parameterized momentum fluxes are 

not well constrained by observations, they do respond in a 

physically consistent manner to changes in the large-scale 

circulation. It should also be noted that the similarity of 

the model results might be due in part to the similarity of 

the orographic gravity wave parameterizations.

Another advance has been our understanding of the 

role of resolved wave drag. However, due to a lack of the 

necessary diagnostic data available for model intercomparison studies, only individual model studies have been 

undertaken. In comparing these studies, some caution is 

required since different definitions are used to characterize tropical upwelling. Using the net upward mass flux 

at 70hPa, McLandress and Shepherd (2009) examined the 

contributions from different zonal wavenumbers. They 

found that 60% of the net upward mass flux trend due to 

resolved wave drag resulted from planetary waves, with 

the remainder coming from synoptic waves. On the other 

hand, Garcia and Randel (2008) and Calvo and Garcia 

(2009) employed a fixed latitude range that did not encompass the subtropics to define tropical upwelling. In 

the former study the trend in tropical upwelling resulted 

from changes in resolved wave drag in the subtropical 

1950 1960 1970 1980 1990 2000

Year

-2

-1

0

1

2

3

4

Temperature Anomaly (K)

AOGCMs with fixed ozone

AOGCMs with ozone changes

Radiosonde climatology

Figure 4-15. Time series of 

globally averaged annual temperature anomalies (K) at 50 hPa 

from radiosonde observations 

(black lines) and simulations of 

atmosphere-ocean general circulation models (AOGCMs; colored lines). Climate model simulations that included specified 

stratospheric ozone depletion, 

as well as increases in greenhouse gases and anthropogenic 

aerosols, are shown with blue 

lines, while model simulations 

that did not include stratospheric 

ozone depletion are shown with 

red dashed lines. Some, but 

not all, the models also include 

specified stratospheric volcanic 

aerosol amounts. Short vertical 

lines on the abscissa denote 

years of major volcanic eruptions. Figure modified from Cordero and Forster (2006).
4.24

Chapter 4

lower stratosphere, which in turn resulted from enhanced 

wave propagation due to changes in the zonal mean zonal 

winds. Calvo and Garcia (2009) concluded that enhanced 

dissipation of quasi-stationary planetary waves forced by 

increased tropical convection were responsible for the upwelling trends, at least in a fixed latitude band. A similar 

conclusion about the role of convectively forced stationary waves was made by Deckert and Dameris (2008), although they did not quantify the impact of those waves on 

upwelling trends. 

Consistent with the predicted increase in net upward 

mass flux, the mean age of stratospheric air is predicted to 

decrease (Austin and Li, 2006; Garcia and Randel, 2008; 

Oman et al., 2009; Butchart et al., 2010). Oman et al. 

(2009) examined the impact of a number of different factors affecting mean age and concluded that their impact 

depended upon the time period examined. Over the past 40 

years they found that ozone depletion was responsible for 

the simulated decrease in mean age, while later in this century after ozone has returned to 1980 levels, decreasing age 

was attributed to increasing greenhouse gas concentrations.

Another advance has been our understanding of the 

role of ozone depletion and recovery in the simulated net 

upward mass flux trends. Li et al. (2008) found that in the 

past nearly half of the trend in their simulations occurred 

in December–February. They attributed this to increased 

net downward mass flux in the Southern Hemisphere resulting from the delay in breakdown of the Antarctic polar 

vortex, which allowed Rossby waves to propagate higher 

into the summer stratosphere, thus driving the increased 

downwelling. The impact of ozone depletion (and recovery) on the upward mass flux trends was confirmed in 

McLandress et al. (2010) using a set of chemistry-climate 

model simulations in which greenhouse gas concentrations were held fixed in time and only ODSs were allowed 

to vary.

The role of parameterized orographic gravity wave 

drag in model simulations is better understood than in the 

previous Assessment (WMO, 2007). However, much 

more needs to be done to validate such parameterizations, particularly since they play a key role in the simulated accelerations of the Brewer-Dobson circulation. 

Mass Flux (1

09 kg 

s-1)

Year

1960 1980 2000 2020 2040 2060 2080 2100

10

9

8

7

6

5

4

CMAM

MRI

ULAQ

UMUKCA-UCAM

WACCM

CNRM-ACM

GEOSCCM

CCSRNIES

UMUKCA-METO

SOCOL

Figure 4-16. Annual mean upward mass flux (kg s-1) at 70 hPa, calculated from residual mean vertical velocity 

between the turnaround latitudes of the Brewer-Dobson circulation, for the simulations prepared for the CCMVal SPARC report. See Table 3-1 in Chapter 3 of this Assessment for a description of the chemistry-climate 

models shown in the figure. Figure from Chapter 4 of SPARC CCMVal (2010).
4.25

Stratospheric Changes and Climate

Various prescriptions exist for parameterizing the influences of unresolved gravity waves, but they are based on 

specifications of certain parameters that are not well constrained by observations at the present time. As long as 

chemistry-climate models are strongly dependent on the 

parametrization of gravity waves, the simulated changes 

in the Brewer-Dobson circulation will remain similarly 

dependent on the model treatment of such unresolved 

physical processes.

The mechanisms responsible for the increased 

resolved wave drag are also unclear. As discussed in 

McLandress and Shepherd (2009), differences in the way 

tropical upwelling is computed have made it difficult to 

come to a clear consensus on which types of resolved 

waves are responsible for the increase in the BrewerDobson circulation. Quantifying whether the increased 

resolved wave driving of the Brewer-Dobson circulation 

results from changes in wave sources in the troposphere 

or wave propagation in the stratosphere is an outstanding 

issue.

4.4 EFFECTS OF VARIATIONS IN 

STRATOSPHERIC CLIMATE ON THE 

TROPOSPHERE AND SURFACE

In this section we assess the linkages between 

stratospheric variability and climate change in the troposphere and at the Earth’s surface. The section includes an 

assessment of radiative, dynamical, and chemical coupling 

between the stratosphere and troposphere, but focuses 

primarily on the dynamical effects of the Antarctic ozone 

hole on surface climate. The evidence for a robust linkage between the Antarctic ozone hole and the surface flow 

in the Southern Hemisphere was assessed in Chapter 5 of 

WMO, 2007 (Baldwin and Dameris et al., 2007). Here we 

extend that assessment to include recent advances in our 

understanding of the linkages between the Antarctic ozone 

hole and surface climate change throughout the Southern 

Hemisphere. Key advances include a deeper understanding of the likely effects of Antarctic ozone depletion on 

surface weather, the circulation of the Southern Ocean, the 

distribution of Antarctic sea ice, and the Southern Hemisphere carbon cycle.

The primary novel effects of stratospheric climate 

variability on surface climate assessed in this section are 

summarized in Table 4-1. The text in Section 4.4 is organized into five components as follows:

Section 4.4.1, stratospheric composition and globalmean surface temperatures

Section 4.4.2, the Antarctic ozone hole and changes in 

Southern Hemisphere surface climate

Section 4.4.3, stratospheric climate change and the 

tropical tropospheric circulation

Section 4.4.4, stratospheric variability and tropospheric chemistry

Section 4.4.5, solar-induced variability at stratospheric levels and tropospheric climate

4.4.1 Effects of Stratospheric Composition 

Changes on Global-Mean Surface 

Temperature and Tropospheric

Temperature

Previous World Meteorological Organization 

(WMO) Ozone Assessments (e.g., WMO, 2007) and the 

recent IPCC Assessment (IPCC, 2007) have evaluated radiative forcings associated with changes in stratospheric 

climate. Radiative forcings give an indication of surface 

temperature change, whereby positive forcings enhance 

the global mean surface temperature and negative forcings 

tend to cool the surface. Stratospheric ozone was most 

recently assessed to give a net radiative forcing of −0.05 ± 

0.1 W/m2 between 1750 and 2005 (Forster et al., 2007a). 

Water vapor concentration changes associated with methane oxidation between 1750 and 2005 were assessed to 

contribute a forcing of +0.07 ± 0.05 W/m2. The amount of 

stratospheric aerosol following volcanic eruptions can exert a strongly negative forcing on a short timescale, estimated to be roughly −3 W/m2

 following the eruption of 

Mt. Pinatubo. Note that volcanic eruptions can also have 

indirect forcings by affecting concentrations of ozone (see 

Chapter 3) and stratospheric water vapor (e.g., Joshi and 

Jones, 2009). The direct forcing of solar changes since 

1750 is estimated to be small, although uncertainties exist 

in long-term trends in solar irradiance, in variations in its 

spectral composition, and also in its indirect effects on 

stratospheric ozone.

Figure 4-17 shows the ozone forcing since the 

1970s evaluated from chemistry-climate models and observations using a single radiation model from Chapter 

10 of SPARC CCMVal (2010). The 1970s-to-2004 radiative forcing is estimated to be −0.03 ± 0.2 W/m2

 (90% 

confidence range) from stratospheric ozone changes in 17 

chemistry-climate model runs and +0.03 W/m2 employing 

Randel and Wu (2007) observations. Negative net radiative forcings arise from models with ozone decline in the 

lowermost stratosphere, particularly at or near the tropopause. This stratospheric ozone forcing is not entirely of 

anthropogenic origin, but includes components due to the 

indirect effect of volcanic eruptions and solar irradiance 

change. The large spread is possibly due to an incorrect 

ozone response to volcanic eruptions in certain chemistryclimate models (Chapter 10 of SPARC CCMVal, 2010). 

Including the uncertainty in radiative transfer modeling 

would be unlikely to increase the uncertainty range (Chapter 3 of SPARC CCMVal, 2010). As differences between 
4.26

Chapter 4

Affected 

Tropospheric 

Climate Parameters

Currently Observed Effects Currently Understood 

Causes

Level of 

Scientific 

Understanding

1. Global radiative 

balance 

affecting global 

mean surface 

temperature 

(Section 4.4.1)

•	 Slight negative long-term radiative forcing 

from stratospheric ozone variations

•	 Slight positive long-term radiative forcing 

from stratospheric water vapor from 

methane oxidation

•	 Large decadal fluctuations in radiative 

forcing from stratospheric water vapor 

Stratospheric ozone trends, 

stratospheric water vapor 

trends, stratospheric aerosol 

trends, and other greenhouse 

gas trends affect stratospheric 

temperatures. These affect 

both the shortwave and longwave components of the radiative budget.

Very Strong

2. Tropical 

tropospheric 

temperature 

(Section 4.4.1)

Tropical upper-tropospheric cooling trend

Reduced tropical stratospheric 

ozone cools stratospheric 

temperatures. This reduces 

longwave radiative flux into 

the tropical troposphere.

Medium

3. Antarctic 

tropospheric 

temperature 

(Section 4.4.1)

Late spring and summer Antarctic

tropospheric cooling

ODS-induced stratospheric 

ozone loss cools Antarctic 

stratosphere. This reduces 

longwave radiative flux into 

the Antarctic troposphere.

Medium-Strong

4. Southern Annular 

Mode (SAM) 

(Section 4.4.2.1)

Trend to positive index of SAM in Southern 

Hemisphere summer. Evident as falls in geopotential height over the pole and an eastward 

acceleration of the surface winds over the 

Southern Ocean.

ODS-induced stratospheric 

ozone loss cools the Antarctic 

stratosphere. This induces 

a stratospheric dynamical 

response that couples to the 

tropospheric SAM.

Strong

5. Southern 

Hemisphere 

extratropical 

winds, storm 

tracks, and 

precipitation 

(Section 4.4.2.2)

•	 Poleward shift of Southern Hemisphere 

extratropical jet in summer

•	 Reduced low pressure systems at low 

latitude and increased at high latitudes

•	 Poleward shift of Southern Hemisphere 

precipitation

Effects 5–9 are all related to 

the summertime tropospheric 

SAM trend (climate parameter 

4), which is caused by ODSinduced Antarctic stratospheric cooling.

Medium

6. Antarctic surface 

temperatures 

(Section 4.4.2.3)

•	 Summertime warming of Antarctic 

Peninsula and cooling trend of East 

Antarctic High Plateau

Medium-Strong

7. Antarctic sea ice 

(Section 4.4.2.3)

•	 Increase in summertime Antarctic sea ice 

extent Medium

8. Southern Ocean 

temperatures 

and circulation 

(Section 4.4.2.4)

•	 Warming of Southern Ocean subsurface 

and poleward shift of Southern Ocean 

density structure

Low-Medium

9. Southern Ocean 

carbon cycle 

(Section 4.4.2.4)

•	 Reduction in Southern Ocean carbon 

uptake Low-Medium

10. Width of tropical 

belt (Section 4.4.3)

Widening of tropical belt, with accompanying 

poleward shift of the jet streams, the downward 

branches of the Hadley Cell, and the region of 

high tropical tropopause

Attributed both to ODSinduced Antarctic ozone 

depletion and greenhouse 

warming.

Low

Table 4-1. Highlights of Section 4.4: How stratospheric climate variations affect the troposphere and 

surface.
4.27

Stratospheric Changes and Climate

Global mean solar (shortwave) forcing

Radiative Forcing (W/m2)

Global mean thermal infrared (longwave) forcing

Global mean net forcing

1960 1970 1980 1990 2000

Year

Radiative Forcing (W/m2) Radiative Forcing (W/m2)

Year

1960 1970 1980 1990 2000

1960 1970 1980 1990 2000

Year

0.4

0.2

0.0

-0.2

-0.4

-0.6

0.4

0.2

0.0

-0.2

-0.4

-0.6

0.4

0.2

0.0

-0.2

-0.4

-0.6

AMTRAC3

CMAM

EMAC

MRI

ULAQ

UMUKCA-UCAM

CAM3.5

CNRM-ACM

GEOSCCM

NIWA_SOCOL

UMSLIMCAT

WACCM

CCSRNIES

E39CA

LMDZrepro

SOCOL

UMUKCA-METO

Randel&Wu

Figure 4-17. Stratospheric ozone global-mean radiative forcing (W/m2) for shortwave (top), longwave (middle) 

and net (shortwave plus longwave) compared to the 1970s average, evaluated from the CCMVal REF-B1 scenario ozone fields (colored lines) and from the Randel and Wu (2007) observation-based data set (thick black 

line). See Table 3-1 in Chapter 3 of this Assessment for a description of the chemistry-climate models shown 

in the figure. Based on Chapter 10 of SPARC CCMVal (2010).
4.28

Chapter 4

the SPARC CCMVal calculations and the IPCC assessment of the stratospheric ozone forcing are small, the 

IPCC assessment of this forcing is retained.

Decadal radiative effects of stratospheric composition changes are significant when compared to that from 

carbon dioxide, whose forcing increased by roughly 0.2 

W/m2 per decade since 1980 (Forster et al., 2007b). Solomon et al. (2010) estimated that the drop in stratospheric 

water vapor levels near the tropopause following 2000 (see 

Section 4.1.3) contributed a forcing of −0.1 W/m2. They 

show that decadal variation of stratospheric water vapor 

concentrations likely contributed to an enhanced rate of 

surface warming in the 1990s (see also Forster and Shine, 

2002) and a reduced rate of surface warming in the 2000s 

(see Figure 4-18).

Volcanic aerosols directly affect surface climate by 

reflecting solar radiation to space. Major volcanic eruptions cool global average surface temperatures for a few 

years following an eruption (Hansen et al., 1992; Robock, 

2000), and also influence the patterns of Northern Hemisphere surface climate through effects on the Northern Annular Mode (e.g., Shindell et al., 2004; Stenchikov et al., 

2006). Several studies have underscored that a major volcanic eruption not only perturbs the atmosphere but also 

cools the oceans, thereby affecting ocean heat uptake, sea 

level, and climate for many decades (Church et al., 2005; 

Delworth et al., 2005; Gleckler et al., 2006; Domingues et 

al., 2008; Gregory et al., 2006; Stenchikov et al., 2009). 

Thus the timescale over which volcanic effects need to be 

considered in climate studies is much longer than the time 

over which the aerosols are present in the stratosphere, and 

is controlled by multiple factors, including not only the 

stratospheric timescales but also the slow transport time 

within the deep ocean.

Recent studies have shown that the changes in radiative forcing from changes in stratospheric composition manifest themselves by directly changing temperatures in the 

free troposphere. This is in addition to any surface-mediated 

temperature change. Parts of the observed tropospheric temperature trend in both the tropics and Southern Hemisphere 

polar regions have been linked to stratospheric ozone depletion and the associated stratospheric cooling trends (Forster 

et al., 2007b; Keeley et al., 2007; Grise et al., 2009).

In the tropics, observed cooling trends over 1985–

2005 above 12 km in the upper troposphere do not appear to be linked to the local radiative impacts of ozone 

depletion in the upper troposphere but rather to the reduced downwelling longwave radiation associated with 

ozone depletion in the lower stratosphere (Forster et al., 

2007b). The observed tropical upper-tropospheric temperature trends above 12 km are not well reproduced in 

climate model simulations, especially those with simple 

ozone prescriptions, suggesting the importance of properly prescribing tropical lower-stratospheric ozone in future 

climate modeling studies of the tropical upper troposphere 

(Forster et al., 2007b; Allen and Sherwood, 2008).

Several recent studies have suggested that the observed summertime cooling trends in the Antarctic free 

troposphere are at least partially attributable to decreases 

in downwelling longwave radiation emitted from the Antarctic lower stratosphere (Keeley et al., 2007; Grise et al., 

2009). The decreases in downwelling longwave radiation 

appear to be driven primarily by the stratospheric temperature changes associated with ozone depletion (Grise 

et al., 2009), particularly the ozone depletion above the 

lowermost stratosphere (Keeley et al., 2007). A caveat to 

the above studies is that they prescribe zonally symmetric stratospheric ozone concentrations for their radiative 

calculations. Zonal asymmetries in prescribed ozone concentrations at Southern Hemisphere high latitudes have 

been shown to cause additional cooling in the stratosphere 

and upper troposphere (e.g., Waugh et al., 2009; Gillett 

et al., 2009), but most of this additional cooling appears 

to be driven by dynamical rather than radiative processes 

(Crook et al., 2008). Dall’Amico et al. (2010b) also show 

the importance of representing global ozone variability 

including the stratospheric quasi-biennial oscillation for 

simulating ozone-climate effects.

4.4.2 Surface Climate Impacts of the 

Antarctic Ozone Hole

As discussed in Section 4.2 and in Chapter 5 of 

the previous Ozone Assessment (Baldwin and Dameris 

Impact of Stratospheric H2O Changes

on Decadal Temperature Trends

0.0

0.1

0.2

0.3

K/decade

LLGHG and Aerosols Only

Plus Increasing Water Trend Through 2000

Plus Decrease in 2001

'80s '90s '80s '90s '00s

'00s

{

Figure 4-18. Decadal warming rates (K/decade) 

arising from (i) the long-lived greenhouse gases 

(LLGHG) and aerosols alone (black), as well as 

(ii) that obtained including the stratospheric water 

decline after 2000 (red) and (iii) including both the 

stratospheric water vapor decline after 2000 and the 

increase in the 1980s and 1990s (cyan). Based on 

Solomon et al. (2010).
4.29

Stratospheric Changes and Climate

et al., 2007), the robust stratospheric cooling in the Antarctic spring associated with the ozone hole is coincident 

with an anomalous eastward3 acceleration of the stratospheric polar vortex, and hence a delay in the springtime 

breakdown of the Southern Hemisphere stratospheric vortex. The trends in the Southern Hemisphere extratropical 

flow are not limited to stratospheric levels, but extend to 

Earth’s surface during the spring/summer months of December and January, when the trends in the atmospheric 

flow are dominated by a poleward shift of the tropospheric zonal winds of the Southern Hemisphere (Thompson 

and Solomon, 2002; WMO, 2007). In WMO (2007), 

this tropospheric response was attributed to the Antarctic 

ozone hole, primarily based on the observational study of 

Thompson and Solomon (2002) and the modeling study of 

Gillett and Thompson (2003). In this section, we assess 

work since WMO (2007) on this stratosphere-troposphere 

connection and its implication for the surface climate of 

high southern latitudes.

In the stratosphere and troposphere, the spatial 

structure of the observed extratropical trends projects 

strongly upon the dominant and most persistent pattern 

of natural variability in the extratropical circulation: the 

Southern Annular Mode (Thompson and Wallace, 2000; 

Thompson and Solomon, 2002; Gillett and Thompson, 

2003)4. In fact, the observational and modeling evidence 

outlined in Chapter 5 of the previous Ozone Assessment 

(Baldwin and Dameris et al., 2007) strongly suggests that 

the dynamical influence of the ozone hole on surface climate is manifested almost entirely in the structure of the 

Southern Annular Mode. However, we emphasize that 

the projection is clearest in the zonal mean and that the 

Southern Annular Mode does not account for all aspects 

of Southern Hemisphere climate trends: the trend in the 

3 When used in the context of the wind, “eastward” (toward the 

east) is synonymous with “westerly” (from the west). We use the 

“eastward” terminology throughout this chapter since the term 

“westerly” has an ambiguous definition depending on what it is 

used to describe. 

4 The Southern Annular Mode (SAM) is the dominant pattern 

of variability in the Southern Hemisphere extratropical circulation (e.g., see Thompson and Wallace 2000 and discussions in 

Chapter 5 of WMO, 2007). Typically, it is defined to be the leading 

empirical orthogonal function (EOF) of the Southern Hemisphere 

geopotential or zonal wind. For an EOF calculation using monthly 

mean time series and using all months of the year, the SAM based 

on (1) zonally varying 850 hPa geopotential, (2) zonal mean 1000-

50 hPa geopotential, and (3) zonal mean 1000-50 hPa zonal wind 

respectively explains (1) 27%, (2) 47%, and (3) 45% of the variance (Table 2 of Thompson and Wallace, 2000). In the troposphere, SAM variability is characterized by meridional excursions 

of the tropospheric jet. During the Southern Hemisphere spring, 

the SAM also corresponds to variations in the strength of the polar 

vortex in the stratosphere. SAM dynamics is understood to reflect 

interactions between waves and the zonal mean flow (e.g., Limpasuvan and Hartmann, 2000).

annular mode is largest during summer (as noted below), 

it does not account for many zonally varying aspects of 

recent Southern Hemisphere climate change, and it may 

not fully account for zonally asymmetric forcing by ozone 

depletion (e.g., Crook et al., 2008).

That the response to the ozone hole is related to a 

persistent and dominant mode of internal variability like 

the Southern Annular Mode is expected based on general 

theoretical grounds (e.g., Leith, 1975; Gerber et al., 2008; 

Ring and Plumb, 2008) and also the fact that the annular 

mode dynamically connects the stratospheric and tropospheric circulations (Thompson and Wallace, 2000). The 

detailed dynamics of this stratosphere-to-troposphere link 

are still being explored; they have been linked to direct driving of the circulation by diabatic heating and changes to the 

eddy-induced stresses, to eddy mean-flow interactions in 

which the mean flow changes feed back onto the transient 

eddy driving field, to barotropic-mode dynamics, and to radiative driving (e.g., see Song and Robinson, 2004; Esler 

et al., 2006; Thompson et al., 2006; Simpson et al., 2009; 

Grise et al., 2009). But at this point, the precise dynamical mechanism(s) by which the tropospheric circulation responds to ozone depletion have not been fully elucidated.

Since WMO (2007), the robustness and seasonality of the tropospheric circulation response to Antarctic 

ozone depletion has been established in numerous climate 

simulations (e.g., Arblaster and Meehl, 2006; Miller et al., 

2006; Cai and Cowan, 2007; Karpechko et al., 2008; Perlwitz et al., 2008; Fogt et al., 2009; Son et al., 2008, 2009b, 

2010; Chapter 10 of SPARC CCMVal, 2010). Figure 4-19

SAM Trends 1957-2005

Trend and Condence Interval (decade−1)

(summer) (fall) (winter) (spring)

DJF MAM JJA SON

Figure 4-19. Seasonal trends over 1957–2005 in the 

SAM index (in units of decade−1) from observations and 

the CMIP3 climate models with (“OZONE”) and without 

(“NO OZONE”) time-varying ozone forcing. The 95% 

confidence intervals reflect the number of time series 

involved in the trend analysis (see Fogt et al. 2009 for 

details). Data taken from Table 3 of Fogt et al. (2009).
4.30

Chapter 4

(based on Table 3 of Fogt et al., 2009) provides an updated 

view of observed and simulated tropospheric circulation 

trends expressed in terms of the Southern Annular Mode. 

The positive trends in the Southern Annular Mode are most 

robust in observations and models during the austral summer and fall seasons (labeled DJF and MAM, respectively). 

During the austral summer season, a statistically significant 

positive Southern Annular Mode response is simulated only 

in those integrations that include the observed Antarctic 

ozone depletion (labeled “OZONE” in the figure). During 

the austral fall season, the simulated trends are half the observed value and occur whether or not ozone forcing is present. In the austral winter (labeled JJA), the observed and 

simulated trends are not significant. In the austral spring 

(labeled SON), there is a poorly understood discrepancy 

between the models and observations: the “OZONE” models produce a positive Southern Annular Mode trend when 

none is observed (Fogt et al., 2009).

The following caveat must be considered when interpreting the results in Figure 4-19: different models are 

included in the “OZONE” and “NO OZONE” sets in the 

figure, so it is difficult to cleanly attribute the Southern 

Annular Mode trend to ozone-related forcing in these simulations. We can begin to assess the separate effects of 

ozone and greenhouse gas forcing from a limited number 

of studies. For example, when ozone and greenhouse gas 

forcings are prescribed separately in the National Center 

for Atmospheric Research Community Climate System 

Model, it is clear that ozone forcing drives the primary 

Southern Annular Mode response in summer (Arblaster 

and Meehl, 2006). In addition, Sigmond et al. (2010) use 

timeslice integrations of the Canadian Middle Atmosphere 

Model to examine the effect of coupling to a dynamical 

ocean and sea ice model on the summertime Southern Annular Mode response to the ozone hole in the absence of 

greenhouse gas forcing. They find that the summertime 

Southern Annular Mode response is similar whether the 

ocean and sea ice state is prescribed or predicted. This 

implies that the ocean and sea ice response to the ozone 

hole have not caused further changes to the surface wind 

response to the ozone hole.

In the following subsections we assess our improved understanding of the surface climate impacts of the 

ozone-forced trends in the tropospheric circulation. We 

discuss how the trends in Figure 4-19 have affected the 

high-latitude climate of the Southern Hemisphere, including surface circulation and precipitation (Section 4.4.2.1), 

surface temperatures and sea ice (Section 4.4.2.2), Southern Ocean temperatures and circulation (Section 4.4.2.3), 

and the Southern Ocean’s contribution to the global carbon cycle (Section 4.4.2.4). Research in this area has advanced rapidly since the previous Assessment. In the absence of Arblaster and Meehl (2006)-type studies in which 

ozone and greenhouse gas forcing are separately applied 

within coupled ocean-atmosphere climate models, we are 

not able to quantitatively assess the net impact of ozone 

forcing on the annual mean forcing of the Southern Annular Mode trends for the broader set of coupled climate 

models. There is an urgent need for such simulations to 

more precisely connect ODS-related ozone depletion to 

Southern Hemisphere climate change.

Climate impacts of Northern Hemisphere ozone 

depletion are not covered in this Assessment, since no 

definitive connection has been made between Northern 

Hemisphere ozone depletion and surface climate (e.g., Figure 4-9; Thompson and Solomon, 2005), despite the presence of pronounced stratosphere-troposphere coupling in 

association with dynamic variability in the stratospheric 

vortex (e.g., Baldwin and Dunkerton, 2001). The lack of 

an ozone-surface climate response in the Northern Hemisphere is consistent with the weak ozone losses observed 

there (Chapter 2; Solomon et al., 2007) and thus the weak 

radiative forcing and temperature trends in the Arctic stratosphere that are masked by relatively high natural variability 

in some seasons (Grise et al., 2009; Section 4.2).

4.4.2.1 Effects on Winds, Storm Tracks, and

Precipitation

The austral summer trends in the Southern Annular 

Mode and strengthening of the eastward winds over the 

Southern Ocean are associated with a poleward shift of the 

extratropical jets in the Southern Hemisphere troposphere 

in summer (Gillett and Thompson, 2003; Archer and Caldeira, 2008). As would be expected from such a shift, 

observational studies reveal reductions in the number of 

low pressure systems and extratropical storms in Southern 

Hemisphere middle latitudes (e.g., Simmonds and Keay, 

2000; Rao et al., 2003; Vera, 2003; Bengtsson et al., 2006; 

Pezza et al., 2007, 2008) and increases at high latitudes 

(e.g., Lynch et al., 2006). These changes in the storm 

tracks are apparent in both winter and summer. There are 

some suggestions of increases in the intensity of storms in 

the storm track (Simmonds and Keay, 2000) but a more 

recent analysis suggests that this result might be an artifact 

of changes in the available input data for the reanalyses 

considered (Bengtsson et al., 2006).

Onshore winds and extratropical weather systems 

are the causes of much of the precipitation over land in 

Southern Hemisphere middle latitudes. Hence, the changes in the Southern Annular Mode and the Southern Hemisphere storm track are expected to have led to changes in 

precipitation. Many recent studies have considered observed changes in regional precipitation over the past few 

decades and links to changes in Southern Hemisphere 

circulation. They have tried to separate the influences of 

the Southern Annular Mode from those due to decadal 

variations in El Niño-Southern Oscillation, the two major 
4.31

Stratospheric Changes and Climate

factors affecting circulation variations in the Southern 

Hemisphere middle latitudes. No studies have directly 

attributed observed regional precipitation changes in summer to stratospheric ozone depletion yet, probably due to 

the large variability of regional precipitation and the sparse 

network of observing stations. However, a recent analysis 

of the CMIP3 ensemble of climate model simulations 

(CMIP3 is defined in Footnote 2 in Section 4.3.1) suggests 

a signal due to stratospheric ozone depletion in the multimodel average that includes increases in summer precipitation in the Southern Hemisphere high latitudes and 

decreases in middle latitudes due to stratospheric ozone 

depletion (Son et al., 2009b). This signal is likely to have 

contributed to observed changes in precipitation but might 

not be apparent relative to natural variability.

During the summer season, the positive polarity 

of the Southern Annular Mode is linked to increased precipitation in southeastern Australia (Gillett et al., 2006; 

Hendon et al., 2007; Meneghini et al., 2007) and South 

Africa (Gillett et al., 2006) and reduced precipitation in 

southwest New Zealand (Renwick and Thompson, 2006; 

Griffiths, 2006; Ummenhofer et al., 2009) and southwest 

South America (Gillett et al., 2006).

The increase in wind speeds over the Southern 

Ocean can generate more sea spray and lead to increases 

in natural cloud condensation nuclei for the formation of 

reflective low clouds in Southern Hemisphere summer 

(Korhonen et al., 2010). In a modeling study using observed wind trends input into an global aerosol model, 

Korhonen et al. (2010) find that the Southern Hemisphere 

wind trends since the 1980s give rise to a more than 20% 

increase in cloud condensation nuclei concentrations in 

the 50°S–65°S latitude band and a negative cloud radia- °S–65°S latitude band and a negative cloud radia- –65°S latitude band and a negative cloud radia- °S latitude band and a negative cloud radiative forcing of −0.7 W/m2

. The latter, when added to the 

negative radiative forcing from stratospheric ozone loss, 

gives an ozone-related regional negative radiative forcing 

comparable to the positive radiative forcing from greenhouse gas increases over the same period.

4.4.2.2 Effects on Surface Temperatures and

Sea Ice

The ozone hole-related summertime trends in the 

Southern Annular Mode are also linked to changes in 

Antarctic surface temperatures and sea ice. The Antarctic Peninsula region has experienced exceptional recent 

warming in austral summer (in excess of 1°C/decade at 

some stations) since the late 1970s, greatly exceeding the 

global average and making it one of the fastest changing 

regions on Earth (e.g., Marshall et al., 2006; Zazulie et al., 

2010). In contrast, a summer seasonal cooling trend has 

been noted at stations on the high plateau of east Antarctica (e.g., Thompson and Solomon, 2002; Turner et al., 

2005), and these are linked to the previously noted circulation changes. The stronger eastward flow associated 

with the recent summertime trend in the Southern Annular Mode has reduced the effectiveness of blocking by 

the topography of the peninsula, allowing more maritime 

airmasses to penetrate across the peninsula and causing 

summer season warming on the east side of the peninsula 

that is three times stronger than that observed on the west 

side (Marshall et al., 2006). Thus both the spatial and seasonal patterns of the observed summertime surface climate 

changes in the past several decades support the identification of the Southern Annular Mode and ozone loss as a 

key factor in summertime climate change over much of 

Antarctica. The breakup of the Larsen-B ice shelf in 2002 

was likely due at least in part to the remarkable warming 

in the Antarctic Peninsula region (Marshall et al., 2006).

The effect of changes in circulation on Antarctic sea 

ice has been probed using numerical analyses constrained 

by observed (reanalyzed) winds and temperatures (Goosse 

et al., 2009). This approach suggests that strengthening 

of the eastward winds has likely affected the spatial pattern of sea ice distributions, mainly through changes in sea 

ice advection (Goosse et al., 2009). The spatial pattern of 

satellite-observed recent changes in sea ice, with increased 

ice extent in the Ross Sea region and decreases near the 

peninsula, is well reproduced when atmospheric circulation and temperature changes are considered (as shown 

in Figure 4-20, taken from Goosse et al., 2009). Atmospheric circulation changes in response to ozone loss have 

been probed by several models, including, e.g., Gillett 

and Thompson (2003) and Turner et al. (2009). Natural 

variability in the Southern Annular Mode, the El NiñoSouthern Oscillation, and/or the Pacific Decadal Oscillation may also contribute to recent trends in sea ice (Udagawa et al., 2009), as could other factors such as changes in 

precipitation or ocean heat transport (Zhang, 2007).

The very sparse spatial coverage of long-term 

Antarctic temperature records makes it difficult to derive 

trends averaged over the continent as a whole. Recently, 

reconstruction methods employing statistical approaches 

to data filling have been employed to derive trends averaged across Antarctica, drawing from composites using a 

range of station data, ice cores, and available automated 

weather stations (Chapman and Walsh, 2007; Monaghan 

et al., 2008; Steig et al., 2009). Strong warming has been 

derived for the west Antarctic region in some of these 

analyses (e.g., Steig et al., 2009), although this is heavily dependent on data from a single record (Byrd station). 

The most recent reconstruction suggests that Antarctica 

taken as a whole shows warming when the time period 

since 1957 is considered (Steig et al., 2009), although 

weaker warming is seen in some other reconstructions 

from 1960 (see Figure 4-21). Steig et al. (2009) and Gillett et al. (2008) suggest an important role for greenhouse 

gases in driving Antarctic temperatures especially prior 
4.32

Chapter 4

to onset of the ozone hole. This finding does not contradict the recent summer season cooling observed over east 

Antarctica since 1980 in association with the ozone hole; 

indeed those observations are included in the analysis (see 

Figure 4-21).

4.4.2.3 Effects on Southern Ocean Temperatures and Circulation

The summertime trend in the Southern Annular 

Mode implies an associated poleward shift and intensification of the surface zonal wind stress field and its curl (Oke 

and England, 2004; Cai and Cowan, 2007; Fyfe et al., 

2007). In this subsection, we discuss the implications of 

the poleward intensification of the surface wind stress for 

the Southern Ocean temperature and circulation. Greenhouse gas forcing and ozone forcing both appear to affect the tropospheric surface winds, with the ozone-forced 

signal peaking in austral summer (e.g., Kushner et al., 

2001; Thompson and Solomon, 2002; Gillett and Thompson, 2003; Miller et al., 2006; Arblaster and Meehl, 2006; 

Haigh and Roscoe, 2009). The following discussion is 

thus relevant to projections of Southern Ocean change under increasing greenhouse gases and ozone recovery (Section 4.5). However, we will focus on effects that can be 

plausibly linked to tropospheric impacts of stratospheric 

ozone depletion. Southern Ocean impacts also have implications for global climate and the Southern Ocean carbon 

cycle (Section 4.4.2.4).

Some of the influence of surface wind stress on the 

Southern Ocean can be understood by using surface wind 

stress and its curl to diagnose Ekman transport and pumping (e.g., Fyfe and Saenko, 2006; Cai and Cowan, 2007). 

In the classical description, the surface eastward winds 

drive (1) the eastward Antarctic Circumpolar Current and 

(2) a meridional overturning circulation that consists of 

upwelling to the south of the Antarctic Circumpolar Current, northward flow at the surface at the latitude of the current, and downwelling to the north of the current. When 

ozone forcing results in a poleward intensification of the 

surface wind stress field, the Antarctic Circumpolar Current and wind-driven overturning circulation are expected 

to similarly undergo a poleward intensification. From the 

preindustrial period to the end of the 20th century, Fyfe 

and Saenko (2006) diagnose, on the basis of wind stress 

output from the CMIP3 models, a poleward shift of about 

0.9 degrees latitude and a 5% strengthening of the Antarctic Circumpolar Current for the ensemble mean of the 

CMIP3 models. Cai and Cowan (2007) find that the ensemble mean of the CMIP3 “OZONE” simulations (which 

include ozone depletion) produce a zonal wind stress trend 

(expressed as a Southern Annular Mode trend) that is consistent with trends inferred from National Centers for Environmental Prediction/National Center for Atmospheric 

Research reanalyzed winds, and diagnose an accompanying change to the wind-driven ocean circulation in the subtropics and extratropics for these simulations. Consistent 

with Miller et al. (2006) and Son et al. (2009b), Cai and 

Cowan (2007) find that the “NO OZONE” simulations do 

not, on average, produce a realistic wind stress/Southern 

Annular Mode trend.

a b

Observations Model

Figure 4-20. The trend of annual-mean sea ice concentration over the period 1980–2000 for (a) the observa The trend of annual-mean sea ice concentration over the period 1980–2000 for (a) the observations (Rayner et al., 2003) and (b) for the model results averaged over the six simulations with data assimilation 

(units are the fractional change per decade). The black line represents the location of the climatological average September ice edge, as defined by an annual ice concentration equal to 15%, for the observations (a) and 

the model (b). Based on Goosse et al. (2009).
4.33

Stratospheric Changes and Climate

Figure 4-21. Spatial plots 

of the temporal trends (K/

decade) of annual nearsurface temperature for 

three different periods: 

(a and b) 1960–2002, (c 

and d) 1970–2002, and 

(e and f) 1980–2002. Figures a, c, and e are from 

the RECON temperature 

reconstruction, and Figures b, d, and f are from 

the CHAPMAN temperature reconstruction (figure 

from Monaghan et al., 

2008; reconstructions described in that paper).
4.34

Chapter 4

Key oceanographic observations5 and ocean model 

analysis support the viewpoint that the ocean circulation 

responds to the wind stress changes in a manner consistent with wind-driven ocean circulation theory. In particular, a significant temperature trend to depths up to 

1200 m has been observed in the Antarctic Circumpolar 

Current (Gille, 2002, 2008; Aoki et al., 2003; Sprintall, 

2008; Böning et al., 2008). Figure 4-22a shows observed 

oceanic temperature trends as a function of depth and dynamic height, which serves as a meridional coordinate, 

from Gille (2008). There is strong subsurface warming 

in the Antarctic Circumpolar Current, which is located 

in the 0.4–1.2 m dynamic height range. The subsurface 

warming is in large part consistent with a poleward shift 

of the ocean temperature structure by 1° latitude every 40 

years (Figure 4-22b): 87% of the trend below 200m in the 

Antarctic Circumpolar Current region is explained by this 

construction. Important aspects of trends in the Antarctic Circumpolar Current region are reproduced in ocean 

models and in the CMIP3 coupled models, which allow 

a significant portion of the subsurface warming to be attributed to changes in the wind stress field (Oke and England, 2004; Fyfe, 2006; Fyfe et al., 2007; Gille, 2008). For 

example, Oke and England (2004), using an ocean model 

driven by prescribed atmospheric forcings, and Fyfe et al. 

(2007), using an ocean model coupled to a simplified atmospheric model, obtain intensified subsurface warming 

when realistic surface wind stress trends are imposed.

Since we can attribute much of the poleward shift 

in the surface wind stress to ozone forcing (Gillett and 

Thompson, 2003; Cai and Cowan, 2007; Son et al., 2009b; 

Fogt et al., 2009), we can infer that ozone forcing in the 

1970–2000 period contributes to the observed subsurface 

warming in the Southern Ocean. That is, the stratospheric 

ozone-induced change in the Southern Ocean opposes the 

effects of global warming of the Earth’s surface; greenhouse gas forcing warms the ocean surface and increases 

high latitude precipitation, which stratifies the ocean surface and thereby reduces oceanic mixing of heat (and, by 

extension, the uptake of anthropogenic carbon dioxide). In 

the Geophysical Fluid Dynamics Laboratory coupled models CM2.0 and CM2.1, Russell et al. (2006) find that the 

enhanced Ekman transport signal associated with positive 

Southern Annular Mode-related wind trends counteracts 

this and contributes to an increased heat uptake. The implications of the circulation changes for the Southern Ocean 

carbon cycle are discussed in Section 4.4.2.4. Under ozone 

recovery (Section 4.5), the ozone-forced Southern Annular 

5 It is outside the scope of this report to assess the calibration 

and sampling of long-term ocean observations; we will place more 

confidence in conclusions that have been repeatedly verified in 

the literature.

Mode trend in December, January, and February is expected to weaken or change sign (see Section 4.5); this effect 

would tend to reduce ocean heat uptake, but the net effect 

will depend on the combined effects of CO2 increases and 

ozone recovery on the Southern Annular Mode.

Recent oceanographic literature provides some caveats about the straightforward connection between wind 

stress and Southern Ocean circulation. The complication 

relates to the effects of oceanic mesoscale (10 to 50 km 

scale) eddies, which are responsible for much of the transport of heat and constituents in the Southern Ocean. The 

CMIP3 generation ocean models cannot explicitly resolve 

mesoscale eddies and instead parameterize their effects 

with dynamically motivated mixing schemes (Gent and 

McWilliams, 1990). A developing body of work (e.g., 

Radko and Marshall, 2003; Hallberg and Gnanadesikan, 

2006; Meredith and Hogg, 2006; Böning et al., 2008; 

Screen et al., 2009) shows that when they are explicitly resolved, mesoscale eddy effects could modify some of the 

conclusions presented above. In particular, for relatively 

fine ocean model resolution, the mesoscale eddy field 

can respond strongly to changes in wind stress and can 

counteract some of the changes to the Ekman circulation 

among other effects (e.g., Hallberg and Gnanadesikan, 

2006; Meredith and Hogg, 2006; Screen et al., 2009). In 

a sensitivity test, Fyfe et al. (2007) impose this kind of effect in a coarse-resolution ocean model, but their approach 

has not been independently validated. Furthermore, not

(a) (b)

Depth (m)

Dynamic Height (m) Dynamic Height (m)

Figure 4-22. (a) Rate of change of ocean temperature (°C/yr) as a function of dynamic height (m), 

which serves as a meridional coordinate, from shipbased and autonomous float data. The Antarctic 

Circumpolar Current is in the range 0.4–1.2 m. (b) 

Effective rate of change of ocean temperature (°C/yr) 

that would result if ocean temperature structure were 

shifted poleward at a rate of 1° latitude per 40 years. 

Based on Gille (2008).
4.35

Stratospheric Changes and Climate

all aspects of observed changes in the density structure in 

the Antarctic Circumpolar Current are represented by the 

wind-driven ocean circulation theory. In particular, intensification of the wind-driven Antarctic Circumpolar Current is associated with an increased slope (baroclinicity) of 

density surfaces. This increased slope is seen in CMIP3 

class ocean models (e.g., Russell et al., 2006) but has not 

been detected in observations, which might be the result of 

mesoscale eddy effects (Böning et al., 2008). Solid conclusions will require a better understanding of uncertainties in trend analysis of multidecadal ocean observations. 

Nevertheless, since this recent work has implications for 

Southern Ocean circulation, density structure, temperatures, and carbon cycle, and connects to global climate 

issues, it appears that high-resolution ocean models will 

be required to fully evaluate the role of Antarctic ozone 

depletion in Southern Ocean climate change (see also Ito 

et al., 2010).

To summarize, the dynamical coupling between the 

stratosphere and troposphere in conjunction with forcing 

by the ozone hole suggests that the wind-driven Southern Ocean circulation should intensify as a result of polar 

stratospheric ozone depletion. This effect may contribute 

to observed subsurface warming in the Southern Hemisphere and has the potential to attenuate some of the surface 

global warming associated with greenhouse gas increases. 

Some predictions of the changes to the Antarctic Circumpolar Current and the Southern Hemisphere overturning 

could be dependent on ocean model resolution; thus, highresolution ocean models will likely be required to assess 

with confidence how the Antarctic ozone hole and ozone 

recovery are likely to affect the Southern Ocean.

4.4.2.4 Stratospheric Links to Southern Ocean

Carbon

The summertime trend in the Southern Annular 

Mode also provides a potential link between stratospheric 

ozone depletion and the flux of carbon at the Earth’s surface: the ozone hole is linked to changes in the surface 

winds, and those changes in the surface winds are expected 

to affect the flux of carbon at the ocean surface. Our current understanding of these linkages is assessed below.

The primary anthropogenic driver of climate change 

is the emission of more than 9 gigatonnes of carbon (GtC) 

annually in the form of CO2, mainly due to fossil fuel 

burning with smaller contributions from land use change 

(Forster et al., 2007a). Thus the processes that remove 

atmospheric carbon dioxide are critical to understanding 

current climate change.

The Southern Ocean plays a key role in the uptake 

of global carbon dioxide. As carbon dioxide increases, the 

net sink of CO2 in the Southern Ocean would be expected 

to increase in the absence of other changes. However, recent studies suggest that the Southern Ocean carbon sink 

has not strengthened over about the past two decades, and 

this behavior has been linked to changes in winds. Strong 

eastward winds drive a northward Ekman circulation in 

the Southern Ocean, which pushes surface waters away 

from the Antarctic continent and drives a strong ocean circulation that brings up water and dissolved carbon from 

the deep ocean.

As evidenced in Figure 4-23, the regions displaying broad reductions in carbon uptake correspond to those 

that underlie the largest tropospheric wind changes associated with the Southern Annular Mode, as obtained in a 

model driven by ozone depletion. Numerous studies have 

recently probed how changes in wind stress linked to the 

Southern Annular Mode have influenced carbon uptake in 

the Southern Ocean using a range of different approaches, 

some relying on global climate models, some on models 

forced by National Centers for Environmental Prediction/

National Center for Atmospheric Research reanalyzed 

winds, some making use of inversion methods based on 

CO2 observations, and some using station-based observations of atmospheric CO2 (Wetzel et al., 2005; Butler et 

al., 2007; Le Quéré et al., 2007; Lenton and Matear, 2007; 

Lovenduski et al., 2008; Lenton et al., 2009; Khatiwala et 

al., 2009).

The observation-based inversion studies are subject 

to uncertainties in sampling and in the winds prescribed 

from reanalysis (Baker et al., 2006; Law et al., 2008; Le 

Quéré et al., 2007). Nevertheless these different types of 

analyses, including both models and inversion methods, 

support recent reductions in Southern Ocean carbon uptake. Taken together, this body of work suggests that the 

changes in wind stress over about the past three decades 

have reduced the net Southern Ocean carbon uptake by 

about 0.6–5 Gt (about 0.02–0.16 GtC/yr) compared to 

the trend that should be expected due to increasing atmospheric CO2. One context for these numbers is provided 

by comparison to full compliance with the Kyoto Protocol, which corresponds to a reduction of global carbon 

emissions of about 0.5 GtC/yr by 2012 (see Velders et al., 

2007, and Chapter 5 of this Assessment). The dominant 

mechanism for a reduction in carbon uptake is generally 

thought to be stronger upwelling and subsequent outgassing of natural carbon from the deep ocean induced by the 

wind stress changes (Le Quéré et al., 2007; Lovenduski et 

al., 2008; Hall and Visbeck, 2002). Note that both natural and anthropogenic carbon fluxes must be considered 

in assessing changes in the net carbon flux. Lenton et al. 

(2009) emphasized the dominant role of ozone depletion 

in driving the change in net carbon uptake in model simulations that explicitly tested the role of ozone versus other 

greenhouse gas forcings. Khatiwala et al. (2009) noted 

that only a small change (about 5%) in the ocean carbon 

sink is required to explain the observations.
4.36

Chapter 4

The downward transport of anthropogenic carbon 

by stronger overturning and the increasing thermal stratification of the Southern Ocean are also important to carbon 

and heat uptake, but analyses of the data suggest that these 

effects have been overwhelmed by the effect of strengthening eastward flow on natural carbon fluxes (Le Quéré 

et al., 2007). Ito et al. (2010) suggest that higher spatial 

resolution may be needed to ensure accurate evaluation of 

Trends in Ocean Carbon

Calculated and Observed Trends in Near-Surface Winds

Longitude

Latitude

Sea-Air CO2 Trend ( atm/yr)

Model

Observations

m/s

m/s

Figure 4-23. Top panel shows trends in the observed partial pressure of CO2 for ocean minus air, for 1981–

2007 (μatm/yr). The observed trends shown by the circles are calculated by fitting a linear trend to repeated 

measurements of surface-ocean and air CO2. Positive (red) values indicate regions where the partial pressure 

of CO2 in the ocean is increasing faster than atmospheric CO2, indicating a weakening of the net ocean uptake 

(from Le Quéré et al., 2009). Large, medium, and small circles are plotted for trends with errors of <0.25, 0.25–

0.50, and >0.50 μatm yr

−1, respectively. Trends are also shown for three broader areas where only a single 

estimate can be made (areas shown by hatching). As indicated on the figure, the trends in southern circumpolar waters were estimated from austral winter data only, whereas the trends in the South Indian Ocean were 

estimated for 1991–2007 only. No seasonal cycle was removed from the data in the western equatorial Pacific 

region. The amplitudes of the trends in the hatched regions are: equatorial Pacific warm pool (near Indonesia), 

+0.3 ± 2 μatm/yr; Southern Ocean, +0.5 ± 0.6 μatm/yr; South Indian Ocean, +0.4 ± 0.1μatm/yr . See Le Quéré 

et al. (2009) for details. The middle panel shows December–February calculated trends in near-surface winds 

at southern middle and high latitudes, from global model simulations forced by observed ozone depletion from 

1979 to 1997, while the bottom panel shows observed December–May wind trends in the same region (based 

on Gillett and Thompson, 2003). Units on the bottom two panels are meters/second.
4.37

Stratospheric Changes and Climate

the oceanic processes that could contribute to the carbon 

uptake and its changes.

Future changes in the Southern Ocean carbon sink 

are subject to an uncertain combination of processes. As 

recovery of the ozone hole occurs, the circumpolar eastward wind changes related to ozone depletion should reverse (Section 4.5). However, the weakening of the circumpolar eastward flow due to ozone recovery is expected 

to be roughly compensated for by a strengthening of the 

circumpolar eastward flow due to increases in carbon dioxide and other greenhouse gases by late 21st century in 

summer (see Section 4.5). The effect of the eastward flow 

on the ocean carbon uptake this century will depend on 

the rate of increase of anthropogenic carbon in the surface 

ocean and the relative strength of the downward transport 

of anthropogenic carbon compared to the upward transport 

of natural carbon, which are both linked to changes in the 

ocean circulation (see, e.g., Zickfeld et al., 2008). Further, the extent to which the ocean circulation will adjust 

to altered stratification and winds over long timescales is 

uncertain.

4.4.3 Stratospheric Variations and the 

Width of the Tropical Belt

Changes in the tropical and extratropical tropopause layer, and their relation to changes in stratosphere/

troposphere exchange of trace constituents, particularly 

water vapor, were addressed in the 2006 Ozone Assessment (Chapters 2 and 5: Law and Sturges et al., 2007, and 

Baldwin and Dameris et al., 2007). Since that Assessment, 

research has linked tropopause changes with changes in 

several other stratospheric and upper-tropospheric features. 

For example: 

• the region of low column ozone values typical of the 

tropics has expanded in the Northern Hemisphere 

(Hudson et al., 2006);

• the jet streams have moved poleward in both hemispheres (Fu et al., 2006; Archer and Caldeira, 2008);

• the downward branches of the Hadley cell (where 

deep clouds are relatively rare compared with the upward branch) have moved poleward, as seen in outgoing longwave radiation, a surrogate for cloud top 

temperature (Hu and Fu, 2007). Mean meridional 

mass flux changes independently support this finding 

of Hadley cell expansion (Hu and Fu, 2007); and

• the region of high tropical tropopause expanded poleward in both hemispheres (Seidel and Randel, 2007).

The qualitative consistency of these observed 

changes in independent data sets suggests a widening of 

the tropical belt of between 2 and 5 degrees latitude between 1979 and 2005 (Figure 4-24), as summarized by 

Seidel et al. (2008) and Reichler (2009). Combined with 

an increase in the height of the tropical tropopause (Zhou 

et al., 2001; Santer et al., 2003; Seidel and Randel, 2006), 

the widening contributes to an increase in the volume of 

the tropical troposphere of ~5%.

Widening of the tropical belt (i.e., the region of 

globe with tropical characteristics), including poleward 

migration of the jet streams and associated migration of 

storm tracks, could be associated with changes in global 

precipitation patterns. Lu et al. (2007) found a significant 

correlation between expansion of the tropical belt, as determined by the location of the downward branches of the 

Hadley circulation, and the location of the subtropical dry 

zones in climate models. Poleward movement of precipitation patterns in climatically sensitive subtropical regions 

might be among the most significant surface climate impacts of tropical expansion, and may already have begun 

in North America (Seager et al., 2007). Zhou et al. (2001) 

also note that circulation changes related to tropical beltwidening impacting troposphere-to-stratosphere transport 

Ozone (NH only, Hudson et al.)

Outgoing longwave radiation (Hu and Fu)

Hadley circulation (Hu and Fu)

Jet stream separation (Reichler)

Tropopause (Seidel and Randel)

70

60

50

Width of Tropics (degrees latitude)

Width of NH Tropics (degrees lattitude)

1980 1985 1990 1995 2000 2005

35

30

25

Year

Figure 4-24. Changes in several estimates of the 

width of the tropical belt since 1979 (in degrees latitude). These include the width of the Hadley circulation, based on both outgoing longwave radiation and 

horizontal winds streamfunction (Hu and Fu, 2007); 

the separation of the Northern and Southern Hemisphere subtropical jet-stream cores (Reichler, 2009); 

the width of the region of frequent high tropopause 

levels (Seidel and Randel, 2007); and the width of 

the region with tropical column ozone levels (Northern Hemisphere only, right axis, Hudson et al., 2006). 

(Based on Seidel et al., 2008.)
4.38

Chapter 4

may produce changes in the average stratospheric entry 

value of water vapor.

The cause(s) of these observed changes, their relation to changes in the lower troposphere and at the surface, 

and the likelihood that they will continue or be reversed 

in the future, are areas of active research. Insights into 

these issues have been obtained from two types of global 

climate models. One is coupled atmosphere-ocean global 

climate models, particularly the CMIP3 models, which 

consider ocean feedback but do not represent stratospheric processes in great detail. A second type is coupled 

chemistry-climate models, with more complete representation of the stratosphere but without an interactive ocean 

model. However, results to date from both CMIP3 and 

coupled chemistry-climate model simulations are uncertain because neither type of model includes all relevant 

processes and because the horizontal resolution of these 

global model systems is generally several degrees of latitude, which may make it difficult to resolve smaller but 

significant changes of the horizontal extent of the tropics.

Simulations of past climate change by CMIP3 models that include increases in greenhouse gas concentrations 

and stratospheric ozone depletion reveal changes in the 

width of the tropics as measured by changes in the region 

of high (tropical) tropopause heights (Lu et al., 2009), and 

modeled tropopause height changes have been linked to 

poleward migration of the jet streams (Lorenz and DeWeaver, 2007; Son et al., 2009b). More fundamentally, 

model simulations have linked imposed perturbations 

in heating of the stratosphere to tropospheric dynamical 

changes, including latitudinal shifts in jet streams (Simpson et al., 2009). But the widening simulated by CMIP3 

models, for both the 20th and 21st centuries, is smaller 

than the observed widening over the last several decades 

(Johanson and Fu, 2009), suggesting both that the widening is not the result of natural variability alone (because it 

is not found in control simulations without anthropogenic 

forcing) and that the responsible mechanisms may not be 

well represented by those models.

The importance of stratospheric processes, particularly those related to ozone loss and recovery, is highlighted by comparison of CMIP3 and chemistry-climate 

model simulations (and projections) of tropopause height 

changes (Son et al., 2009a). Son et al. (2010) find that the 

increase in width of the Southern Hemisphere tropics in 

the CCMVal-2 models is linearly related to the simulated 

amount of ozone depletion. Lu et al. (2009) also show 

that widening of the tropics only occurs when they include 

the radiative impact of ozone depletion in their simulation. Different mechanisms may be responsible for different aspects of the observed changes in the width of the 

tropical zone, with halocarbons influencing ozone changes 

and tropopause height change at the subtropical edges of 

the tropical belt, and with sea surface temperatures and 

CO2 influencing tropopause changes within the tropical 

region (Deckert and Dameris, 2008; Lamarque and Solomon, 2010).

4.4.4 Effects of Stratospheric Variations on 

Tropospheric Chemistry

Changes in the stratosphere can also affect chemical processes in the troposphere in two principal ways. 

First, the return of air from the stratosphere at midlatitudes 

to high latitudes is the sink of many stratospheric chemical species. This stratosphere-troposphere transport is 

an important term in the overall budget of tropospheric 

ozone. Second, the overhead column of stratospheric 

ozone affects the penetration of UV radiation to the troposphere. Changes in column ozone are particularly 

important for modulating UV-B radiation (wavelength 

less than ~320 nanometers) which causes photolysis of O3

to yield excited O(1D) atoms. O(1D) reacts with H2O in 

the main formation reaction of tropospheric hydroxyl radicals (OH), the most important tropospheric oxidant. Any 

changes in the tropospheric circulation, which are driven 

from the stratosphere, can also change the distribution of 

tropospheric ozone.

The stratosphere can affect tropospheric chemistry 

through other mechanisms. For example, stratospheric radiative or circulation changes that contribute to different 

surface temperatures may impact emissions. However, 

these contributions will be very small compared to the 

large tropospheric climate changes induced by increasing atmospheric greenhouse gases. Therefore, this report 

concentrates on the impact of stratosphere-troposphere 

transport and column ozone changes. Isaksen et al. (2009) 

have recently reviewed tropospheric chemistry-climate 

interactions.

Tropospheric ozone levels are determined by the 

complex interaction between photochemistry and transport. The primary sources of tropospheric ozone are transport from the stratosphere and in situ photochemical production, and the main sinks are photochemistry and surface 

deposition. The net global tropospheric ozone burden is a 

small residual of these terms. Although it is recognized 

that input of ozone-rich air from the stratosphere (via 

stratosphere-troposphere transport) is an important term in 

this overall burden, there are significant variations in its 

quantification, which is usually done through diagnosing 

the stratosphere-troposphere transport term in global 3-D 

models (e.g., Stevenson et al., 2006). Wild (2007) and Wu 

et al. (2007) updated estimates of terms in the tropospheric 

ozone budget and investigated the causes of model-model 

differences from sensitivity runs of their own models. 

Wild (2007) concluded that some previous differences in 

modeled budgets could be reconciled by using updated 
4.39

Stratospheric Changes and Climate

chemical mechanisms, but uncertainties in stratospheretroposphere transport remained. Wu et al. (2007) found 

that the model-model differences were determined largely 

by differences in emissions and stratosphere-troposphere 

transport. The current best estimate of this term from observations is 550 ± 140 teragrams (Tg) per year (Olsen et 

al., 2001), which is in good agreement with model estimates of 550 ± 170 Tg/yr (Stevenson et al., 2006). The 

mean lifetime of ozone in the troposphere is around 25 

days. Variations in the modeled tropospheric ozone burden, e.g., 340 ± 40 Tg (corresponding to a mean mixing 

ratio of approximately 45 ± 5 parts per billion by volume, 

ppbv) from Stevenson et al. (2006), are much smaller than 

variations in stratosphere-troposphere transport, although 

similar to variations in other ozone budget terms. Therefore, large variations in modeled flux from the stratosphere 

are not reflected in such large variations in the mean tropospheric ozone burden; the large photochemical and 

deposition terms damp out these variations.

Observations showed substantial declines in ozone 

throughout the troposphere in 1992–1993 following the 

eruption of Mt. Pinatubo (see Fusco and Logan, 2003) and 

a good correlation between trends in lower-tropospheric 

and lower-stratospheric ozone from 1992–2004 (Ordóňez 

et al., 2007). These declines were generally largest in the 

upper troposphere and linked to ozone loss in the lower 

stratosphere (which is discussed in Chapter 2). Terao et 

al. (2008) showed that stratospheric ozone contributes 

significantly (30–40% at midlatitudes to high latitudes) to 

the spring O3 maximum at 500 hPa. The effects of past 

decreases in stratospheric ozone are not usually included 

in models used to estimate the tropospheric ozone budget. 

Fusco and Logan (2003) looked at the impact of decreases 

in lower-stratospheric ozone in their study of tropospheric 

ozone changes from 1970–1995. They found that decreases in lower stratosphere ozone might have reduced 

the stratospheric source by 30% over this time period. 

More recently, Hsu and Prather (2009) diagnosed the contribution of stratosphere-troposphere transport to the tropospheric ozone budget and included the effect of stratospheric ozone depletion from 1979 to 2004. They found 

a reduction of up to 10% in the stratosphere-troposphere 

transport ozone flux, but this caused only a small decrease 

in mean tropospheric ozone of 1 ppb (about 2%).

A number of older studies investigated how decreases in column ozone in the 1980s and 1990s affected 

tropospheric OH trends (e.g., Madronich and Granier, 

1992; Bekki et al., 1994; Fuglestvedt et al., 1994; Granier 

et al., 1996). More recently, Duncan and Logan (2008) 

used a three-dimensional (3-D) model to investigate the 

impact of various forcings, including changing overhead 

column ozone, on tropospheric OH and carbon monoxide 

(CO) from 1988–1997. During this relatively short period, 

overhead column ozone decreased due to the solar cycle 

phase and eruption of Mt. Pinatubo, as well as the longterm trends in ozone-depleting substances. The decreased 

ozone column led to important year-round increases in 

OH, where a −3% change in the column caused a +4% 

change in OH. In spring/summer at higher latitudes, column ozone changes of −10% caused an increase in OH of 

about +8%. This contribution from stratospheric ozone 

variations makes an important contribution to the overall 

variability in tropospheric OH of ±10% over the past few 

decades inferred from methyl chloroform (Bousquet et al., 

2005; Prinn et al., 2005). This contribution to increased 

OH made a small impact on the model CO trends, about 

0.1–0.2 Tg/yr averaged over the period considered compared to a CO burden of about 500 Tg. However, this 

change was compensated for by an increase from oxidation of increasing CH4, while changes in fossil fuel emissions had a much larger impact regionally.

Isaksen et al. (2005) studied the impact of decreased 

column ozone on tropospheric chemistry under different pollution conditions. They imposed a uniform global 

decrease in column ozone, affecting only their model’s 

photolysis calculation, and studied the impact on surface 

ozone. They found that in polluted regions, i.e., regions 

of net tropospheric ozone production, the increased photochemical activity from faster photolysis led to more ozone 

production which in turn led to an additional 1 ppb surface 

ozone for the 10% column reduction over continental Europe in January. In contrast, in cleaner remote regions the 

enhanced photochemical activity increased the net sink of 

ozone. The largest effect was a decrease of 3 ppb at low 

latitudes to midlatitudes, but a change of 10% in the ozone 

column represents a large perturbation for these latitudes. 

Overall, these changes in surface ozone are therefore 

small.

4.4.5 Influence of the Stratosphere on the 

Impact of Solar Variability on Surface 

Climate

Changes in the stratosphere in response to solar 

variability have the potential to influence tropospheric climate by many of the radiative and dynamical processes 

outlined in Sections 4.4.1–4.4.4 above. Some specific 

linkages between solar variability, the stratosphere, and 

surface climate are assessed below. More detailed discussions of many of the issues related to the impact of solar 

variability on climate are presented by Gray et al. (2010).

The signals in stratospheric ozone and temperature 

produced by 11-year solar cycle variations in incoming radiation have been discussed in Section 2.4.3.2 of Chapter 

2 and Section 4.3.1 of this chapter, respectively. While 

some details of the responses are not fully understood, a 

picture is emerging in the tropics in which the temperature 
4.40

Chapter 4

of the tropical upper stratosphere is warmer by 1.1–1.8 K 

and the lower stratosphere warmer by 0.5–0.8 K when the 

sun is more active, relative to when it is quiet (Randel et 

al., 2009; Gray et al., 2010). There is a related signal in 

zonal winds in a strengthening of the subtropical lower 

mesospheric jet in the winter hemisphere during high solar 

activity (Frame and Gray, 2010). A solar signal has also 

been detected in the stratospheric polar night jet in both 

hemispheres, and thus in the annular modes, (e.g., Kodera, 

2002; Boberg and Lundstedt, 2002; Thejll et al., 2003; 

Kuroda and Kodera, 2004, 2005; Kuroda et al., 2007; Lee 

and Hameed, 2007; Barriopedro et al., 2008; Lee et al., 

2008). Most of these studies, however, have used simple linear regression or confined their discussions to correlation coefficients. Hence they detect only the linear 

component of the response, and have not considered the 

impact of other potential forcing factors such as volcanic 

eruptions. Analyses using data sets that include the most 

recent solar maximum period, during which there was no 

major coincident volcanic activity, have now shown that a 

separation is indeed possible (Frame and Gray, 2010).

In a multiple regression analysis, Keckhut et al. 

(2005) find a temperature response in the tropics in rocketsonde and lidar data similar to that in the reanalysis studies. The response includes a seasonally varying out-ofphase signal in midlatitudes that Hampson et al. (2005) 

relate to the occurrence of stratospheric warmings. Haigh 

and Roscoe (2006) carried out a multiple regression analysis of time series of the Northern and Southern Annular 

Mode indices throughout the depth of the atmosphere. No 

significant response to the 11-year solar cycle was found if 

the solar and stratospheric quasi-biennial oscillation terms 

were included separately, but when they were combined 

into a single term (in which the phase of solar activity 

modulates the stratospheric quasi-biennial oscillation) to 

represent their interaction, a strongly significant result was 

found. This finding is consistent with the original results 

of Labitzke (1987), in which polar stratospheric height 

data were sorted by phase of the stratospheric quasibiennial oscillation and level of solar activity.

Variations in incoming total solar irradiance of order 

0.1% occur over the 11-year solar activity cycle (Fröhlich 

and Lean, 2004) resulting in a top-of-atmosphere radiative forcing of approximately 0.2 W/m2

. On multidecadal 

timescales, underlying variations in total solar irradiance 

are less certain, but assessments suggest that there has been 

an overall increase in the range 1–4 W/m2 since the year 

1700, giving a top-of-atmosphere radiative forcing of 0.17–

0.67 W/m2 since that date (Gray et al., 2010). The amount 

of radiation reaching the tropopause, which provides a better indicator of climate radiative forcing, is dependent on 

the spectral composition of the radiation and the response 

of stratospheric ozone. This effect has been assessed to be 

a small addition to the radiative forcing value (Larkin et 

al., 2000), but recent measurements from the Solar Radiation and Climate Experiment satellite (Harder et al., 2009) 

indicate that there may be compensating changes at UV 

and visible wavelengths which imply that solar radiative 

forcing at the tropopause would be out of phase with that 

at the top of the atmosphere due to absorption of the UV in 

the stratosphere (Haigh et al., 2010).

As well as direct radiative effects, variations in solar activity may produce a solar signal in the troposphere 

“top-down” from an initial response in the stratosphere. 

Early model studies of UV variations in the stratosphere 

(Haigh, 1996, 1999; Shindell et al., 1999; Balachandran 

et al., 1999; Larkin et al., 2000) obtained a response in the 

troposphere even though the near-surface in these model 

runs was constrained by imposed, seasonally varying, seasurface temperatures. The main response to enhanced UV 

and stratospheric ozone in these models is (1) an expansion of the Hadley cell and (2) a poleward movement of 

the tropical convective maxima, the tropospheric jets, and 

the midlatitude storm tracks. The pattern of these anomalies was similar to the signal found in tropospheric zonal 

winds in reanalysis data (Haigh, 2003; Frame and Gray, 

2010) and also in upper air temperatures and geopotential 

heights (Brönnimann et al., 2006) using a data set based on 

radiosonde and aircraft observations dating back to 1922. 

The amplitude of the modeled pattern depended on the 

magnitude of the ozone change but was generally about 

half the amplitude of that observed. The new Solar Radiation and Climate Experiment data, however, imply that 

the ozone response, and thus presumably the tropospheric 

response to the stratospheric warming, might be significantly larger.

This signal has the same spatial structure as, and 

thus presumably involves similar eddy-mean flow feedbacks to, the dominant pattern of atmospheric variability, i.e., the atmospheric annular modes. Feedback of the 

tropospheric zonal wind changes on the tropospheric eddy 

momentum fluxes appears to be important in establishing 

and maintaining the response to stratospheric forcing, as 

discussed in Section 4.4.2 (see also Polvani and Kushner, 2002; Kushner and Polvani, 2004, 2006; Song and 

Robinson, 2004). Coupling between the Hadley circulation and midlatitude eddies may also play a key part; in a 

mechanistic study focusing on the processes involved in 

a response to solar forcing, Haigh et al. (2005) obtained 

a zonal mean tropospheric response, qualitatively similar to the observed 11-year solar response, by imposing 

anomalous diabatic heating in the low-latitude lower stratosphere. Consistent with this result, the enhanced Hadley circulation response in the coupled chemistry-climate 

model simulations of Shindell et al. (2006) was linked to 

the additional condensational heating in the upper tropical 

troposphere and lower stratosphere relative to simulations 

with fixed ozone. Imposing a generic stratospheric heat-
4.41

Stratospheric Changes and Climate

ing in a simplified climate model, Simpson et al. (2009) 

have shown that it is the response of the eddy momentum 

fluxes to changes in structure of the tropopause region 

that drives the tropospheric response, so that details of the 

stratospheric response to solar variability may be important in determining the tropospheric signal.

Matthes et al. (2006), using a general circulation 

model with prescribed sea surface temperatures, found 

the response to UV variability in tropical vertical velocity 

was not uniformly distributed in longitude but was largest 

over the Indian and West Pacific Oceans—indicating an 

influence on the Walker circulation similar to that found in 

observations by Kodera (2004) and Kodera et al. (2007). 

The weakened ascent during solar maximum in the tropical 

troposphere of their model may result from the increased 

static stability in the tropopause region suppressing convection (Kodera and Shibata, 2006; Matthes et al., 2006).

Other studies have also identified solar influences 

on the strength and extent of the Walker circulation. van 

Loon et al. (2007) and Meehl et al. (2008) show a strengthening, at peak years of the 11-year solar cycle, which they 

identify as distinct from any El Niño-Southern Oscillation signal (van Loon and Meehl, 2008) although their 

approach has been questioned by Roy and Haigh (2010) 

and Zhou and Tung (2010). The associated sea surface 

temperature response at the peak years of the 11-year solar cycle is a cooler-than-normal equatorial eastern Pacific 

and poleward-shifted intertropical and South Pacific convergence zones with a warming following with a lag of a 

couple of years (Meehl et al., 2008; White and Liu, 2008).

The mechanisms invoked to explain this sea surface 

temperature response involve changes in the visible radiation absorbed by the tropical oceans, along with coupled 

air-sea interactions, and thus may be termed “bottom-up” 

rather than the “top-down” processes involving the stratosphere that are the focus of this report. Recent modeling 

studies (Rind et al., 2008; Meehl et al., 2009) have suggested that both pathways may combine to produce an 

amplified response to a small solar forcing. However, the 

responses probably depend on complex nonlinear interactions between different processes that have yet to be elaborated. Furthermore, all the mechanisms depend crucially 

on the spectral composition of the solar irradiance variability of which current understanding is currently being 

challenged by data from the Solar Radiation and Climate 

Experiment Spectra Irradiance Monitor (SIM) instrument 

(see the discussion in Section 4.3.1).

4.5 WHAT TO EXPECT IN THE FUTURE

The stratosphere is expected to change over the rest 

of this century in response to a range of forcing factors. 

Concentrations of the main anthropogenic greenhouse 

gases (CO2, N2O and CH4, hydrofluorocarbons (HFCs), 

and SF6) are expected to continue to rise for at least several decades, but concentrations of halogenated ODSs have 

leveled off and are expected to decrease. Future stratospheric aerosol loading (especially from volcanoes) is unpredictable. Chapter 5 of this Assessment provides a number of possible scenarios of emissions and concentrations 

of anthropogenic greenhouse gases. These atmospheric 

concentration changes will affect both ozone concentrations and stratospheric climate in the future. The evolution of the future ozone layer is discussed extensively in 

Chapter 3. Here we briefly assess the predicted changes 

through 2100 in the stratospheric climate changes assessed 

in this chapter.

4.5.1 Stratospheric Temperatures

The evolution of stratospheric temperature is expected to depend on three main factors: the recovery of the 

ozone layer, increasing concentrations of most greenhouse 

gases, and the resulting changes to the mean overturning 

circulation in the stratosphere (the Brewer-Dobson circulation). At least one experiment predicts that changes 

in Arctic sea ice will drive changes in polar stratospheric 

wave forcing and thus temperatures (Scinocca et al., 2009). 

The recovery of the Antarctic ozone hole is expected to 

lead to higher temperatures in the Antarctic stratosphere 

during spring, and a reduction in the strength of the stratospheric polar vortex as the effects of ozone loss diminish. Outside the Antarctic stratosphere, future changes in 

stratospheric ozone are expected to lead to a reversal of 

the ozone-induced cooling observed in both the lowermost 

and upper stratosphere (Chapter 10 of SPARC CCMVal, 

2010). However, in general the warming associated with 

ozone recovery will be outweighed by a larger cooling effect due to continued greenhouse gas increases (Chapter 

5 of WMO, 2007; Chapter 10 of SPARC CCMVal, 2010; 

McLandress et al., 2010; Jonsson et al., 2009).

The projected evolution of stratospheric temperatures in the 21st century depends on both the choice of scenario and climate model used to project that scenario. All 

scenarios have CO2 continuing to increase, with ODSs and 

stratospheric chlorine loading decreasing. An example 

of trends at 1 hPa is shown in Figure 4-25 for the NASA 

Goddard Space Flight Center (GSFC) Goddard Earth Observing System chemistry-climate model (GEOS) under 

the CCMVal REF-B2 scenario6. CO2-induced cooling 

(including its effect on ozone) dominates the response, 

although ozone recovery reduces the rate of cooling over 

the next few decades. Under this scenario all the CCMVal 

models exhibit global cooling in the middle and upper 

6 See SPARC CCMVal 2010, Chapter 2, for scenario details.
4.42

Chapter 4

stratosphere by several degrees, with much smaller temperature changes of undetermined sign in the lower stratosphere (typically less than 0.5 K) (Chapter 3 of SPARC 

CCMVal, 2010).

Temperatures in the lower tropical stratosphere 

and tropopause region could be significantly affected 

if atmospheric concentrations of HFCs rise. At the one 

ppb level HFCs give an additional warming of 0.2–0.5 

K to the lower tropical stratosphere region. This range 

depends on the spectral absorption characteristics of the 

HFC (Forster and Joshi, 2005). Note the HFCs with small 

atmospheric lifetimes are unlikely to reach high enough 

concentrations to have an appreciable effect (see Chapter 

5 of this Assessment).

4.5.2 Brewer-Dobson Circulation

Atmospheric model simulations forced with increased greenhouse gas concentrations exhibit an acceleration of the Brewer-Dobson circulation (e.g., Rind 

et al., 1998; Butchart and Scaife, 2001; Butchart et al., 

2010). This is an extremely robust response in the current generation of chemistry-climate models, as discussed 

in Section 4.3. The implications of a future increase in 

the Brewer-Dobson circulation are substantial. Such an 

increase would change the spatial distribution of stratospheric ozone, with increased total ozone at high latitudes 

and decreased total ozone in the tropics (see Figure 3-3 

in Chapter 3 of this Assessment). It would also increase 

the stratosphere-to-troposphere ozone flux (Hegglin 

and Shepherd, 2009), and it would decrease the net age 

of stratospheric air (e.g., Austin and Li, 2006; Garcia 

and Randel, 2008; Oman et al., 2009; Butchart et al., 

2010).

However, as noted in Section 4.2, the observational evidence for an increase in the Brewer-Dobson circulation is unclear. The observed decreases in tropical 

stratospheric temperatures are consistent with increased 

upwelling in the lower tropical stratosphere (e.g., Thompson and Solomon, 2005, 2009; Randel et al., 2009) as are 

the observed decreases in lower tropical stratospheric 

ozone (Randel and Wu, 2007; Chapter 2). But the reliability of the ozone and temperature data for assessing 

relatively small trends in this region of the atmosphere 

is still under investigation (e.g., Randel and Wu, 2006; 

Chapter 2). Estimates of long-term variations in age of 

air based on midlatitude balloon measurements of SF6

and CO2 over 1975–2005 have substantial uncertainties, 

but suggest no significant trends (Engel et al., 2009). 

Some effort has been made to use reanalyses to diagnose 

trends in the Brewer-Dobson circulation but without success due to inhomogeneities in the observational record 

and reanalyzed products (Iwasaki et al., 2009).

4.5.3 Stratospheric Water Vapor

Modeling studies of future greenhouse gas increases project that the tropical tropopause will warm and 

rise in altitude (Gettelman et al., 2009). One effect of a 

warmer tropical tropopause is an increase in stratospheric 

water vapor globally. Assumed increases in anthropogenic contributions to tropospheric methane, along with the 

possibility that natural sources of methane may increase 

in a warming climate, would also contribute to increases 

in stratospheric water vapor. Projected decreases in the 

net age of stratospheric air (Section 4.5.2) could lead to a 

slight decrease in stratospheric water vapor.

As presented in Gettelman et al. (2010), there is 

significant spread in chemistry-climate models’ representation of both tropical cold point temperatures (Figure 

4-12) and 80-hPa water vapor (used to represent stratospheric entry values), as shown in Figure 4-13. The trend 

in the historical record of tropical radiosonde cold point 

temperatures also does not agree with the multi-model 

1950 2000 2050 2100

Year

252

254

256

258

260

262

264

266

Temperature (K)

Latitude = 40°N

Pressure = 1 hPa

Deviation from Mean (K)

0

2

-2

4

-4

EESC

CO2

CH4

Total

Figure 4-25. Time series of simulated temperatures (K) at 40°N and 1 hPa (thin black line), with 

the statistical fit as the heavy black line. Data are 

for the NASA GSFC GEOS chemistry-climate model 

under the CCMVal REF-B2 scenario. Scale on left 

shows temperatures for this time series. Scale on 

right shows deviations from the mean temperature 

due to changes in equivalent effective stratospheric 

chlorine (blue line), CO2 (red line), and CH4 (orange 

line). Based on Stolarski et al. (2010).
4.43

Stratospheric Changes and Climate

mean; the amplitude of the observed trend is more than 

an order of magnitude larger than the modeled trend 

(Gettelman et al., 2009). Given the discrepancies between modeled and measured tropical temperatures and 

lower-stratospheric water, we cannot provide any reliable 

prediction as to what the future evolution of stratospheric 

water vapor will be in a climate with increasing greenhouse gas concentrations.

4.5.4 Tropopause Height and Width of the 

Tropical Belt

Chemistry-climate models project that the tropopause height will continue to increase in the future, but 

with a trend weaker than that in the recent past (Son et al., 

2009a; Gettelman et al., 2009). The reduced trend appears 

to be directly associated with stratospheric ozone recovery 

in the Southern Hemisphere in summer. With recovery 

of the ozone layer, chemistry-climate models also suggest 

that the observed trend toward widening of the tropics 

may weaken in the Southern Hemisphere in summer (Son 

et al., 2009a; Son et al., 2009b). However, our confidence 

in the magnitude of projected changes in the width of the 

tropics is limited by (1) discrepancies in such predictions 

derived from CMIP3 models and chemistry-climate models (Johanson and Fu, 2009; Son et al., 2009a) and (2) an 

incomplete understanding of the dynamics of the changes 

(Reichler, 2009).

4.5.5 Radiative Effects and Surface 

Temperature

By the middle of this century, global average concentrations of ozone-depleting substances may return 

to levels prior to 1960, but increasing concentrations 

of greenhouse gases are expected to alter stratospheric 

temperatures and circulation. Consequently, the ozone 

distribution will not simply return to its state prior to 

1960. The altered ozone concentrations will continue to 

exert an anomalous radiative forcing on surface climate. 

Portmann and Solomon (2007) computed an ozone forcing at 2100 from CO2, N2O, and CH4 changes under a 

range of IPCC Special Report on Emissions (i.e., SRES) 

scenarios. Although the induced ozone column changes 

were relatively large (up to a 5% change in the column, 

globally averaged) the overall radiative forcing since pre1970 ranged between −0.1 W/m2

 and +0.1 W/m2. This 

small forcing is a result of most of the associated ozone 

change occurring in the upper stratosphere, where it only 

contributed a small radiative forcing (see Section 4.2.2 

and Chapter 3).

4.5.6 Tropospheric Annular Modes and 

Stratosphere-Troposphere Coupling

As discussed in Section 4.4.2, only climate models that include observed changes in stratospheric ozone 

are able to reproduce the observed positive trend in the 

tropospheric Southern Annular Mode during austral summer over the past few decades. We infer that polar stratospheric ozone losses have been the dominant driver of 

the observed trend in the Southern Hemisphere surface 

circumpolar flow, as has been confirmed in those models 

that separately include ozone forcing with other radiative 

forcings held fixed (Gillett and Thompson, 2003; Shindell 

and Schmidt, 2004; Arblaster and Meehl, 2006).

In the absence of other forcings, stratospheric ozone 

recovery is expected to drive an equatorward shift in the 

Southern Hemisphere extratropical jet in the 21st century 

in austral summer that projects onto the negative polarity 

of the annular modes (e.g., Son et al., 2008). In contrast, 

future increases in atmospheric carbon dioxide concentrations are expected to lead to positive trends in the annular 

modes in both hemispheres (e.g., Fyfe et al., 1999; Kushner et al., 2001; Cai et al., 2003; Shindell and Schmidt, 

2004; Brandefelt and Källén, 2004; Yin, 2005; Miller et 

al., 2006; Arblaster and Meehl, 2006; Lu et al., 2008), 

although the amplitudes of the greenhouse gas-induced 

trends vary from simulation to simulation (Yin, 2005; 

Miller et al., 2006). The effects of ozone recovery and 

increasing carbon dioxide on the Southern Annular Mode 

are expected to approximately cancel through the middle 

of the 20th century during austral summer, as evidenced in 

both CCMVal-2 simulations and CMIP3 simulations (e.g., 

Figure 4-26; SPARC CCMVal, 20107). CMIP3 models 

without ozone recovery (Figure 4-26d) exhibit a weak 

positive Southern Annular Mode trend, which is presumably dominated by the greenhouse gas forcing.

There are several important caveats in the interpretation of annular mode trends. First, simulations of 20th

century climate tend to place the Southern Hemisphere jet 

farther equatorward than is observed (Fyfe and Saenko, 

2006; Son et al., 2010; Kidston and Gerber, 2010). The 

reasons for these biases are not well understood. But they 

seem to be important, as differences in the projected magnitude of the trend in the Southern Hemisphere are well 

correlated with biases in the latitude of the jet in the simulation of 20th century climate; models with a larger equatorward bias simulate a larger poleward shift in jet location, and vice versa (Son et al., 2010; Kidston and Gerber, 

7 Note that the CCMVal-2 results shown in Figure 4-26 differ 

somewhat from those based on the previous generation CCMVal-1 

models (Son et al., 2008; Perlwitz et al., 2008). Those runs indicate an equatorward shift of the jet at the surface in December, 

January, and February. 
4.44

Chapter 4

2010). This suggests that climate models may tend to 

overestimate the 21st century poleward shift in jet location 

(Kidston and Gerber, 2010).

An additional caveat involves the definition of the 

annular modes in the presence of a changing climate. Normally, the annular mode patterns are based on geopotential anomalies relative to a fixed climatology, such as for 

1960–2010. In long model integrations including climate 

change, the future annular mode index is calculated by 

projecting anomalies (relative to 1960–2010) onto the annular mode patterns. However, the projected widespread 

warming of the troposphere will lift geopotential heights 

at all latitudes in the free troposphere, with the largest 

increases in the tropics. Because the annular mode patterns emphasize changes over the polar caps, much of the 

projected changes to geopotential will not be reflected in 

the annular mode index. Thus, the annular mode patterns 

are not a good proxy for long-term projected changes to 

hemispheric geopotential (e.g., Fyfe, 2003; Gerber et al., 

2010). The zonal-mean response to anthropogenic forcing also appears to have a distinctive baroclinic signature 

that is very different from the annular modes (Woollings, 

2008). Shindell and Schmidt (2004) note a similar decoupling of the signatures of global warming and the Southern 

Annular Mode.

It is possible to modify the definition of the annular modes to compensate for the effects of tropospheric 

warming, but the resulting annular mode trends are not 

comparable to annular mode trends calculated using standard methodologies. For example, Morgenstern et al. 

(2010b) used a modified definition of the Northern Annular Mode (NAM) that partially accounts for the effects 

of tropospheric warming. With this modified definition, 

in the CCMVal-2 models, increasing CO2 is generally associated with trends towards the negative polarity of the 

NAM in winter. An alternate methodology is to remove 

the slowly varying trend in geopotential heights from the 

climatology, resulting in annular mode indices that have 

no long-term trends (the geopotential height trends are 

contained in the slowly varying climatology). Using this 

method, Gerber et al. (2010) found that the structure and 

month-to-month variability of the annular modes remain 

CMIP3

CMIP3 CMIP3

Figure 4-26. Trends in Decemberto-February (DJF) zonal-mean zonal 

wind. The multi-model mean trends 

between 2001 and 2050 are shown 

for the CCMVal-2 models (A), the 

CMIP3 models (B), the CMIP3 models with prescribed ozone recovery 

(C), and the CMIP3 models with no 

ozone recovery (D). Shading and 

contour intervals are 0.05 m s−1

decade−1. Deceleration and acceleration are indicated with blue and 

red colors, respectively, and trends 

weaker than 0.05 m s−1 decade−1

are omitted. Superimposed black 

solid lines are DJF zonal-mean 

zonal wind averaged from 2001 to 

2010, with a contour interval of 10 m 

s−1, starting at 10 m s−1. From Son 

et al. (2008, 2010).
4.45

Stratospheric Changes and Climate

relatively robust over time in CMIP3 and CCMVal-2 

model integrations of the 21st century (again, the trends in 

the mean of the annular mode indices are contained in the 

slowly varying climatology).

4.5.7 Tropospheric Chemistry

The future increases in stratospheric ozone will 

also impact tropospheric chemistry. The effects can be 

expected to be the reverse of the impacts of stratospheric 

ozone depletion over the past few decades discussed in 

Section 4.4.4. If stratospheric cooling leads to a thicker 

ozone layer than in the past (“super-recovery”) then future 

impacts may be larger. Decreasing UV radiation reaching 

the lower atmosphere is expected to decrease tropospheric 

OH. Duncan and Logan (2008) analyzed past CO trends. 

Based on their analysis, OH decreases due to ozone recovery would augment rather than cancel the positive impact 

of likely increasing CH4 on CO trends. In any case, based 

on calculations during the period of ozone loss (Section 

4.4.4), the impact is likely to be small but it remains to be 

quantified.

The flux of stratospheric ozone into the troposphere 

may also change in a future climate. Even with the same 

rate of mass transport, ozone increases in the lower stratosphere would lead to larger ozone transport into the troposphere. However, the stratosphere-troposphere transport 

mass flux may also change. Stevenson et al. (2006) studied the climate of 2030 using nine different models. They 

found ozone stratosphere-troposphere transport will increase by 0–19%. Zeng et al. (2008) investigated the impact of climate change on the tropospheric ozone budget 

using a 3-D chemistry-climate model with a relatively low 

lid and simple representation of stratospheric chemistry. 

For a 2100 atmosphere, they found that a doubling of CO2

caused an 80% increase in ozone transport from the stratosphere, which led to increases in tropospheric ozone down 

to the surface. Similar studies using low-lid tropospheric 

chemistry models by Collins et al. (2003) and Sudo et al. 

(2003) showed large stratosphere-troposphere transport 

increases by 2100 of 40% and 130%, respectively.

Recently, Hegglin and Shepherd (2009) investigated the impact of climate change on ozone in the lower 

atmosphere in a model with full stratospheric chemistry. 

Under a moderate greenhouse gas emissions scenario they 

found the stratosphere-troposphere transport of ozone 

increased by 23% between 1965 and 2095 due to changes 

in the Brewer-Dobson circulation, i.e., a smaller value 

than the low-lid tropospheric models cited above, but for a 

different greenhouse gas scenario. Hegglin and Shepherd 

(2009) also investigated the impact of the changes in vertical and latitudinal distribution of ozone on surface UV. 

Between 1965 and 2095 they found surface clear-sky UV 

fluxes decreased by 9% at northern high latitudes, increased by 4% in the tropics and increased by up to 20% at 

high southern latitudes. These changes in the clear-sky 

UV flux, due to circulation-induced changes in lowerstratospheric ozone, are different from those expected 

based on ozone recovery from halocarbon-induced depletion, and illustrate the importance of dynamics in controlling total ozone in the future as halocarbons decrease. 

Future UV levels will also be affected by changes in cloud 

cover (see Chapter 3, Section 3.4).

Future changes in the stratosphere-troposphere 

transport of ozone from ten chemistry-climate models 

have been assessed in Chapter 10 of SPARC CCMVal 

(2010). The report provided a comparison of the 

chemistry-climate model stratosphere-troposphere transport ozone flux with observations and other model studies 

for past and future time periods related to ozone depletion 

(1965–2000), ozone recovery (2000–2035), and climate 

change (1965–2095). The chemistry-climate model multimodel mean of the change in the stratosphere-troposphere 

transport ozone flux attributable to climate change (1965–

2095) was slightly larger in the Northern Hemisphere 

(17%) than in the Southern Hemisphere (14%). The 

chemistry-climate model multi-model mean was generally 

consistent in terms of relative changes with the Canadian 

model results shown in Hegglin and Shepherd (2009), but 

showed a reduction in Northern Hemisphere stratospheretroposphere transport ozone flux from 1965–2000 due to 

ozone depletion, as well as a weaker long-term increase in 

Northern Hemisphere stratosphere-troposphere transport 

ozone flux over 1965–2095 due to climate change. Over 

the period 2000–2030, the CCMVal-2 models showed 

increases in the global stratosphere-troposphere transport 

ozone flux of 10.3 (±2.0)% (see CCMVal, 2010, Chapter 

10, Figure 10-22), which is toward the upper end of the 

range from tropospheric models reported by Stevenson et 

al. (2006).

4.5.8 Solar and Volcanic Influences

The existing observational record is short, and 

hence we do not fully understand the limits of natural variability in the stratosphere. Based on our current understanding, the influence of explosive volcanic eruptions on 

stratospheric ozone depletion is linked to heterogeneous 

chemistry on the surface of particles, and is very sensitive 

to the amount of chlorine loading (Section 3.2.5). Hence 

potential future volcanic eruptions are expected to be associated with increasingly less ozone depletion as chlorine declines, and are even expected to result in ozone increases by midcentury (Tie and Brasseur, 1995). On the 

other hand, the direct effects of major volcanic eruptions 

on climate (including surface cooling, stratospheric warming, winter season weather patterns, and changes in ocean 

heat content and sea level rise) are expected to continue to 
4.46

Chapter 4

occur in a manner that scales with the stratospheric aerosol 

abundance.

Future stratospheric temperatures and ozone levels 

will depend on variations in the solar radiation incident at 

the Earth. The sun is currently in an anomalously quiescent state (Fröhlich, 2009), which makes prediction of its 

future behavior and total solar irradiance highly uncertain. 

Furthermore, recent space-based measurements (Harder 

et al., 2009) have cast doubt on currently accepted models of the ozone response to solar variability. That said, 

while the impact of future variations in solar irradiance 

is highly uncertain, it is unlikely that solar variability will 

have more than a minor effect on ozone levels relative to 

the effects of the reduction in chlorine loading.

REFERENCES

AchutaRao, K., and K.R. Sperber, ENSO simulation in 

coupled ocean-atmosphere models: Are the current 

models better?, Clim. Dyn., 27, 1-15, doi: 10.1007/

s00382-006-0119-7, 2006.

Allen, R.J., and S.C. Sherwood, Warming maximum in the 

tropical upper troposphere deduced from thermal 

winds, Nature Geosci., 1, 399-403, doi: 10.1038/

ngeo208, 2008.

Aoki, S., M. Yoritaka, and A. Masuyama, Multidecadal 

warming of subsurface temperature in the Indian 

sector of the Southern Ocean, J. Geophys Res., 108, 

8081, doi: 10.1029/2000JC000307, 2003.

Arblaster, J.M., and G.A. Meehl, Contributions of external 

forcings to Southern Annular Mode trends, J. Clim.,

19 (12), 2896-2905, doi: 10.1175/JCLI3774.1, 2006.

Archer, C.L., and K. Caldeira, Historical trends in the 

jet streams, Geophys. Res. Lett., 35, L08803, doi: 

10.1029/2008GL033614, 2008.

Austin, J., and F. Li, On the relationship between the 

strength of the Brewer–Dobson circulation and the 

age of stratospheric air, Geophys. Res. Lett., 33, 

L17807, doi: 10.1029/2006GL026867, 2006.

Austin, J., K. Tourpali, E. Rozanov, H. Akiyoshi, S. Bekki, G. Bodeker, C. Brühl, N. Butchart, M. Chipperfield, M. Deushi, V.I. Fomichev, M.A. Giorgetta, L. 

Gray, K. Kodera, F. Lott, E. Manzini, D. Marsh, K. 

Matthes, T. Nagashima, K. Shibata, R.S. Stolarski, 

H. Struthers, and W. Tian, Coupled chemistry-climate model simulations of the solar cycle in ozone 

and temperature, J. Geophys. Res., 113, D11306, 

doi: 10.1029/2007JD009391, 2008.

Austin, J., R.J. Wilson, H. Akiyoshi, S. Bekki, N. Butchart, 

C. Claud, V.I. Fomichev, P. Forster, R.R. Garcia, 

N.P. Gillett, P. Keckhut, U. Langematz, E. Manzini, T. Nagashima, W.J. Randel, E. Rozanov, K. Shibata, K.P. Shine, H. Struthers, D.W.J. Thompson, 

F. Wu, and S. Yoden, Coupled chemistry climate 

model simulations of stratospheric temperatures and 

their trends for the recent past, J. Geophys. Res., 36, 

L13809, doi: 10.1029/2009GL038462, 2009.

Baker, D.F., R.M. Law, K.R. Gurney, P. Rayner, P. Peylin, A.S. Denning, P. Bousquet, L. Bruhwiler, Y.-

H. Chen, P. Ciais, I.Y. Fung, M. Heimann, J. John, 

T. Maki, S. Maksyutov, K. Masarie, M. Prather, B. 

Pak, S. Taguchi, and Z. Zhu, TransCom 3 inversion intercomparison: Impact of transport model errors on the interannual variability of regional CO2

fluxes, 1988–2003, Global Biogeochem. Cycles,

20, GB1002, doi: 10.1029/2004GB002439, 2006.

Balachandran N., D. Rind, P. Lonergan, and D. Shindell, 

Effects of solar cycle variability on the lower stratosphere and the troposphere, J. Geophys. Res., 104, 

(D22) 27321-27339, 1999.

Baldwin, M., and M. Dameris (Lead Authors), 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, and D. Thompson, Climate-ozone connections, Chapter 5 in Scientific Assessment of Ozone Depletion: 2006, Global Ozone 

Research and Monitoring Project–Report No. 50, 

572 pp., World Meteorological Organization, Geneva, Switzerland, 2007.

Baldwin, M.P., and T.J. Dunkerton, Stratospheric harbingers of anomalous weather regimes, Science, 244 

(5542), 581-584, 2001.

Barriopedro, D., R. Garcia-Herrera, and R. Huth, Solar modulation of Northern Hemisphere winter 

blocking, J. Geophys. Res., 113, D14118, doi: 

10.1029/2008JD009789, 2008.

Bekki, S., K.S. Law, and J.A. Pyle, Effect of ozone depletion on atmospheric CH4 and CO concentrations, 

Nature, 371, 595-597, doi: 10.1038/371595a0, 1994.

Bengtsson, L., K. Hodges, and E. Roeckner, Storm tracks 

and climate change, J. Clim., 19 (15), 3518-3543, 

doi: 10.1175/JCLI3815.1, 2006.

Boberg, F., and H. Lundstedt, Solar wind variations related to 

fluctuations of the North Atlantic oscillation, Geophys. 

Res. Lett., 29, doi: 10.1029/2002GL014903, 2002.

Böning, C.W., A. Dispert, M. Visbeck, S.R. Rintoul, and 

F.U. Schwarzkopf, The response of the Antarctic Circumpolar Current to recent climate change, Nature 

Geosci., 1, 864-869, doi: 10.1038/ngeo362, 2008.

Bousquet, P., D.A. Hauglustaine, P. Peylin, C. Carouge, 

and P. Ciais, Two decades of OH variability as 

inferred by an inversion of atmospheric transport 

and chemistry of methyl chloroform, Atmos. Chem. 

Phys., 5, 2635-2656, doi: 10.5194/acp-5-2635-

2005, 2005.

Brandefelt, J., and E. Källén, The response of the Southern Hemisphere atmospheric circulation to an en-
4.47

Stratospheric Changes and Climate

hanced greenhouse gas forcing, J. Clim., 17 (22), 

4425-4442, doi: 10.1175/3221.1, 2004.

Brewer, A.W., Evidence for a world circulation provided 

by the measurements of helium and water vapour 

distribution in the stratosphere, Quart. J. Roy. 

Meteorol. Soc., 75 (326), 351-363, doi: 10.1002/

qj.49707532603, 1949.

Brönnimann, S., T. Ewen, T. Griesser, and R. Jenne, 

Multidecadal signal of solar variability in the upper troposphere during the 20th century, Space Sci. 

Rev., 125 (1-4), 305-315, doi: 10.1007/s11214-006-

9065-2, 2006.

Butchart, N., and A.A. Scaife, Removal of chlorofluorocarbons by increased mass exchange between the 

stratosphere and troposphere in a changing climate, 

Nature, 410, 799-802, 2001.

Butchart, N., A.A. Scaife, M. Bourqui, J. de Grandpré, 

S.H.E. Hare, J. Kettleborough, U. Langematz, E. 

Manzini, F. Sassi, K. Shibata, D. Shindell, and M. 

Sigmond, Simulations of anthropogenic change 

in the strength of the Brewer-Dobson circulation, 

Clim. Dyn., 27 (7-8), 727-741, 2006.

Butchart, N., I. Cionni, V. Eyring, T.G. Shepherd, D.W. 

Waugh, H. Akiyoshi, J. Austin, C. Brühl, M.P. 

Chipperfield, E. Cordero, M. Dameris, R. Deckert, 

S. Dhomse, S.M. Frith, R.R. Garcia, A. Gettelman, 

M.A. Giorgetta, D.E. Kinnison, F. Li, E. Mancini, 

C. McLandress, S. Pawson, G. Pitari, D.A. Plummer, E. Rozanov, F. Sassi, J.F. Scinocca, K. Shibata, B. Steil, and W. Tian, Chemistry-climate 

model simulations of 21st century stratospheric climate and circulation changes, J. Clim., 23 (20), doi: 

10.1175/2010JCLI3404.1, 2010.

Butler, A.H., D.W.J. Thompson, and K.R. Gurney, Observed relationships between the Southern Annular Mode and atmospheric carbon dioxide, 

Global Biogeochem. Cycles, 21, GB4014, doi: 

10.1029/2006GB002796, 2007. 

Cai, W., and T. Cowan, Trends in Southern Hemisphere circulation in IPCC AR4 models over 1950–99: Ozone 

depletion versus greenhouse forcing, J. Clim., 20

(4), 681-693, doi: 10.1175/JCLI4028.1, 2007.

Cai, W., P.H. Whetton, and D.J. Karoly, The response of the 

Antarctic oscillation to increasing and stabilized atmospheric CO2, J. Clim., 16 (10), 1525-1538, 2003.

Calvo, N., and R.R. Garcia, Wave forcing of the tropical 

upwelling in the lower stratosphere under increasing 

concentrations of greenhouse gases, J. Atmos. Sci., 66

(10), 3184-3196, doi: 10.1175/2009JAS3085.1, 2009.

Canziani, P.O., F.E. Malanca, and E.A. Agosta, Ozone 

and upper troposphere/lower stratosphere variability and change at southern midlatitudes 1980–2000: 

Decadal variations, J. Geophys. Res., 113, D20101, 

doi: 10.1029/2007JD009303, 2008.

CCSP (U.S. Climate Change Science Program), Temperature Trends in the Lower Atmosphere: Steps for 

Understanding and Reconciling Differences, edited 

by T.R. Karl, S.J. Hassol, C.D. Miller, and W.L. 

Murray, Synthesis and Assessment Product 1.1, A 

Report by the Climate Change Science Program 

and the Subcommittee on Global Change Research, 

164 pp., Washington, D.C., 2006.

Chapman, W.L., and J.E. Walsh, A synthesis of Antarctic temperatures, J. Clim., 20 (16), 4096-4117, doi: 

10.1175/JCLI4236.l, 2007.

Chin, M., and D. Davis, A reanalysis of carbonyl sulfide 

as a source of stratospheric background sulfur aerosol, J. Geophys. Res., 100 (D5), 8993-9005, 1995.

Christy, J.R., R.W. Spencer, W.B. Norris, W.D. Braswell, 

and D.E. Parker, Error estimates of version 5.0 of 

MSU-AMSU bulk atmospheric temperatures, J. Atmos. Oceanic Technol., 20 (5), 613-629, 2003.

Church, J. A., N.J. White, and J.M. Arblaster, Significant 

decadal-scale impact of volcanic eruptions on sea 

level and ocean heat content, Nature, 438, 74-77, 

doi: 10.1038/nature04237, 2005.

Clough, S., and M. Iacono, Line-by-line calculation of at , S., and M. Iacono, Line-by-line calculation of atmospheric fluxes and cooling rates. 2. Application 

to carbon dioxide, ozone, methane, nitrous oxide 

and the halocarbons, J. Geophys. Res., 100 (D8), 

16519-16535, 1995.

Collins, W.J., R.G. Derwent, B. Garnier, C.E. Johnson, 

M.G. Sanderson, and D.S. Stevenson, Effect of 

stratosphere-troposphere exchange on the future 

tropospheric ozone trend, J. Geophys. Res., 108 

(D12), 8528, doi: 10.1029/2002JD002617, 2003.

Cook, P.A., and H.K. Roscoe, Variability and trends in 

stratospheric NO2 in Antarctic summer, and implications for stratospheric NOy, Atmos. Chem. Phys.,

9, 3601-3612, doi: 10.5194/acp-9-3601-2009, 2009.

Cordero, E.C., and P.M.deF. Forster, Stratospheric variability and trends in models used for the IPCC 

AR4, Atmos. Chem. Phys., 6, 5369-5380, 2006.

Crook, J.A., N.P. Gillett, and S.P.E. Keeley, Sensitivity 

of Southern Hemisphere climate to zonal asymmetry in ozone, Geophys. Res. Lett., 35, L07806, doi: 

10.1029/2007GL032698, 2008.

Crutzen, P.J., The possible importance of CSO for the sulfate layer of the stratosphere, Geophys. Res. Lett., 3 

(2), 73-76, 1976.

Dall’Amico, M., L.J. Gray, K.H. Rosenlof, A.A. Scaife, 

K.P. Shine, and P.A. Stott, Stratospheric temperature trends: Impact of ozone variability and the 

QBO, Clim. Dyn., 34 (2-3), doi: 10.1007/s00382-

009-0604-x, 2010a.

Dall’Amico, M., P.A. Stott, A.A. Scaife, L.J. Gray, K.H. 

Rosenlof, and, A.Y. Karpechko, Impact of stratospheric variability on tropospheric climate change, 
4.48

Chapter 4

Clim. Dyn., 34 (2-3), 399-417, doi: 10.1007/s00382-

009-0580-1, 2010b.

Deckert, R., and M. Dameris, Higher tropical SSTs 

strengthen the tropical upwelling via deep convection, Geophys. Res. Lett., 35, L10813, doi: 

10.1029/2008GL033719, 2008.

Delworth, T.L., V. Ramaswamy, and G.L. Stenchikov, The 

impact of aerosols on simulated ocean temperature and 

heat content in the 20th century, Geophys. Res. Lett., 

32, L24709, doi: 10.1029/2005GL024457, 2005.

Deshler, T., A review of global stratospheric aerosol: 

Measurements, importance, life cycle, and local 

stratospheric aerosol, Atmos. Res., 90 (2-4), 223-

232, doi: 10.1016/j.atmosres.2008.03.016, 2008.

Dhomse, S., M. Weber, and J. Burrows, The relationship between tropospheric wave forcing and tropical lower stratospheric water vapor, Atmos. Chem. 

Phys., 8, 471-480, 2008.

Dlugokencky, E.J., L. Bruhwiler, J.W.C. White, L.K. Emmons, P.C. Novelli, S.A. Montzka, K.A. Masarie, 

P.M. Lang, A.M. Crotwell, J.B. Miller, and L.V. 

Gatti, Observational constraints on recent increases 

in the atmospheric CH4 burden, Geophys. Res. Lett.,

36, L18803, doi: 10.1029/2009GL039780, 2009.

Dobson, G.M.B., Origin and distribution of the polyatomic 

molecules in the atmosphere, Proc. Roy. Soc. London, 236, 1205, 187-193, 1956.

Domingues, C.M., J.A. Church, N.J. White, P.J. Gleckler, S.E. Wijffels, P.M. Barker, and J.R. Dunn, 

Improved estimates of upper-ocean warming and 

multi-decadal sea-level rise, Nature, 453, 1090-

1093, doi: 10.1038/nature07080, 2008.

Duncan, B.N., and J.A. Logan, Model analysis of the factors regulating the trends and variability of carbon 

monoxide between 1988 and 1997, Atmos. Chem. 

Phys., 8, 7389-7403, 2008.

Engel, A., T. Möbius, H. Bönisch, U. Schmidt, R. Heinz, 

I. Levin, E. Atlas, S. Aoki, T. Nakazawa, S. Sugawara, F. Moore, D. Hurst, J. Elkins, S. Schauffler, 

A. Andrews, and K. Boering, Age of stratospheric 

air unchanged within uncertainties over the past 

30 years, Nature Geosci., 2, 28-31, doi: 10.1038/

ngeo388, 2009.

Ern, M., and P. Preusse, Wave fluxes of equatorial Kelvin waves and QBO zonal wind forcing derived 

from SABER and ECMWF temperature space-time 

spectra, Atmos. Chem. Phys., 9, 3957-3986, 2009.

Esler, J.G., L.M. Polvani, and R.K. Scott, The Antarctic 

stratospheric sudden warming of 2002: A selftuned resonance?, Geophys. Res. Lett., 33, L12804, 

doi: 10.1029/2006GL026034, 2006.

Eyring, V., N. Butchart, D.W. Waugh, H. Akiyoshi, J. 

Austin, S. Bekki, G.E. Bodeker, B.A. Boville, C. 

Brühl, M.P. Chipperfield, E. Cordero, M. Dameris, 

M. Deushi, V.E. Fioletov, S.M. Frith, R.R. Garcia, 

A. Gettelman, M.A. Giorgetta, V. Grewe, L. Jourdain, D.E. Kinnison, E. Mancini, E. Manzini, M. 

Marchand, D.R. Marsh, T. Nagashima, P.A. Newman, J.E. Nielsen, S. Pawson, G. Pitari, D.A. Plummer, E. Rozanov, M. Schraner, T.G. Shepherd, K. 

Shibata, R.S. Stolarski, H. Struthers, W. Tian, and 

M. Yoshiki, Assessment of temperature, trace species, and ozone in chemistry-climate model simulations of the recent past, J. Geophys. Res., 111, 

D22308, doi: 10.1029/2006JD007327, 2006.

Eyring, V., A. Gettelman, N.R.P. Harris, S. Pawson, T.G. 

Shepherd, D.W. Waugh, H. Akiyoshi, N. Butchart, 

M.P. Chipperfield, M. Dameris, D.W. Fahey, 

P.M.deF. Forster, P.A. Newman, M. Rex, R.J. 

Salawitch, and B.D. Santer, Report on the Third 

SPARC CCMVal Workshop, SPARC Newsletter 

No. 30, 17-19, available: http://www.atmosp.physics.utoronto.ca/SPARC/Newsletter30Web/CCMVal.html, 2008.

Feck, T., J.-U. Grooβ, and M. Riese, Sensitivity of Arctic 

ozone loss to stratospheric H2O, Geophys. Res. Lett., 

35, L01803, doi: 10.1029/2007GL031334, 2008.

Fogt, R.L., A.J. Monaghan, J. Perlwitz, D.H. Bromwich, J.M. 

Jones, and G.J. Marshall, Historical SAM variability. 

Part II: Twentieth-century variability and trends from 

reconstructions, observations, and the IPCC AR4 

models, J. Clim., 22 (20), 5346-5355, 2009.

Forster, P.M.deF., and K.P. Shine, Assessing the climate impact of trends in stratospheric water vapor, Geophys. 

Res. Lett., 29, doi: 10.1029/2001GL013909, 2002.

Forster, P.M.deF., and M. Joshi, The role of halocarbons 

in the climate change of the troposphere and stratosphere, Clim. Change, 71 (1-2), 249-266, doi: 

10.1007/s10584-005-5955-7, 2005.

Forster, P., V. Ramaswamy, P. Artaxo, T. Berntsen, R. 

Betts, D.W. Fahey, J. Haywood, J. Lean, D.C. 

Lowe, G. Myhre, J. Nganga, R. Prinn, G. Raga, 

M. Schulz, and R. Van Dorland, Changes in atmospheric constituents and in radiative forcing, 

Chapter 2 in Climate Change 2007: The Physical 

Science Basis. Contribution of Working Group I 

to the Fourth Assessment Report of the Intergovernmental Panel on Climate Change, edited by S. 

Solomon, D. Qin, M. Manning, Z. Chen, M. Marquis, K.B. Averyt, M.Tignor, and H.L. Miller, 996 

pp., Cambridge University Press, Cambridge, U.K. 

and New York, NY, U.S.A., 2007a.

Forster, P.M., G. Bodeker, R. Schofield, S. Solomon, 

and D.W.J. Thompson, Effects of ozone cooling 

in the tropical lower stratosphere and upper troposphere, Geophys. Res. Lett., 34, L23813, doi: 

10.1029/2007GL031994, 2007b.

Frame, T.H.A., and L.J. Gray, The 11-year solar cycle in 
4.49

Stratospheric Changes and Climate

ERA-40 data: An update to 2008, J. Clim., 23 (8), 

doi: 10.1175/JCLI3150.1, 2213-2222, 2010.

Free, M., and D.J. Seidel, Comments on “Biases in stratospheric and tropospheric temperature trends derived 

from historical radiosonde data,” J. Clim., 20 (14), 

3704-3709, doi: 10.1175/JCLI4210.1, 2007.

Free, M., and J. Lanzante, Effect of volcanic eruptions on 

the vertical temperature profile in radiosonde data 

and climate models, J. Clim., 22 (11), 2925-2939, 

doi: 10.1175/2008JCLI2562.1, 2009.

Free, M., S.J. Seidel, J.K. Angell, J. Lanzante, I. Durre, and T.C. Peterson, Radiosonde Atmospheric 

Temperature Products for Assessing Climate 

(RATPAC): A new data set of large-area anomaly 

time series, J. Geophys. Res., 110, D22101, doi: 

10.1029/2005JD006169, 2005.

Fröhlich, C., Evidence of a long-term trend in solar irradiance, Astron. Astophys., 501 (3), L27-L30, doi: 

10.1051/0004-6361/200912318, 2009.

Fröhlich, C., and J. Lean, Solar radiative output and its 

variability: Evidence and mechanisms, Astron Astrophys Rev 12 (4), 273-320, doi: 10.1007/s00159-

004-0024-1, 2004.

Fu, Q., C.M. Johanson, S.G. Warren, and D.J. Seidel, 

Contribution of stratospheric cooling to satelliteinferred tropospheric temperature trends, Nature,

429, 55-58, doi: 10.1038/nature02524, 2004.

Fu, Q., C.M. Johanson, J.M. Wallace, and T. Reichler, 

Enhanced mid-latitude tropospheric warming in 

satellite measurements, Science, 312 (5777), 1179-

1179, doi: 10.1126/science.1125566, 2006.

Fu, Q., S. Solomon, and P. Lin, On the seasonal dependence of tropical lower-stratospheric temperature 

trends, Atmos. Chem. Phys., 10, 2643-2653, doi: 

10.5194/acp-10-2643-2010, 2010.

Fueglistaler, S., and P.H. Haynes, Control of interannual and longer-term variability of stratospheric 

water vapor, J. Geophys. Res., 110, D24108, doi: 

10.1029/2005JD006019, 2005.

Fueglistaler, S., A.E. Dessler, T.J. Dunkerton, I. Folkins, Q. Fu, and P.W. Mote, Tropical tropopause layer, Rev. Geophys., 47, RG1004, doi: 

10.1029/2008RG000267, 2009.

Fuglestvedt, J.S., J.E. Jonson, and I.S.A. Isaksen, Effects 

of reductions in stratospheric ozone on tropospheric 

chemistry through changes in photolysis rates, Tellus, 46B (3), 172-192, 1994.

Funatsu, B., C. Claud, P. Keckhut, and A. Hauchecorne, 

Cross-validation of Advanced Microwave Sounding Unit and lidar for long-term upper-stratospheric 

temperature monitoring, J. Geophys. Res., 113, 

D23108, doi: 10.1029/2008JD010743, 2008.

Fusco, A.C., and J.A. Logan, Analysis of 1970-1995 

trends in tropospheric ozone at Northern Hemisphere midlatitudes with the GEOS-CHEM model, J. Geophys. Res., 108, 4449, doi: 10.1029/ 

2002JD002742, 2003.

Fyfe, J.C., Separating extratropical zonal wind variability 

and mean change, J. Clim., 16 (5), 863-874, 2003.

Fyfe, J.C., Southern Ocean warming due to human influence, Geophys. Res. Lett., 33, L19701, doi: 

10.1029/2006GL027247, 2006.

Fyfe, J.C., and O.A. Saenko, Simulated changes in the 

extratropical Southern Hemisphere winds and 

currents, Geophys. Res. Lett., 33, L06701, doi: 

10.1029/2005GL025332, 2006.

Fyfe, J.C., G.J. Boer, and G.M. Flato, The Arctic and Antarctic oscillations and their projected changes under 

global warming, Geophys. Res. Lett., 26 (11), 1601-

1604, 1999.

Fyfe, J.C., O.A. Saenko, K. Zickfeld, M. Eby, and A.J. 

Weaver, The role of poleward-intensifying winds 

on Southern Ocean warming, J. Clim., 20 (21), 

5391-5400, doi: 10.1175/2007JCLI1764.1, 2007.

Garcia, R.R., and W.J. Randel, Acceleration of the Brewer-Dobson circulation due to increases in greenhouse gases, J. Atmos. Sci., 65 (8), 2731-2739, doi: 

10.1175/2008JAS2712.1, 2008.

Geller, M.A., X. Zhou, and M. Zhang, Simulations of the 

interannual variability of stratospheric water vapor, 

J. Atmos. Sci., 59 (6), 1076-1085, 2002.

Gent, P.R., and J.C. McWilliams, Isopycnal mixing in 

ocean circulation models, J. Phys. Ocean., 20 (1), 

150-155, 1990.

Gerber, E.P., S. Voronin, and L.M. Polvani, Testing the annular mode autocorrelation timescale in simple atmospheric general circulation 

models, Mon. Wea. Rev., 136, 1523-1536, doi: 

10.1175/2007MWR2211.1, 2008.

Gerber, E.P., M.P. Baldwin, H. Akiyoshi, J. Austin, S. 

Bekki, P. Braesicke, N. Buchart, M. Chipperfield, 

M. Dameris, S. Dhomse, S.M. Firth, R.R. Garcia, 

H. Garney, A. Gettelman, S.C. Hardiman, O. Morgenstern, J.E. Nielsen, S. Pawson, T. Peter, D.A. 

Plummer, J.A. Pyle, E. Rozanov, J.F. Scinocca, 

T.G. Shepherd, and D. Smale, Stratosphere-troposphere coupling and annular mode variability in 

chemistry-climate models, J. Geophys. Res., 115, 

D00M06, doi: 10.1029/2009JD013770, 2010.

Gettelman, A., T. Birner, V. Eyring, H. Akiyoshi, S. Bekki, C. Brühl, M. Dameris, D.E. Kinnison, F. Lefevre, F. Lott, E. Mancini, G. Pitari, D.A. Plummer, 

E. Rozanov, K. Shibata, A. Stenke, H. Struthers, 

and W. Tian, The tropical tropopause layer 1960-

2100, Atmos. Chem. Phys., 9, 1621-1637, 2009.

Gettelman, A., M.I. Hegglin, S.-W. Son, J. Kim, M. Fujiwara, T. Birner, S. Kremser, M. Rex, J.A. Añel, 

H. Akiyoshi, J. Austin, S. Bekki, P. Braesike, C. 
4.50

Chapter 4

Brühl, N. Butchart, M. Chipperfield, M. Dameris, 

S. Dhomse, H. Garny, S.C. Hardiman, P. Jöckel, 

D.E. Kinnison, J.F. Lamarque, E. Mancini, M. 

Marchand, M. Michou, O. Morgenstern, S. Pawson, G. Pitari, D. Plummer, J.A. Pyle, E. Rozanov, 

J. Scinocca, T.G. Shepherd, K. Shibata, D. Smale, 

H. Teyssèdre, and W. Tian, Multimodel assessment 

of the upper troposphere and lower stratosphere: 

Tropics and global trends, J. Geophys. Res., 115, 

D00M08, doi: 10.1029/2009JD013638, 2010.

Gille, S.T., Warming of the Southern Ocean since the 

1950s, Science, 295 (5558), 1275-1277, doi: 

10.1126/science.1065863, 2002.

Gille, S.T., Decadal-scale temperature trends in the Southern Hemisphere ocean, J. Clim., 21 (18), 4749-

4765, doi: 10.1175/2008JCLI2131.1, 2008.

Gillett, N.P., and D.W.J. Thompson, Simulation of recent 

Southern Hemisphere climate change, Science, 302 

(5643), 273-275, 2003.

Gillett, N.P., T.D. Kell, and P.D. Jones, Regional climate impacts of the Southern Annular Mode, Geophys. Res. 

Lett., 33, L23704, doi: 10.1029/2006GL027721, 2006.

Gillett, N.P., D.A. Stone, P.A. Stott, T. Nozawa, A.Yu. 

Karpechko, G.C. Hegerl, M.F. Wehner, and P.D. 

Jones, Attribution of polar warming to human influence, Nature Geosci., 1, 750-754, doi: 10.1038/

ngeo338, 2008.

Gillett, N.P., J.F. Scinocca, D.A. Plummer, and M.C. Reader, Sensitivity of climate to dynamically-consistent 

zonal asymmetries in ozone, Geophys. Res. Lett., 

36, L10809, doi: 10.1029/2009GL037246, 2009.

Gillett, N.P., H. Akiyoshi , S. Bekki , V. Eyring , R. Garcia, 

C.A. McLinden, A. Yu. Karpechko, D.A. Plummer, 

E. Rozanov, J. Scinocca, and K. Shibata, Attribution 

of observed changes in stratospheric ozone and temperature, Atmos. Chem. Phys. Discuss., 10, 17341-

17367, doi: 10.5194/acpd-10-17341-2010, 2010.

Giorgetta, M.A., E. Manzini, E. Roeckner, M. Esch, 

and L. Bengtsson, Climatology and forcing of the 

quasi-biennial oscillation in the MAECHAM5 

model, J. Clim., 19 (16), 3882-3901, doi: 10.1175/

JCLI3830.1, 2006.

Gleckler, P.J., K. AchutaRao, J.M. Gregory, B.D. Santer, 

K.E. Taylor, and T.M.L. Wigley, Krakatoa lives: 

The effect of volcanic eruptions on ocean heat content and thermal expansion, Geophys. Res. Lett., 33, 

L17702, doi: 10.1029/2006GL026771, 2006.

Goosse, H., W. Lefebvre, A. de Montety, E. Crespin, 

and A.H. Orsi, Consistent past half-century trends 

in the atmosphere, the sea ice, and the ocean at 

high southern latitudes, Clim. Dyn., 33 (7-8), doi: 

10.1007/s00382-008-0500-9, 2009.

Granier, C., J.-F. Müller, S. Madronich, and G.P. Brasseur, Possible causes for the 1990-1993 decrease in 

the global tropospheric CO abundances: A threedimensional sensitivity study, Atmos. Env., 30 (10-

11), 1673-1682, doi: 10.1016/1352-2310(95)00397-

5, 1996.

Gray, L.J., S.T. Rumbold, and K. Shine, Stratospheric 

temperature and radiative forcing response to 

11-year solar cycle changes in irradiance and 

ozone, J. Atmos. Sci., 66 (8), 2402-2417, doi: 

10.1175/2009JAS2866.1, 2009.

Gray, L.J., J. Beer, M. Geller, J.D. Haigh, M. Lockwood, 

K. Matthes, U. Cubasch, D. Fleitmann, R.G. Harrison, L. Hood, J. Luterbacher, G.A. Meehl, D. Shindell, B. van Geel, and W. White, Solar influence on 

climate, Rev. Geophys., in press, 2010.

Gregory, J.M., J.A. Lowe, and S.F.B. Tett, Simulated 

global-mean sea level change over the last halfmillennium, J. Clim., 19 (18), 4576-4591, doi: 

10.1175/JCLI3881.1, 2006.

Griffiths, G.M., Changes in New Zealand daily rainfall extremes 1930–2004, Weather and Climate, 26, 3-46, 

2006.

Grise, K.M., D.W.J. Thompson, and P.M. Forster, On the 

role of radiative processes in stratosphere-troposphere coupling, J. Clim., 22 (15), 4154-4161, doi: 

10.1175/2009JCLI2756.1, 2009.

Haigh, J.D., The role of stratospheric ozone in modulating 

the solar radiative forcing of climate, Nature, 370 

(6490), 544-546, 1994.

Haigh, J.D., The impact of solar variability on climate, 

Science, 272 (5264), 981-984, doi: 10.1126/science272.5264.981, 1996.

Haigh, J.D., A GCM study of climate change in response 

to the 11-year solar cycle, Quart. J. Roy. Meteorol. 

Soc., 125 (555), 871-892, 1999.

Haigh, J.D., The effects of solar variability on the Earth’s 

climate, Phil. Trans. Roy. Soc. A, 361 (1802), 95-

111, 2003.

Haigh, J.D., and J.A. Pyle, Ozone perturbation experiments 

in a two-dimensional circulation model, Quart. J. 

Roy. Meteorol. Soc., 108, 551-574, 1982.

Haigh, J.D., and H.K. Roscoe, Solar influences on polar 

modes of variability, Meteorol. Z., 15 (3), 371-

378, 2006.

Haigh, J.D., and H.K. Roscoe, The final warming date 

of the Antarctic polar vortex and influences on its 

interannual variability, J. Clim., 22 (22), 5809-

5819, 2009.

Haigh, J.D., M. Blackburn, and R. Day, The response 

of tropospheric circulation to perturbations in 

lower-stratospheric temperature, J. Clim., 18, 

3672-3685, 2005.

Haigh, J.D., A.R. Winning, R. Toumi, and J.W. Harder, 

An influence of solar spectral variations on radiative forcing of climate, Nature, 467, 696-699, doi: 
4.51

Stratospheric Changes and Climate

10.1038/nature09426, 2010.

Haimberger, L., Homogenization of radiosonde temperature time series using innovation statistics, J. Clim., 20 (7), 1377-1403, doi: 10.1175/

JCLI4050.1, 2007.

Haimberger, L., C. Tavolato, and S. Sperka, Toward 

elimination of the warm bias in historic radiosonde 

temperature records—Some new results from 

a comprehensive intercomparison of upper-air 

data, J. Clim., 21 (18), 4587-4606, doi: 10.1175/ 

2008JCLI1929.1, 2008.

Hall, A., and M. Visbeck, Synchronous variability in the Southern Hemisphere atmosphere, 

sea ice, and ocean resulting from the Annular Mode, J. Clim., 15 (21), 3043-3057, doi: 

10.1175/1520-0442(2002)015<3043:SVITSH>2.0.

CO;2, 2002.

Hallberg, R., and A. Gnanadesikan, The role of eddies in 

determining the structure and response of the winddriven Southern Hemisphere overturning: Results 

from the Modeling Eddies in the Southern Ocean 

(MESO) project, J. Phys. Ocean., 36 (12), 2232-

2252, doi: 10.1175/JPO2980.1, 2006.

Hampson J., P. Keckhut, A. Hauchecorne, and M.L. Chanin, The effect of the 11-year solar-cycle on the 

temperature in the upper-stratosphere and mesosphere: Part II: Numerical simulation and role of 

planetary waves, J. Atmos. Sol. Terr. Phys., 67, 

948-958, 2005.

Hansen, J., A. Lacis, R. Ruedy, and M. Sato, Potential climate impact of Mount Pinatubo eruption, Geophys. Res. Lett., 19 (2), 215-218, doi: 

10.1029/91GL02788, 1992.

Harder, J.W., J.M. Fontenla, P. Pilewskie, E.C. Richard, and 

T.N. Woods, Trends in solar spectral irradiance variability in the visible and infrared, Geophys. Res. Lett.,

36, L07801, doi: 10.1029/2008GL036797, 2009.

Hegerl, G.C., F.W. Zwiers, P. Braconnot, N.P. Gillett, Y. 

Luo, J.A. Marengo Orsini, N. Nicholls, J.E. Penner, 

and P.A. Stott, Understanding and attributing climate change, Chapter 9 in Climate Change 2007: 

The Physical Science Basis. Contribution of Working Group I to the Fourth Assessment Report of the 

Intergovernmental Panel on Climate Change, edited by S. Solomon, D. Qin, M. Manning, Z. Chen, 

M. Marquis, K.B. Averyt, M. Tignor, and H.L. 

Miller, 996 pp., Cambridge University Press, Cambridge, U.K. and New York, NY, U.S.A., 2007.

Hegglin, M.I., and T.G. Shepherd, Large climate-induced 

changes in ultraviolet index and stratosphere-totroposphere ozone flux, Nature Geosci., 2, 687-

691, doi: 10.1038/ngeo604, 2009.

Hendon, H.H., D.W.J. Thompson, and M.C.Wheeler, Australian rainfall and surface temperature variations 

associated with the Southern Hemisphere Annular 

Mode, J. Clim., 20 (11), 2452-2467, doi: 10.1175/

JCLI4134.1, 2007.

Hofmann, D.J., Increase in the stratospheric background 

sulfuric acid aerosol mass in the past 10 years, 

Science, 248 (4958), 996-1000, doi: 10.1126/science.248.4958.996, 1990.

Hofmann, D.J., J. Barnes, M. O’Neill, M. Trudeau, and R. 

Neely, Increase in background stratospheric aerosol observed with lidar at Mauna Loa Observatory 

and Boulder, Colorado, Geophys. Res. Lett., 36, 

L15808, doi: 10.1029/2009GL039008, 2009.

Holton, J.R., and A. Gettelman, Horizontal transport and 

the dehydration of the stratosphere, Geophys. Res.

Lett., 28 (14), 2799-2802, 2001.

Hsu J., and M.J. Prather, Stratospheric variability and tropospheric ozone, J. Geophys. Res., 114, D06102, 

doi: 10.1029/2008JD010942, 2009.

Hu, Y., and Q. Fu, Observed poleward expansion of the 

Hadley circulation since 1979, Atmos. Chem. Phys.,

7, 5229-5236, 2007.

Hudson, R.D., M.F. Andrade, M.B. Follette, and A.D. 

Frolov, The total ozone field separated into meteorological regimes—Part II: Northern Hemisphere 

mid-latitude total ozone trends, Atmos. Chem. 

Phys., 6, 5183-5191, 2006.

IPCC, Climate Change 2007: The Physical Science Basis. 

Contribution of Working Group I to the Fourth Assessment Report of the Intergovernmental Panel on 

Climate Change, edited by S. Solomon, D. Qin, M. 

Manning, Z. Chen, M. Marquis, K.B. Averyt, M. 

Tignor, and H.L. Miller, 996 pp., Cambridge University Press, Cambridge, U.K. and New York, NY, 

U.S.A., 2007.

IPCC/TEAP (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, prepared by Working Groups I 

and III of the Intergovernmental Panel on Climate 

Change, and the Technical and Economic Assessment Panel, edited by B. Metz, L. Kuijpers, S. Solomon, S.O. Andersen, O. Davidson, J. Pons, D. de 

Jager, T. Kestin, M. Manning, and L. Meyer, 488 

pp., Cambridge University Press, Cambridge, U.K. 

and New York, NY, U.S.A., 2005.

Isaksen, I.S.A., C. Zerefos, K. Kourtidis, C. Meleti, S.B. 

Dalsøren, J.K. Sundet, A. Grini, P. Zanis, and D. 

Balis, Tropospheric ozone changes at unpolluted 

and semipolluted regions induced by stratospheric 

ozone changes, J. Geophys. Res., 110, D02302, doi: 

10.1029/2004JD004618, 2005.

Isaksen, I., C. Grainer, G. Myhre, T.K. Berntsen, S.B. Dal-
4.52

Chapter 4

seren, M. Gauss, Z. Klimont, R. Benestad, P. Bousquet, W. Collins, T. Cox, V. Eyring, D. Fowler, S. 

Fuzzi, P. Jöckel, P. Laj, U. Lohmann, M. Malone, 

P. Monks, A.S.H. Prevot, F. Raes, A. Richter, B. 

Rognerud, M. Schulz, D. Shindell, D.S. Stevenson, T. Storelvmo, W.-C. Wang, M. van Weele, M. 

Wild, and D. Wuebbles, Atmospheric composition 

change: Climate-Chemistry interactions, Atmos. 

Env., 43 (33), 5138-5192, doi: 10.1016/j.atmosenv.2009.08.003, 2009.

Ito, T., M. Woloszyn, and M. Mazloff, Anthropogenic carbon dioxide transport in the Southern Ocean driven 

by Ekman flow, Nature, 463, 80-83, doi: 10.1038/

natureo8687, 2010.

Iwasaki, T., H. Hamada, and K. Miyazaki, Comparisons of 

Brewer-Dobson circulations diagnosed from reanalyses, J. Meteor. Soc. Jap., 87 (6), 997-1006, 2009.

Jäger, H., Long-term record of lidar observations of the 

stratospheric aerosol layer at Garmisch-Partenkirchen, J. Geophys. Res., 110, D08106, doi: 

10.1029/2004JD005506, 2005.

Jensen, E., and L. Pfister, Transport and freeze-drying in 

the tropical tropopause layer, J. Geophys. Res., 109, 

D02207, doi: 10.1029/2003JD004022, 2004.

Johanson, C.M., and M. Fu, Hadley cell widening: Model 

simulations versus observations, J. Clim., 22 (10), 

2713-2725, doi: 10.1175/2008JCLI2620.1, 2009.

Jones, R.L., J.A. Pyle, J.E. Harries, A.M. Zavody, J.M. Russell III, and J.C. Gille, The water vapour budget of 

the stratosphere studied using LIMS and the SAMS 

satellite data, Quart. J. Roy. Meteor. Soc., 112 (474), 

1127-1143, doi: 10.1256/smsqj.47411, 1986.

Jonsson, A.I., V.I. Fomichev, and T.G. Shepherd, The 

effect of nonlinearity in CO2 heating rates on the 

attribution of stratospheric ozone and temperature 

changes, Atmos. Chem. Phys., 9, 8447-8452, 2009.

Joshi, M.M., and G.S. Jones, The climatic effects of the direct injection of water vapour into the stratosphere 

by large volcanic eruptions, Atmos. Chem. Phys., 9, 

6109-6118, 2009.

Karpechko, A.Yu., N.P. Gillett, G.J. Marshall, and A.A. 

Scaife, Stratospheric influence on circulation 

changes in the Southern Hemisphere troposphere 

in coupled climate models, Geophys. Res. Lett., 35, 

L20806, doi: 10.1029/2008GL035354, 2008.

Karpetchko, A., E. Kyrö, and B.M. Knudsen, Arctic and 

Antarctic polar vortices 1957-2002 as seen from the 

ERA-40 reanalyses, J. Geophys. Res., 110, D21109, 

doi: 10.1029/2005JD006113, 2005.

Keckhut, P., C. Cagnazzo, M.-L. Chanin, C. Claud, and 

A. Hauchecorne, The 11-year solar-cycle effects on 

the temperature in the upper-stratosphere and mesosphere: Part I–Assessment of observations, J. Atmos. Sol. Terr. Phys , 67, 940-947, doi: 10.1016/j.

jastp.2005.01.008, 2005.

Keeley, S.P.E., N.P. Gillett, D.W.J. Thompson, S. Solomon, and P.M. Forster, Is Antarctic climate most 

sensitive to ozone depletion in the mid or lower 

stratosphere?, Geophys. Res. Lett., 34, L22812, doi: 

10.1029/2007GL031238, 2007.

Khatiwala, S., F. Primeau, and T. Hall, Reconstruction of 

the history of anthropogenic CO2 concentrations in 

the ocean, Nature, 462, 346-349, doi: 10.1038/nature08526, 2009.

Khaykin, S., J.-P. Pommereau, L. Korshunov, V. Yushkov, J. Nielsen, N. Larsen, T. Christensen, A. Garnier, A. Lukyanov, and E. Williams, Hydration of 

the lower stratosphere by ice crystal geysers over 

land convective systems, Atmos. Chem. Phys., 9, 

2275-2287, 2009.

Kidston, J., and E.P. Gerber, Intermodel variability of 

the poleward shift of the austral jet stream in the 

CMIP3 integrations linked to biases in 20th century 

climatology, Geophys. Res. Lett., 37, L09708, doi: 

10.1029/2010GL042873, 2010.

Kirk-Davidoff, D.B., E.J. Hintsa, J.G. Anderson, and D.W. 

Keith, The effect of climate change on ozone depletion through changes in stratospheric water vapour, 

Nature, 402, 399-401, doi: 10.1038/46521, 1999.

Kley, D., J.M. Russell III, and C. Phillips, SPARC Assessment of Upper Tropospheric and Stratospheric Water Vapor, WCRP-No. 113, WMO/TD-No. 1043, 

SPARC Report No.2, 2000.

Kodera, K., Solar cycle modulation of the North Atlantic 

Oscillation: Implication in the spatial structure of 

the NAO, Geophys. Res. Lett., 29 (8), 1218, doi: 

10.1029/2001GL014557, 2002.

Kodera, K., Solar influence on the Indian Ocean Monsoon 

through dynamical processes, Geophys. Res. Lett.,

31, L24209, doi: 10.1029/2004GL020928, 2004.

Kodera, K., and K. Shibata, Solar influence on the tropical stratosphere and troposphere in the northern 

summer, Geophys. Res. Lett., 33, L19704, doi: 

10.1029/2006GL026659, 2006.

Kodera, K., K. Coughlin, and O. Arakawa, Possible 

modulation of the connection between the Pacific and Indian Ocean variability by the solar cycle, Geophys. Res. Lett., 34, L03710, doi: 

10.1029/2006GL027827, 2007.

Korhonen, H., K.S. Carslaw, P.M. Forster, S. Mikkonen, 

N.D. Gordon, and H. Kokkola, Aerosol climate 

feedback due to decadal increases in Southern 

Hemisphere wind speeds, Geophys. Res. Lett., 37, 

L02805, doi: 10.1029/2009GL041320, 2010.

Kuroda, Y., and K. Kodera, Role of the Polar-night Jet Oscillation on the formation of the Arctic Oscillation in 

the Northern Hemisphere winter, J. Geophys. Res.,

109, D11112, doi: 10.1029/2003JD004123, 2004.
4.53

Stratospheric Changes and Climate

Kuroda, Y., and K. Kodera, Solar cycle modulation of the 

Southern Annular Mode, Geophys. Res. Lett., 32, 

L13802, doi: 10.1029/2005GL022516, 2005.

Kuroda, Y., M. Deushi, and K. Shibata, Role of solar activity in the troposphere-stratosphere coupling in the 

Southern Hemisphere winter, Geophys. Res. Lett., 

34, L21704, doi: 10.1029/2007GL030983, 2007.

Kushner, P.J., and L.M. Polvani, Stratosphere-troposphere 

coupling in a relatively simple AGCM: The role of 

eddies, J. Clim., 17 (3), 629-639, 2004.

Kushner, P.J., and L.M. Polvani, Stratosphere-troposphere coupling in a relatively simple AGCM: 

Impact of the seasonal cycle, J. Clim., 19 (21), 

5721-5727, 2006.

Kushner, P., I.M. Held, and T.L. Delworth, Southern Hemisphere atmospheric circulation response to global 

warming, J. Clim., 14 (10), 2238-2249, 2001.

Labitzke, K., Sunspots, the QBO and the stratospheric 

temperature in the north polar region, Geophys. 

Res. Lett., 14 (5), 535-537, 1987.

Lamarque, J.-F., and S. Solomon, Impact of changes in climate and halocarbons on recent lower stratosphere 

ozone and temperature trends, J. Clim., 23 (10), 

2599-2611, doi: 10.1175/2010JCLI3179.1, 2010.

Lambert, A., W.G. Read, N.J. Livesey, M.L. Santee, G.L. 

Manney, L. Froidevaux, D.L. Wu, M.J. Schwartz, 

H.C. Pumphrey, C. Jimenez, G.E. Nedoluha, R.E. 

Cofield, D.T. Cuddy, W.H. Daffer, B.J. Drouin, 

R.A. Fuller, R.F. Jarnot, B.W. Knosp, H.M. Pickett, 

V.S. Perun, W.V. Snyder, P.C. Stek, R.P. Thurstans, 

P.A. Wagner, J.W. Waters, K.W. Jucks, G.C. Toon, 

R.A. Stachnik, P.F. Bernath, C.D. Boone, K.A. 

Walker, J. Urban, D. Murtagh, J.W. Elkins, and 

E. Atlas, Validation of the Aura Microwave Limb 

Sounder middle atmosphere water vapor and nitrous oxide measurements, J. Geophys. Res., 112, 

D24S36, doi: 10.1029/2007JD008724, 2007.

Lanzante, J.R., S.A. Klein, and D.J. Seidel, Temporal homogenization of monthly radiosonde temperature 

data. Part I: Methodology, J. Clim., 16 (2), 224-

240, 2003.

Larkin, A., J.D. Haigh, and S. Djavidnia, The effect of solar UV irradiance variations on the Earth’s atmosphere, Space Sci. Rev., 94 (1-2), 199-214, 2000.

Law, K.S., and W.T. Sturges (Lead Authors), 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, and S.A. Yvon-Lewis, Halogenated 

very short-lived substances, Chapter 2 in Scientific Assessment of Ozone Depletion: 2006, Global 

Ozone Research and Monitoring Project–Report 

No. 50, 572 pp., World Meteorological Organization, Geneva, Switzerland, 2007.

Law, R.M., R.J. Matear, and R.J. Francey, Comment On 

“Saturation of the southern ocean CO2 sink due 

to recent climate change,” Science, 319 (5863), 

570, 2008.

Le Quéré, C., C. Rödenbeck, E.T. Buitenhuis, T.J. Conway, 

R. Langenfelds, A. Gomez, C. Labuschagne, M. Ramonet, T. Nakazawa, N. Metzl, N. Gillett, and M. 

Heimann, Saturation of the Southern Ocean CO2 sink 

due to recent climate change, Science, 316 (5832), 

1735-1738, doi: 10.1126/science.1136188, 2007.

Le Quéré, C., M.R. Raupach, J.G. Canadell, and G. Marland, Trends in the sources and sinks of carbon dioxide, Nature Geosci., 2, 831-836, doi: 10.1038/

ngeo689, 2009.

le Texier, H., S. Solomon, and R.R. Garcia, The role of 

molecular hydrogen and methane oxidation in the 

water vapour budget of the stratosphere, Quart.

J. Roy. Meteorol. Soc., 114 (480), 281-295, doi: 

10.1002/qj.49711448002, 1988.

Lee, H., and A.K. Smith, Simulation of the combined effects of solar cycle, quasi-biennial oscillation, and 

volcanic forcing on stratospheric ozone changes in 

recent decades, J. Geophys. Res., 108 (D2), 4049, 

doi: 10.1029/2001JD001503, 2003.

Lee, J.N., and S. Hameed, Northern Hemisphere annular 

mode in summer: Its physical significance and its 

relation to solar activity variations, J. Geophys. Res., 

112, D15111, doi: 10.1029/2007JD008394, 2007.

Lee, J.N., S. Hameed, and D.T. Shindell, The northern 

annular mode in summer and its relation to solar 

activity variations in the GISS ModelE, J. Atmos. 

Sol. Terr. Phys., 70 (5), 730-741, doi: 10.1016/j.

jastp.2007.10.012, 2008.

Leith, C.E., Climate response and fluctuation dissipation, 

J. Atmos. Sci., 32, 2022-2026, 1975.

Lenton, A., and R.J. Matear, Role of the Southern Annular Mode (SAM) in Southern Ocean CO2 uptake, 

Global Biogeochem. Cycles, 21, GB2016, doi: 

10.1029/2006GB002714, 2007.

Lenton, A., F. Codron, L. Bopp, N. Metzl, P. Cadule, A. 

Tagliabue, and J. Le Sommer, Stratospheric ozone 

depletion reduces ocean carbon uptake and enhances ocean acidification, Geophys. Res. Lett., 36, 

L12606, doi: 10.1029/2009GL038227, 2009.

Li, F., J. Austin, and J. Wilson, The strength of the BrewerDobson circulation in a changing climate: Coupled 

chemistry-climate model stimulation, J. Clim., 21

(1), 40-57, doi: 10.1175/2007.JCLI1663.1, 2008.

Limpasuvan, V., and D.L. Hartmann, Wave-maintained 

annular modes of climate variability, J. Clim., 13, 

4414-4429, 2000.

Lin, P., Q. Fu, S. Solomon, and J.M. Wallace, Temperature trend patterns in Southern Hemisphere high 
4.54

Chapter 4

latitudes: Novel indicators of stratospheric change, 

J. Clim., 22 (23), 2009.

Lorenz, D.J., and E.T. DeWeaver, Tropopause height 

and zonal wind response to global warming in the 

IPCC scenario integrations, J. Geophys. Res., 112, 

D10119, doi: 10.1029/2006JD008087, 2007.

Lovenduski, N.S., N. Gruber, and S.C. Doney, Toward a mechanistic understanding of the decadal trends in the Southern Ocean carbon sink, 

Global Biogeochem. Cycles, 22, GB3016, doi: 

10.1029/2007GB003139, 2008.

Lu, J., G.A. Vecchi, and T. Reichler, Expansion of 

the Hadley cell under global warming, Geophys. Res. Lett., 34, L06805, doi: 10.1029/ 

2006GL028443, 2007.

Lu, J., G. Chen, and D.M.W. Frierson, Response of the 

zonal mean atmospheric circulation to El Niño versus global warming, J. Clim., 21 (22), 5835-5851, 

doi: 10.1175/2008JCLI2200.1, 2008.

Lu, J., C. Deser, and T. Reichler, Cause of the widening of 

the tropical belt since 1958, Geophys. Res. Lett., 36, 

L03803, doi: 10.1029/2008GL036076, 2009.

Lynch, A., P. Uotila, and J.J. Cassano, Changes in synoptic weather patterns in the polar regions in the twentieth and twenty-first centuries, Part 2: Antarctic, 

Int. J. Climatol., 26 (9), 1181-1199, 2006.

Madronich, S., and C. Granier, Impact of recent total ozone 

changes on tropospheric ozone photodissociation, 

hydroxyl radicals, and methane trends, Geophys. 

Res. Lett., 19 (5), 465-467, 1992.

Marshall, G.J., A. Orr, N.P.M. van Lipzig, and J.C. King, 

The impact of a changing Southern Hemisphere 

Annular Mode on Antarctic peninsula summer 

temperatures, J. Clim., 19 (20), 5388-5404, doi: 

10.1175/JCLI3844.1, 2006.

Matthes, K., Y. Kuroda, K. Kodera, and U. Langematz, 

Transfer of the solar signal from the stratosphere to 

the troposphere: Northern winter, J. Geophys. Res.,

111, D06108, doi: 10.1029/2005JD006283, 2006.

McLandress, C., and T.G. Shepherd, Simulated anthropogenic changes in the Brewer-Dobson 

circulation, including its extension to high 

latitudes, J. Clim., 22 (6), 1516-1540, doi: 

10.1175/2008JCLI2679.1, 2009.

McLandress, C., A.I. Jonsson, D.A. Plummer, M.C. 

Reader, J.F. Scinocca, and T.G. Shepherd, Separating the dynamical effects of climate change 

and ozone depletion: Part 1. Southern Hemisphere 

stratosphere, J. Clim., 23 (18), 5002-5020, doi: 

10.1175/2010JCLI3586.1, 2010.

Mears, C.A., and F.J. Wentz, Construction of the Remote 

Sensing Systems V3.2 atmospheric temperature records from the MSU and AMSU microwave sounders, J. Atmos. Oceanic Technol., 26 (6), 1040-1056, 

doi: 10.1175/2008JTECHAI176.1, 2009.

Meehl, G.A., C. Covey, T. Delworth, M. Latif, B. McAvaney, J.F.B. Mitchell, R.J. Stouffer, and K.E. 

Taylor, The WCRP CMIP3 multi-model dataset: 

A new era in climate change research, Bull. Amer. 

Meteorol. Soc., 88 (9), 1383-1394, doi: 10.1175/

BAMS-88-9-1383, 2007.

Meehl, G.A., J.M. Arblaster, G. Branstator, and H. van 

Loon, A coupled air-sea response mechanism to solar forcing in the Pacific region, J. Clim., 21 (12), 

2883-2897, doi: 10.1175/2007JCLI1776.1, 2008.

Meehl, G.A., J.M. Arblaster, K. Matthes, F. Sassi, and H. 

van Loon, Amplifying the Pacific climate system 

response to a small 11-year solar cycle forcing, 

Science, 325 (5944), 1114-1118, doi: 10.1126/science.1172872, 2009.

Meneghini, B., I. Simmonds, and I.N. Smith, Association between Australian rainfall and the Southern 

Annular Mode, Int. J. Climatol., 27, 109-121, doi: 

10.1002/joc:1370, 2007.

Meredith, M.P., and A.M. Hogg, Circumpolar response of 

the Southern Ocean eddy activity to a change in the 

Southern Annular Mode, Geophys. Res. Lett., 33, 

L16608, doi: 10.1029/2006GL026499, 2006.

Miller, R.L., G.A. Schmidt, and D.T. Shindell, Forced 

annular variations in the 20th century Intergovernmental Panel on Climate Change Fourth Assessment Report models, J. Geophys. Res., 111, 

D18101, doi: 10.1029/2005JD006323, 2006.

Monaghan, A.J., D.H. Bromwich, W. Chapman, and J.C. 

Comiso, Recent variability and trends of Antarctic 

near-surface temperature, J. Geophys. Res., 113, 

D04105, doi: 10.1029/2007JD009094, 2008.

Morgenstern, O., M.A. Giorgetta, K. Shibata, V. Eyring, 

D.W. Waugh, T.G. Shepherd, H. Akiyoshi, J. Austin, A.J.G. Baumgaertner, S. Bekki, P. Braesicke, 

C. Brühl, M.P. Chipperfield, D. Cugnet, M. Dameris, S. Dhomse, S.M. Frith, H. Garny, A. Gettelman, S.C. Hardiman, M.I. Hegglin, P. Jöckel, D.E. 

Kinnison, J.-F. Lamarque, E. Mancini, E. Manzini, M. Marchand, M. Michou, T. Nakamura, J.E. 

Nielsen, D. Olivié, G. Pitari, D.A. Plummer, E. 

Rozanov, J.F. Scinocca, D. Smale, H. Teyssèdre, 

M. Toohey, W. Tian, and Y. Yamashita, Review 

of the formulation of present-generation stratospheric chemistry-climate models and associated 

external forcings, J. Geophys. Res., 115, D00M02, 

doi:10.1029/2009JD013728, 2010a.

Morgenstern, O., H. Akiyoshi, S. Bekki, P. Braesicke, N. 

Butchart, M.P. Chipperfield, D. Cugnet, M. Deushi, 

S.S. Dhomse, R.R. Garcia, A. Gettelman, N.P. Gillett, S.C. Hardiman, J. Jumelet, D.E. Kinnison, J.-F. 

Lamarque, F. Lott, M. Marchand, M. Michou, T. 

Nakamura, D. Olivié, T. Peter, D. Plummer, J.A. 
4.55

Stratospheric Changes and Climate

Pyle, E. Rozanov, D. Saint-Martin, J.F. Scinocca, 

K. Shibata, M. Sigmond, D. Smale, H. Teyssèdre, 

W. Tian, A. Voldoire, and Y. Yamashita, Anthropogenic forcing of the Northern Annular Mode in CCMVal-2 models, J. Geophys. Res., 115 (D00M03), 

doi: 10.1029/2009JD013347, 2010b.

Mote, P.W., K.H. Rosenlof, M.E. McIntyre, E.S. Carr, J.C. 

Gille, J.R. Holton, J.S. Kinnersley, H.C. Pumphrey, 

J.M. Russell, and J.W. Waters, An atmospheric 

tape recorder: The imprint of tropical tropopause 

temperatures on stratospheric water vapor, J. Geophys. Res., 101 (D2), 3989-4006, 1996.

Nielsen, J.K., N. Larsen, F. Cairo, G. Di Donfrancesco, 

J.M. Rosen, G. Durry, G. Held, and J.P. Pommereau, Solid particles in the tropical lowest stratosphere, Atmos. Chem. Phys., 7, 685-695, 2007.

Notholt, J., B.P. Luo, S. Füeglistaler, D. Weisenstein, 

M. Rex, M.G. Lawrence, H. Bingemer, I. Wohltmann, T. Corti, T. Warneke, R. von Kuhlmann, 

and T. Peter, Influence of tropospheric SO2 emissions on particle formation and the stratospheric 

humidity, Geophys. Res. Lett., 32, L07810, doi: 

10.1029/2004GL022159, 2005.

Oke, P.R., and M.H. England, Oceanic response to changes 

in the latitude of the Southern Hemisphere subpolar 

westerly winds, J. Clim., 17 (5), 1040-1054, 2004.

Olsen, S.C., C.A. McLinden, and M.J. Prather, Stratospheric N2O-NOy system: Testing uncertainties in 

a three-dimensional framework, J. Geophys. Res., 

106 (D22), 28771-28784, 2001.

Oltmans, S.J., H. Vömel, D.J. Hofmann, K.H. Rosenlof, 

and D. Kley, The increase in stratospheric water 

vapor from balloonborne frostpoint hygrometer 

measurements at Washington D.C., and Boulder, 

Colorado, Geophys. Res. Lett., 27 (21), 3453-

3456, 2000.

Oman, L., D.W. Waugh, S. Pawson, R.S. Stolarski, 

and P.A. Newman, On the influence of anthropogenic forcings on changes in the stratospheric mean age, J. Geophys. Res., 114, doi: 

10.1029/2008JD010378, 2009.

Ordóňez, C., D. Brunner, J. Staehelin, P. Hadjinicolaou, 

J.A. Pyle, M. Jonas, H. Wernli, and A.S.H. Prevot, 

Strong influence of lowermost stratospheric ozone 

on lower tropospheric background ozone changes 

over Europe, Geophys. Res., Lett., 34, L07805, doi: 

10.1029/2006GL029113, 2007.

Pawson, S., K. Labitzke, and S. Leder, Stepwise changes 

in stratospheric temperature, Geophys. Res. Lett.,

25 (12), 2157-2160, 1998.

Perlwitz, J., S. Pawson, R.L. Fogt, J.E. Nielsen, and W.D. 

Neff, Impact of stratospheric ozone hole recovery on Antarctic climate, Geophys. Res. Lett., 35, 

L08714, doi: 10.1029/2008GL033317, 2008.

Peter, T., C. Marcolli, P. Spichtinger, T. Corti, M.B. 

Baker, and T. Koop, When dry air is too humid, 

Science, 314 (5804), 1399-1402, doi: 10.1126/science.1135199, 2006.

Pezza, A.B., I. Simmonds, and J.A. Renwick, Southern 

hemisphere cyclones and anticyclones: Recent 

trends and links with decadal variability in the 

Pacific Ocean, Int. J. Climatol., 27 (11), 1403-

1419, 2007.

Pezza, A.B., T. Durrant, I. Simmonds, and I. Smith, 

Southern Hemisphere synoptic behavior in extreme 

phases of SAM, ENSO, sea ice extent, and southern 

Australia rainfall, J. Clim., 21 (21), 5566-5584, doi: 

10.1175/2008JCLI2128.1, 2008.

Pitari, G., E. Mancini, V. Rizi, and D.T. Shindell, Impact 

of future climate and emission changes on stratospheric aerosols and ozone, J. Atmos. Sci., 59 (3), 

414-440, 2002.

Plumb, R.A., and J. Eluszkiewicz, The Brewer–Dobson 

circulation: Dynamics of the tropical upwelling, J. 

Atmos. Sci., 56, 868-890, 1999.

Polvani, L.M., and P.J. Kushner, Tropospheric response 

to stratospheric perturbations in a relatively simple 

general circulation model, Geophys. Res. Lett., 29, 

doi: 10.1029/2001GL014284, 2002.

Portmann, R.W., and S. Solomon, Indirect radiative forcing of the ozone layer during the 21st century, 

Geophys. Res. Lett., 34, L02813, doi: 10.1029/ 

2006GL028252, 2007.

Prinn, R.G., J. Huang, R.F. Weiss, D.M. Cunnold, P.J. Fraser, P.G. Simmonds, A. McCulloch, C. Harth, S. Reimann, P. Salameh, S. O’Doherty, R.H.J. Wang, L.W. 

Porter, B.R. Miller, and P.B. Krummel, Evidence for 

variability of atmospheric hydroxyl radicals over the 

past quarter century, Geophys. Res. Lett., 32, L07809, 

doi: 10.1029/2004GL022228, 2005.

Punge, H.J., and M.A. Giorgetta, Net effect of the QBO in 

a chemistry climate model, Atmos. Chem. Phys., 8

(21), 6505-6525, 2008.

Radko, T., and J. Marshall, Equilibration of a warm 

pumped lens on a β plane, J. Phys. Oceanogr., 33

(4), 885-899, 2003.

Ramaswamy, V., M.L. Chanin, J. Angell, J. Barnett, D. 

Gaffen, M. Gelman, P. Keckhut, Y. Koshelkov, K. 

Labitzke, J.J.R. Lin, A. O’Neill, J. Nash, W. Randel, R. Rood, K. Shine, M. Shiotani, and R. Swinbank, Stratospheric temperature trends: Observations and model simulations, Rev. Geophys., 39 (1), 

71-122, 2001.

Ramaswamy, V., M.D. Schwarzkopf, W.J. Randel, B.D. 

Santer, B.J. Soden, and G.L. Stenchikov, Anthropogenic and natural influences in the evolution of 

lower stratospheric cooling, Science, 311 (5764), 

1138-1141, 2006.
4.56

Chapter 4

Randel, W.J., Variability and trends in stratospheric temperature and water vapor, in The Stratosphere: Dynamics, Transport and Chemistry, A Festschrift to 

Celebrate Alan Plumb’s 60th Birthday, American 

Geophysical Union, in press, 2010.

Randel, W.J., and F. Wu, Biases in stratospheric and 

tropospheric temperature trends derived from 

historical radiosonde data, J. Clim., 19 (10), 

2094-2104, 2006.

Randel, W.J., and F. Wu, A stratospheric ozone profile data 

set for 1979–2005: Variability, trends, and comparisons with column ozone data, J. Geophys. Res., 112, 

D06313, doi: 10.1029/2006JD007339, 2007.

Randel, W.J., F. Wu, S.J. Oltmans, K. Rosenlof, and 

G.E. Nedoluha, Interannual changes of stratospheric water vapor and correlations with tropical 

tropopause temperatures, J. Atmos. Sci., 61 (17), 

2133-2148, 2004.

Randel, W.J., F. Wu, H. Vömel, G.E. Nedoluha, and 

P. Forster, Decreases in stratospheric water vapor after 2001: Links to changes in the tropical tropopause and the Brewer-Dobson circulation, J. Geophys. Res., 111, D12312, doi: 

10.1029/2005JD006744, 2006.

Randel, W.J., K.P. Shine, J. Austin, J. Barnett, C. Claud, 

N.P. Gillett, P. Keckhut, U. Langematz, R. Lin, 

C. Long, C. Mears, A. Miller, J. Nash, D.J. 

Seidel, D.W.J. Thompson, F. Wu, and S. Yoden, 

An update of observed stratospheric temperature trends, J. Geophys. Res., 114, D02107, doi: 

10.1029/2008JD010421, 2009.

Rao, V.B., A.M.C. do Carmo, and S.H. Franchito, Interannual variations of storm tracks in the Southern 

Hemisphere and their connections with the Antarctic oscillation, Int. J. Climatol., 23 (12), 1537-1545, 

doi: 10.1002/joc.948, 2003.

Rayner, N.A., D.E. Parker, E.B. Horton, C.K. Folland, 

L.V. Alexander, D.P. Rowell, E.C. Kent, and A. 

Kaplan, Global analyses of sea surface temperature, sea ice, and night marine air temperature since 

the late nineteenth century, J. Geophys. Res., 108, 

4407, doi: 10.1029/2002JD002670, 2003.

Reed, R.J., and C.L. Vleck, The annual temperature variation in the lower tropical stratosphere, J. Atmos. 

Sci., 26 (1), 163-167, 1969.

Reichler, T., Changes in the atmospheric circulation as 

indicator of climate change, in Climate Change: 

Observed Impacts on Planet Earth, edited by T. M. 

Letcher, Elsevier BV, The Netherlands, ISBN: 978-

0-444-53301-2, pp. 145-164, 2009.

Renwick, J., and D. Thompson, The Southern Annular 

Mode and New Zealand climate, Water Atmos., 14 

(2), 24-25, 2006.

Rind, D., D. Shindell, P. Lonergan, and N.K. Balachandran, Climate change and the middle atmosphere. 

Part III: The doubled CO2 climate revisited, J. Clim.

11 (5), 876-894, 1998.

Rind, D., J. Lean, J. Lerner, P. Lonergan, and A. Leboissitier, Exploring the stratospheric/tropospheric 

response to solar forcing, J. Geophys. Res., 113, 

D24103, doi: 10.1029/2008JD010114, 2008.

Ring, M.J., and R.A. Plumb, The response of a simplified 

GCM to axisymmetric forcings: Applicability of 

the fluctuation–dissipation theorem, J. Atmos. Sci.,

65 (12), 3880-3898, 2008.

Robock, A., Volcanic eruptions and climate, Rev. Geophys., 38 (2), 191-219, 2000.

Rohs, S., C. Schiller, M. Riese, A. Engel, U. Schmidt, 

T. Wetter, I. Levin, T. Nakazawa, and S. Aoki, 

Long-term changes of methane and hydrogen 

in the stratosphere in the period 1978-2003 and 

their impact on the abundance of stratospheric 

water vapor, J. Geophys. Res., 111, D14315, doi: 

10.1029/2005JD006877, 2006.

Rosenlof, K.H., Transport changes inferred from HALOE 

water and methane measurements, J. Meteorol. Soc. 

Japan, 80 (48), 831-848, 2002.

Rosenlof, K.H., and J.R. Holton, Estimates of the stratospheric residual circulation using the downward 

control principle, J. Geophys. Res., 98 (D6), 10465-

10479, 1993.

Rosenlof, K.H., and G.C. Reid, Trends in the temperature 

and water vapor content of the tropical lower stratosphere: Sea surface connection, J. Geophys. Res.,

113, D06107, doi: 10.1029/2007JD009109, 2008.

Rosenlof, K.H., S.J. Oltmans, D. Kley, J.M. Russell III, 

E.-W Chiou, W.P. Chu, D.G. Johnson, K.K. Kelly, 

H.A. Michelsen, G.E. Nedoluha, E.E. Remsberg, 

G.C. Toon, and M.P. McCormick, Stratospheric 

water vapor increases over the past half-century, 

Geophys. Res. Lett., 28 (7), 1195-1198, 2001.

Roy, I., and J.D. Haigh, Solar cycle signals in sea level 

pressure and sea surface temperature, Atmos. Chem. 

& Phys., 10, 3147-3153, 2010.

Russell, J.L, K.W. Dixon, A. Gnanadesikan, R.J. Stouffer, 

and J.R. Toggweiler, The Southern Hemisphere 

westerlies in a warming world: Propping open the 

door to the deep ocean, J. Clim., 19 (24), 6382-

6390, doi: 10.1175/JCLI3984.1, 2006.

Santer, B.D., M.F. Wehner, T.M.L. Wigley, R. Sausen, 

G.A. Meehl, K.E. Taylor, C. Ammann, J. Arblaster, W.M. Washington, J.S. Boyle, and W. 

Brüggemann, Contributions of anthropogenic and 

natural forcing to recent tropopause height changes, Science, 301 (5632), 479-483, doi: 10.1126/ 

science.1084123, 2003.

Scaife, A.A., N. Butchart, C.D. Warner, D. Stainforth, 

W. Norton, and J. Austin, Realistic quasi-biennial 
4.57

Stratospheric Changes and Climate

oscillations in a simulation of the global climate, 

Geophys. Res. Lett., 27 (21), 3481-3484, 2000a.

Scaife, A.A., J. Austin, N. Butchart, S. Pawson, M. Keil, 

J. Nash, and I.N. James, Seasonal and interannual variability of the stratosphere diagnosed from 

UKMO TOVS analyses, Quart. J. Roy. Meteorol. 

Soc., 126 (568), 2585-2604, 2000b.

Scaife, A.A., N. Butchart, D.R. Jackson, and R. Swinbank, Can changes in ENSO activity help to 

explain increasing stratospheric water vapor?, 

Geophys. Res. Lett., 30, 1880, doi: 10.1029/ 

2003GL017591, 2003.

Scherer, M., H. Vömel, S. Fueglistaler, S.J. Oltmans, and 

J. Staehelin, Trends and variability of midlatitude 

stratospheric water vapour deduced from the reevaluated Boulder balloon series and HALOE, Atmos. Chem. Phys., 8, 1391-1402, 2008.

Schiller, C., J.-U. Grooß, P. Konopka, F. Plöger, F.H. Silva dos Santos, and N. Spelten, Hydration and dehydration at the tropical tropopause, Atmos. Chem. 

Phys., 9, 9647-9660, doi: 10.5194/acp-9-9647-

2009, 2009.

Schultz, M.G., T. Diehl, G.P. Brasseur, and W. Zittel, Air 

pollution and climate forcing impacts of a global 

hydrogen economy, Science, 302 (5645), 624-627, 

doi: 10.1126/science.1089527, 2003.

Schwarzkopf, M.D., and V. Ramaswamy, Evolution of stratospheric temperature in the 20th

century, Geophys. Res. Lett., 35, L03705, doi: 

10.1029/2007GL032489, 2008.

Scinocca, J.F., M.C. Reader, D.A. Plummer, M. Sigmond, 

P.J. Kushner, T.G. Shepherd, and A.R. Ravishankara, Impact of sudden Arctic sea-ice loss on stratospheric polar ozone recovery, Geophys. Res. Lett.,

36, L24701, doi: 10.1029/2009GL041239, 2009.

Screen, J.A., N.P. Gillett, D.P. Stevens, G.J. Marshall, 

and H.K. Roscoe, The role of eddies in the Southern Ocean temperature response to the Southern 

Annular Mode, J. Clim., 22 (3), 806-818, doi: 

10.1175/2008JCLI2416.1, 2009.

Seager, R., M. Ting, I. Held, Y. Kushnir, J. Lu, G. Vecchi, 

H.P. Huang, N. Harnik, A. Leetmaa, N.C. Lau, C. 

Li, J. Velez and N. Naik, Model projections of an 

imminent transition to a more arid climate in Southwestern North America, Science, 316 (5828), 1181-

1184, doi: 10.1126/science.1139601, 2007.

Seidel, D.J., and J.R. Lanzante, An assessment of three alternatives to linear trends for characterizing global 

atmospheric temperature changes, J. Geophys. Res.,

109, D14108, doi: 10.1029/2003JD004414, 2004.

Seidel, D.J., and W.J. Randel, Variability and trends 

in the global tropopause estimated from radiosonde data, J. Geophys. Res., 111, D21101, doi: 

10.1029/2006JD007363, 2006.

Seidel, D.J., and W.J. Randel, Recent widening of the 

tropical belt: Evidence from tropopause observations, J. Geophys. Res., 112, D20113, doi: 

10.1029/2007JD008861, 2007.

Seidel, D.J., Q. Fu, W.J. Randel, and T.J. Reichler, 

Widening of the tropical belt in a changing climate, Nature Geosci., 1, 21-24, doi: 10.1038/ 

nego.2007.38, 2008.

Seidel, D.J., F.H. Berger, H.J. Diamond, J. Dykema, D. 

Goodrich, F. Immler, W. Murray, T. Peterson, 

D. Sisterson, M. Sommer, P. Thorne, H. Vömel, 

and J. Wang, Reference upper-air observations for climate: Rationale, progress, and plans, 

Bull. Amer. Meteorol. Soc., 90 (3), 361-369, doi: 

10.1175/2008BAMS2540.1, 2009.

Semeniuk, K., and T.G. Shepherd, Mechanism for tropical upwelling in the stratosphere, J. Atmos. Sci., 58

(21), 3097-3115, 2001.

Shepherd, T.G., and A.I. Jonsson, On the attribution of 

stratospheric ozone and temperature changes to 

changes in ozone-depleting substances and well 

mixed greenhouse gases, Atmos. Chem. Phys., 8

(5), 1435-1444, 2008.

Sherwood, S., A microphysical connection among biomass burning, cumulus clouds, and stratospheric 

moisture, Science, 295 (5558), 1272-1275, doi: 

10.1126/science.1065080, 2002.

Sherwood, S.C., J.R. Lanzante, and C. Meyer, Radiosonde 

daytime biases and late-20th century warming, Science, 309 (5740), 1556-1559, doi: 10.1126/science.1115640, 2005.

Sherwood, S.C., C.L. Meyer, R.J. Allen, and H.A. Titchner, 

Robust tropospheric warming revealed by iteratively homogenized radiosonde data, J. Clim., 21 (20), 

5336-5352, doi: 10.1175/2008JCLI2320.1, 2008.

Shindell, D.T., Climate and ozone response to increased 

stratospheric water vapor, Geophys. Res. Lett., 28

(8), 1551-1554, 2001.

Shindell, D.T., and V. Grewe, Separating the influence of 

halogen and climate changes on ozone recovery in 

the upper stratosphere, J. Geophys. Res., 107, doi: 

10.1029/2001JD000420, 2002.

Shindell, D., and G.A. Schmidt, Southern Hemisphere 

climate response to ozone changes and greenhouse 

gas increases, Geophys. Res. Lett., 31, L18209, doi: 

10.1029/2004GL020724, 2004.

Shindell, D.T., D. Rind, N. Balachandran, J. Lean, and P. 

Lonergan, Solar cycle variability, ozone, and climate, Science, 284 (5412), 305-308, 1999.

Shindell, D.T., G.A. Schmidt, M.E. Mann, and G. Faluvegi, Dynamic winter climate response to large tropical volcanic eruptions since 1600, J. Geophys. Res.,

109, D05104, doi: 10.1029/2003JD004151, 2004.

Shindell, D.T., G. Faluvegi, R.L. Miller, G.A. Schmidt, 
4.58

Chapter 4

J.E. Hansen, and S. Sun, Solar and anthropogenic 

forcing of tropical hydrology, Geophys. Res. Lett,

33, L24706, doi: 10.1029/2006GL027468, 2006.

Shine, K.P., M.S. Bourqui, P.M.F. Forster, S.H.E. Hare, 

U. Langematz, P. Braesicke, V. Grewe, M. Ponater, 

C. Schnadt, C.A. Smith, J.D. Haigh, J. Austin, N. 

Butchart, D.T. Shindell, W.J. Randel, T. Nagashima, R.W. Portmann, S. Solomon, D.J. Seidel, J. 

Lanzante, S. Klein, V. Ramaswamy, and M.D. 

Schwarzkopf, A comparison of model-simulated 

trends in stratospheric temperatures, Quart. J. Roy. 

Meteorol. Soc., 129 (590), 1565-1588, 2003.

Shine, K.P., J.J. Barnett, and W.J. Randel, Temperature trends derived from Stratospheric Sounding 

Unit radiances: The effect of increasing CO2 on 

the weighting function, Geophys. Res. Lett., 35, 

L02710, doi: 10.1029/2007GL032218, 2008.

Sigmond, M., J.C. Fyfe, and J.F. Scinocca, Does the ocean 

impact the atmospheric response to stratospheric 

ozone depletion?, Geophys. Res. Lett., 37, L12706, 

doi: 10.1029/2010GL043773, 2010.

Simmonds, I., and K. Keay, Variability of Southern Hemisphere extratropical cyclone behavior, 1958–97, J. 

Clim., 13 (3), 550-561, 2000.

Simpson, I.R., M. Blackburn, and J.D. Haigh, The role 

of eddies in driving the tropospheric response to 

stratospheric heating perturbations, J. Atmos. Sci.,

66 (5), 1347-1365, doi: 10.1175/2008JAS2758.1, 

2009.

Solomon, S., R.W. Portmann, and D.W.J. Thompson, 

Contrasts between Antarctic and Arctic ozone depletion, Proc. Natl. Acad. Sci., 104 (2), 445-449, 

doi: 10.1073/pnas.0604895104, 2007.

Solomon, S., K.H. Rosenlof, R.W. Portmann, J.S. Daniel, 

S.M. Davis, T.J. Sanford, and G.-K. Plattner, Contributions of stratospheric water vapor to decadal changes in the rate of global warming, Science, 327 (5970), 

1219-1223, doi: 10.1126/science.1182488, 2010.

Son, S.-W., L.M. Polvani, D.W. Waugh, H. Akiyoshi, R. 

Garcia, D. Kinnison, S. Pawson, E. Rozanov, T.G. 

Shepherd, and K. Shibata, The impact of stratospheric ozone recovery on the Southern Hemisphere 

westerly jet, Science, 320 (5882), 1486-1489, doi: 

10.1126/science.1155939, 2008.

Son, S.-W., L.M. Polvani, D.W. Waugh, T. Birner, H. Akiyoshi, R.R. Garcia, A. Gettelman, D.A. Plummer, and 

E. Rozanov, The impact of stratospheric ozone recovery on tropopause height trends, J. Clim., 22 (2), 429-

445, doi: 10.1175/2008JCLI2215.1, 2009a.

Son, S.-W., N.F. Tandon, L.M. Polvani, and D.W. Waugh, 

Ozone hole and Southern Hemisphere climate 

change, Geophys. Res. Lett., 36, L15705, doi: 

10.1029/2009GL038671, 2009b.

Son, S.-W., E.P. Gerber, J. Perlwitz, L.M. Polvani, N.P. 

Gillett, K.-H. Seo, V. Eyring, T.G. Shepherd, D. 

Waugh, H. Akiyoshi, J. Austin, A. Baumgaertner, 

S. Bekki, P. Braesicke, C. Brühl, N. Butchart, M.P. 

Chipperfield, D. Cugnet, M. Dameris, S. Dhomse, 

S. Frith, H, Garny, R. Garcia, S.C. Hardiman, P. 

Jöckel, J.F. Lamarque, E. Mancini, M. Marchand, M. Michou, T. Nakamura, O. Morgenstern, 

G. Pitari, D.A. Plummer, J. Pyle, E. Rozanov, J.F. 

Scinocca, K. Shibata, D. Smale, H. Teyssèdre, W. 

Tian, and Y. Yamashita, Impact of stratospheric 

ozone on Southern Hemisphere circulation change: 

A multimodel assessment, J. Geophys. Res., 115, 

D00M07, doi: 10.1029/2010JD014271, 2010. 

Song, Y., and W.A. Robinson, Dynamical mechanisms 

for stratospheric influences on the troposphere, J. 

Atmos. Sci., 61 (14), 1711-1725, 2004.

Soukharev, B.E., and L.L. Hood, Solar cycle variation of 

stratospheric ozone: Multiple regression analysis 

of long-term satellite data sets and comparisons 

with models, J. Geophys. Res., 111, D20314, doi: 

10.1029/2006JD007107, 2006.

SPARC (Stratospheric Processes And their Role in Climate), SPARC Assessment of Stratospheric Aerosol 

Properties, edited by L. Thomason, and Th. Peter, World Climate Research Progam Report 124, 

SPARC Report 4, WMO/TD-No. 1295, 346 pp., 

Verrières le Buisson, France, 2006. 

SPARC CCMVal, SPARC Report on the Evaluation of 

Chemistry-Climate Models, edited by V. Eyring, 

T.G. Shepherd, and D.W. Waugh, Stratospheric 

Processes And their Role in Climate (SPARC) 

Report No. 5, WCRP-132, WMO/TD-No. 1526, 

available: http://www.atmosp.physics.utoronto.ca/

SPARC, 2010.

Sprintall, J., Long-term trends and interannual variability of 

temperature in Drake Passage, Prog. Ocean., 77 (4), 

316-330, doi: 10.1016/j.pocean.2006.06.004, 2008.

Steig, E.J., D.P. Schneider, S.D. Rutherford, M.E. Mann, 

J.C. Comiso, and D.T. Shindell, Warming of the 

Antarctic ice sheet surface since the 1957 International Geophysical Year, Nature, 457, 459-462, 

doi: 10.1038/nature07669, 2009.

Stenchikov, G., K. Hamilton, R.J. Stouffer, A. Robock, 

V. Ramaswamy, B. Santer, and H.-F. Graf, Arctic 

Oscillation response to volcanic eruptions in the 

IPCC AR4 climate models, J. Geophys. Res., 111, 

D07107, doi: 10.1029/2005JD006286, 2006.

Stenchikov, G., T.L. Delworth, V. Ramaswamy, R.J. 

Stouffer, A. Wittenberg, and F. Zeng, Volcanic signals in oceans, J. Geophys. Res., 114, D16104, doi: 

10.1029/2008JD011673, 2009.

Stevenson, D.S., F.J. Dentener, M.G. Schultz, K. Ellingsen, T.P.C. vanNoije, O. Wild, G. Zeng, M. Anann, 

C.S. Atherton, N. Bell, D.J. Bergmann, I. Bey, 
4.59

Stratospheric Changes and Climate

T. Butler, J. Cofala, W.J. Collins, R.G. Derwent, 

R.M. Doherty, J. Drevet, H.J. Eskes, A.F. Fiore, 

M. Gauss, D.A. Hauglustaine, L.W. Horowitz, 

I.S.A. Isaksen, M.C. Krol, J.-F. Lamarque, M.G. 

Lawrence, V. Montanaro, J.-F. Müller, G. Pitari, 

M.J. Prather, S.A. Pyle, S. Rast, J.M. Rodriguez, 

M.G. Sanderson, H.H. Savage, D.T. Shindell, 

S.E. Strahan, K. Sudo, and S. Szopa, Multi-model 

ensemble of present-day and near-future tropospheric ozone, J. Geophys. Res., 111, D8301, doi: 

10.1029/2005JD006338, 2006.

Stolarski, R.S., A.R. Douglass, P.A. Newman, S. Pawson, and M.R. Schoeberl, Relative contribution of 

greenhouse gases and ozone-depleting substances 

to temperature trends in the stratosphere: A chemistry-climate model study, J. Clim., 23 (1), 28-42, 

doi: 10.1175/2009JCLI2955.1, 2010.

Sudo, K., M. Takahashi, and H. Akimoto, Future changes in stratosphere-troposphere exchange and their 

impacts on future tropospheric ozone simulations, Geophys. Res. Lett., 30 (24), 2256, doi: 

10.1029/2003GL018526, 2003.

Takahashi, M., Simulation of the Quasi-Biennial Oscillation in a general circulation model, Geophys. Res. 

Lett., 26 (9), 1307-1310, 1999.

Terao, Y., J.A. Logan, A.R. Douglass, and R.S. Stolarski, 

Contribution of stratospheric ozone to the interannual variability of tropospheric ozone in the northern extratropics, J. Geophys. Res., 113, D18309, 

doi: 10.1029/2008JD009854, 2008.

Thejll, P., B. Christiansen, and H. Gleisner, On correlations 

between the North Atlantic Oscillation, geopotential 

heights, and geomagnetic activity, Geophys. Res. 

Lett., 30, 1347, doi: 10.1029/2002GL016598, 2003.

Thomason, L., L.R. Poole, and T. Deshler, A global climatology of stratospheric aerosol surface area density 

deduced from Stratospheric Aerosol and Gas Experiment II measurements: 1984-1994, J. Geophys. 

Res., 102 (D7), 8967-8976, 1997.

Thompson, D.W.J., and J.M. Wallace, Annular modes in 

the extratropical circulation. Part I: Month-to-month 

variability, J. Clim., 13 (5), 1000-1016, 2000.

Thompson, D.W.J., and S. Solomon, Interpretation of recent Southern Hemisphere climate change, Science,

296 (5569), 895-899, 2002.

Thompson, D.W.J., and S. Solomon, Recent stratospheric 

climate trends as evidenced in radiosonde data: 

Global structure and tropospheric linkages, J. Clim.,

18 (22), 4785-4795, 2005.

Thompson, D.W.J., and S. Solomon, Understanding recent 

stratospheric climate change, J. Clim., 22 (8), 1934-

1943, doi: 10.1175/2008JCLI2482.1, 2009.

Thompson, D.W.J., J.C. Furtado, and T.G. Shepherd, On 

the tropospheric response to anomalous stratospheric wave drag and radiative heating, J. Atmos. 

Sci., 63 (10), 2616-2629, 2006.

Thorne, P.W. (editor), Global climate, Chapter 2 in State 

of the Climate 2008, edited by T.C. Peterson and 

M.O. Baringer, Special Supplement to Bull. Amer. 

Meteorol. Soc., 90 (8), 196 pp., 2009.

Thorne, P.W., D.E. Parker, S.F.B. Tett, P.D. Jones, M. 

McCarthy, H. Coleman, and P. Brohan, Revisiting radiosonde upper air temperatures from 1958 

to 2002, J. Geophys. Res., 110, D18105, doi: 

10.1029/2004JD005753, 2005.

Tian, W.S., M.P. Chipperfield, and D. Lu , Impact of increasing stratospheric water vapor on ozone depletion and temperature change, Adv. Atmos. Sci., 26 (3), 

423-437, doi: 10.1007/s00376-009-0423-3, 2009.

Tie, X.X., and G. Brasseur, The response of stratospheric 

ozone to volcanic eruptions: Sensitivity to atmospheric chlorine loading, Geophys. Res. Lett., 22

(22), 3035-3038, 1995.

Tromp, T.K., R.-L. Shia, M. Allen, J.M. Eiler, and Y.L. 

Yung, Potential environmental impact of a hydrogen 

economy on the stratosphere, Science, 300 (5626), 

1740-1742, doi: 10.1126/science.1085169, 2003.

Turner, J., S.R. Colwell, G.J. Marshall, T.A. LachlanCope, A.M. Carleton, P.D. Jones, V. Lagun, P.A. 

Reid, and S. Iagovkina, Antarctic climate change 

during the last 50 years, Int. J. Climatol., 25 (3), 

279-294, doi: 10.1002/joc.1130, 2005.

Turner, J., J.C. Comiso, G.J. Marshall, T.A. LachlanCope, T. Bracegirdle, T. Maksym, M.P. Meredith, 

Z. Wang, and A. Orr, Non-annular atmospheric circulation change induced by stratospheric ozone depletion and its role in the recent increase of Antarctic sea ice extent, Geophys. Res. Lett., 36, L08502, 

doi: 10.1029/2009GL037524, 2009.

Udagawa, Y., Y. Tachibana, and K. Yamazaki, Modulation in interannual sea ice patterns in the Southern 

Ocean in association with large-scale atmospheric 

mode shift, J. Geophys. Res., 114, D21103, doi: 

10.1029/2009JD011807, 2009.

Ueyama, R., and J.M. Wallace, To what extent does highlatitude wave forcing drive tropical upwelling in the 

Brewer-Dobson Circulation?, J. Atmos. Sci., 67 (4), 

1232-1246, doi: 10.1175/2009JAS3216.1, 2010.

Ummenhofer, C.C., A. Sen Gupta, and M.H. England, 

Causes of late twentieth-century trends in New 

Zealand precipitation, J. Clim., 22 (1), 3-19, doi: 

10.1175/2008JCLI2323.1, 2009.

van Loon, H., and G.A. Meehl, The response in the Pacific 

to the sun’s decadal peaks and contrasts to cold 

events in the Southern Oscillation, J. Atmos. Solar Terr. Phys., 70 (7), 1046-1055, doi: 10.1016/j.

jastp.2008.01.009, 2008.

van Loon, H., G.A. Meehl, and D. Shea, Coupled air-sea 
4.60

Chapter 4

response to solar forcing in the Pacific region during northern winter, J. Geophys. Res., 112, D02108, 

doi: 10.1029/2006JD007378, 2007.

Velders, G.J.M., S.O. Andersen, J.S. Daniel, D.W. Fahey, 

and M. McFarland, The importance of the Montreal Protocol in protecting climate, Proc. Natl. 

Acad. Sci., 104 (12), 4814-4819, doi: 10.1073/

pnas.0610328104, 2007.

Vera, C., Interannual and interdecadal variability of atmospheric synoptic-scale activity in the Southern 

Hemisphere, J. Geophys. Res., 108, 8077, doi: 

10.1029/2000JC000406, 2003.

Vernier, J.P., J.P. Pommereau, A. Garnier, J. Pelon, N. 

Larsen, J. Nielsen, T. Christensen, F. Cairo, L.W. 

Thomason, T. Leblanc, and I.S. McDermid, Tropical stratospheric aerosol layer from CALIPSO lidar 

observations, J. Geophys. Res., 114, D00H10, doi: 

10.1029/2009JD011946, 2009.

Vömel, H., D.E. David, and K. Smith, Accuracy of tropospheric and stratospheric water vapor measurements 

by the cryogenic frost point hygrometer: Instrumental details and observations, J. Geophys. Res., 112, 

D08305, doi: 10.1029/2006JD007224, 2007.

Waugh, D., W. Randel, S. Pawson, P. Newman, and E. Nash, 

Persistence of the lower stratospheric polar vortices, J. 

Geophys. Res., 104 (D22), 27191-27201, 1999.

Waugh, D.W., L. Oman, P.A. Newman, R.S. Stolarski, S. 

Pawson, J.E. Nielsen, and J. Perlwitz, Effect of zonal asymmetries in stratospheric ozone on simulated 

Southern Hemisphere climate trends, Geophys. Res. 

Lett., 36, L18701, doi: 10.1029/2009GL040419, 2009.

Weinstock, E.M., J.B. Smith, D.S. Sayres, J.V. Pittman, J.R. 

Spackman, E.J. Hintsa, T.F. Hanisco, E.J. Moyer, J.M. 

St. Clair, M.R. Sargent, and J.G. Anderson, Validation of the Harvard Lyman-α in situ water vapor instrument: Implications for the mechanisms that control stratospheric water vapor, J. Geophys. Res., 114, 

D23301, doi: 10.1029/2009JD012427, 2009.

Weisenstein, D., G. Yue, M. Ko, N.-D. Sze, J. Rodriguez, 

and C. Scott, A two-dimensional model of sulfur 

species and aerosols, J. Geophys. Res., 102 (D11), 

13019-13035, 1997.

Wetzel, P., A. Winguth, and E. Maier-Reimer, Sea-toair CO2 flux from 1948 to 2003: A model study, 

Global Biogeochem. Cycles, 19, GB2005, doi: 

10.1029/2004GB002339, 2005.

White W.B., and Z. Liu, Non-linear alignment of El Niño 

to the 11-yr solar cycle, Geophys. Res. Lett., 35, 

L19607, doi: 10.1029/2008GL034831, 2008.

Wild, O., Modelling the global tropospheric ozone budget: 

Exploring the variability in current models, Atmos. 

Chem. Phys., 7 (10), 2643-2660, doi: 10.5194/acp7-2643-2007, 2007.

Woollings, T., Vertical structure of anthropogenic zonal-mean 

atmospheric circulation change, Geophys. Res. Lett.,

35, L19702, doi: 10.1029/2008GL034883, 2008.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 2002, Global 

Ozone Research and Monitoring Project–Report 

No. 47, 498 pp., Geneva, Switzerland, 2003.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 2006, Global 

Ozone Research and Monitoring Project–Report 

No. 50, 572 pp., Geneva, Switzerland, 2007.

Wu, S., L.J. Mickley, D.J. Jacob, J.A. Logan, R.M. Yantosca, and D. Rind, Why are there large differences 

between models in global budgets of tropospheric ozone?, J. Geophys. Res., 112, D05302, doi: 

10.1029/2006JD007801, 2007.

Yin, J.H., A consistent poleward shift of the storm tracks in 

simulations of 21st century climate, Geophys. Res. Lett.,

32, L18701, doi: 10.1029/2005GL023684, 2005.

Zazulie, N., M. Rusticucci, and S. Solomon, Changes in 

climate at high southern latitudes: A unique daily 

record at Orcadas spanning 1903-2008, J. Clim., 

23, 189-196, doi: 10.1175/2009JCLI3074.1, 2010.

Zeng, G., and J.A. Pyle, Changes in tropospheric ozone 

between 2000 and 2100 modeled in a chemistryclimate model, Geophys. Res. Lett., 30 (7), 1392, 

doi: 10.1029/2002GL016708, 2003.

Zeng, G., J.A. Pyle, and P.J. Young, Impact of climate 

change on tropospheric ozone and its global budgets, Atmos. Chem. Phys., 8, 369-387, 2008.

Zhang, J., Increasing Antarctic sea ice under warming 

atmospheric and oceanic conditions, J. Clim., 20

(11), doi: 10.1175/JCLI4136.1, 2007.

Zhou, J., and K.-K.Tung, Solar cycles in 150 years of global 

sea-surface temperature data, J. Clim., 23 (12), 3234-

3248, doi: 10.1175/2010JCLI3232.1, 2010.

Zhou, S., M.E. Gelman, A.J. Miller, and J.P. McCormack, 

An inter-hemisphere comparison of the persistent 

stratospheric polar vortex, Geophys. Res. Lett., 27

(8), 1123-1126, 2000.

Zhou, T., M.A. Geller, and K. Hamilton, The roles of the 

Hadley Circulation and downward control in tropical 

upwelling. J. Atmos. Sci., 63 (11), 2740-2757, 2006.

Zhou, X.-L., M.A. Geller, and M. Zhang, Cooling trend of 

the tropical cold point tropopause temperatures and 

its implications, J. Geophys. Res., 106 (D2), 1511-

1522, 2001.

Zickfeld, K., J.C. Fyfe, M. Eby, and A.J. Weaver, Comment on “Saturation of the southern ocean CO2 sink 

due to recent climate change,” Science, 319 (5863), 

570, doi: 10.1126/science.1146886, 2008.

Zou, C.-Z., M. Gao, and M.D. Goldberg, Error structure 

and atmospheric temperature trends in observations 

from the Microwave Sounding Unit, J. Clim., 22 (7),

1661-1681, doi: 10.1175/2008JCLI2233.1, 2009.
Chapter 5

A Focus on Information and 

Options for Policymakers

Coordinating Lead Authors:

J.S. Daniel

G.J.M. Velders

Lead Authors:

O. Morgenstern

D.W. Toohey

T.J. Wallington

D.J. Wuebbles

Coauthors:

H. Akiyoshi

A.F. Bais

E.L. Fleming

C.H. Jackman

L.J.M. Kuijpers

M. McFarland

S.A. Montzka

M.N. Ross

S. Tilmes

M.B. Tully

Contributors:

S.O. Andersen

U. Langematz

P.M. Midgley

SCIENTIFIC SUMMARY .............................................................................................................................................1

5.1 SUMMARY OF PREVIOUS ASSESSMENT AND KEY ISSUES 

TO BE ADDRESSED IN THE CURRENT ASSESSMENT.................................................................................5

5.2 METRICS USED TO QUANTIFY OZONE AND CLIMATE IMPACTS ...........................................................5

5.2.1 Background ...............................................................................................................................................5

5.2.2 Ozone Impacts: ODPs and EESC..............................................................................................................7

5.2.2.1 Fractional Release Values and Global Lifetimes........................................................................7

5.2.2.2 Ozone Depletion Potentials.........................................................................................................9

5.2.2.3 Equivalent Effective Stratospheric Chlorine ..............................................................................9

5.2.3 Climate Impacts: GWPs and Radiative Forcing......................................................................................10

5.2.3.1 Lifetime Updates.......................................................................................................................11

5.2.3.2 Radiative Efficiency Updates ...................................................................................................11

5.2.3.3 Updates to Indirect GWPs from Ozone Destruction.................................................................12

5.3 FUTURE BASELINE SCENARIOS ....................................................................................................................14

5.3.1 Chlorine- and Bromine-Containing Ozone-Depleting Substances .........................................................14

5.3.2 CO2, CH4, and N2O .................................................................................................................................16

5.3.3 ODP- and GWP-Weighted Emissions, EESC, and Radiative Forcing ...................................................16

5.4 IMPACTS OF HUMAN ACTIVITIES RELEVANT TO OZONE POLICY ......................................................19

5.4.1 Background .............................................................................................................................................19

5.4.2 Ozone Impacts.........................................................................................................................................20

5.4.2.1 Chlorine- and Bromine-Containing ODSs................................................................................20

5.4.2.2 CO2, CH4, and N2O...................................................................................................................22

5.4.2.3 N2O from Automotive Biofuels................................................................................................22

5.4.2.4 Geoengineering: Enhancing Earth’s Albedo by Stratospheric Injection of Sulfur...................23

5.4.2.5 Emissions from Aviation and Rockets......................................................................................24

5.4.2.6 Summary...................................................................................................................................25

5.4.3 Climate Impacts.......................................................................................................................................25

5.4.3.1 Major HFCs Used as Replacements for ODSs.........................................................................25

5.4.3.2 Other Replacements for ODSs..................................................................................................30

5.4.4 Other Environmental Impacts .................................................................................................................31

5.4.5 Impact of 2007 Accelerated HCFC Phase-Out .......................................................................................32

5.5 THE WORLD AVOIDED BY OZONE POLICY ................................................................................................33

5.5.1 ODS Production ......................................................................................................................................33

5.5.2 Radiative Forcing ....................................................................................................................................34

5.5.3 Climate and Ozone Impacts Avoided by the Montreal Protocol.............................................................35

5.5.4 UV Impacts of the Avoided Ozone Depletion ........................................................................................37

5.5.5 Summary .................................................................................................................................................37

REFERENCES .............................................................................................................................................................38

APPENDIX 5A

Table 5A-1. Direct Global Warming Potentials for selected gases....................................................................47

Chapter 5

A FOCUS ON INFORMATION AND OPTIONS FOR POLICYMAKERS

Contents
Table 5A-2. Assumptions made in obtaining production and emission estimates for

the baseline (A1) scenario..............................................................................................................50

Table 5A-3. Mixing ratios (ppt) of the ODSs considered in the baseline (A1) scenario ...................................54

Table 5A-4. Halocarbon indirect GWPs from ozone depletion using the EESC-based 

method described in Daniel et al. (1995).......................................................................................56
5.1

A Focus on Information for Policymakers

SCIENTIFIC SUMMARY

Ozone Depletion Potentials (ODPs) and Global Warming Potentials (GWPs) are metrics frequently used to quantify the relative impacts of substances on ozone depletion and climate forcing. In Chapter 5, both ODPs and GWPs have 

been updated. The direct GWPs for some compounds presented here have not appeared previously in WMO/UNEP or 

Intergovernmental Panel on Climate Change (IPCC) assessments. Indirect GWPs have also been re-evaluated.

Information for Policymakers

•	 The Montreal Protocol is working. It has protected the stratospheric ozone layer from much higher levels of

depletion by phasing out production and consumption of ozone-depleting substances (ODSs). Simulations show 

that unchecked growth in the emissions of ODSs would have led to ozone depletion globally in the coming decades 

much larger than has been observed. Solar ultraviolet-B (UV-B) radiation at the surface would also have increased 

substantially.

•	 The Montreal Protocol and its Amendments and Adjustments have made large contributions toward reducing global greenhouse gas emissions. Because many ODSs are potent greenhouse gases, the Montreal Protocol 

has successfully avoided larger climate forcing. In 2010, the decrease of annual ODS emissions under the Montreal 

Protocol is estimated to be about 10 gigatonnes (Gt) of carbon dioxide–equivalent (GtCO2-eq) per year, which is 

about five times larger than the annual emissions reduction target for the first commitment period (2008–2012) of 

the Kyoto Protocol.

•	 The accelerated hydrochlorofluorocarbon (HCFC) phase-out agreed to by the Parties to the Montreal Protocol

in 2007 is projected to reduce cumulative HCFC emissions by 0.6–0.8 million ODP-tonnes between 2011 and

2050 and bring forward the year equivalent effective stratospheric chlorine (EESC) returns to 1980 levels by

4–5 years. In terms relevant to climate, the accelerated HCFC phase-out is projected to reduce emissions by

0.4–0.6 GtCO2-eq per year averaged over 2011 through 2050. The actual climate benefit will be determined, 

in part, by the climate impact of the compounds used to replace the HCFCs. In comparison, global anthropogenic 

emissions of CO2 were greater than 30 Gt per year in 2008.

•	 EESC at midlatitudes is projected to return to 1980 levels in 2046 for the baseline (A1) scenario, 2–3 years

earlier than projected in the previous Assessment. This revision is primarily due to an improved understanding of 

lower stratospheric chlorine and bromine release from ODSs, along with contributions from smaller projected HCFC 

emissions, and despite larger projected emissions of carbon tetrachloride (CCl4) and a smaller 1980 mixing ratio of 

methyl bromide (CH3Br).

•	 EESC in the Antarctic vortex is projected to return to 1980 levels around 2073 for the baseline (A1) scenario,

7–8 years later than projected in the previous Assessment. This is primarily due to an improved understanding of lower stratospheric chlorine and bromine release from ODSs, with smaller contributions from changes in the 

emissions of CCl4 and HCFCs and a smaller 1980 mixing ratio of CH3Br. The return to 1980 levels in the Antarctic 

vortex is about 26 years later than the return of midlatitude EESC to 1980 levels.

•	 Due to the ongoing success of the Montreal Protocol and its Amendments and Adjustments in reducing the

production, emissions, and abundances of controlled ODSs, other compounds and activities not controlled by

the Montreal Protocol are becoming relatively more important to stratospheric ozone levels.

•	 Increasing abundances of radiatively important gases that are not controlled by the Montreal Protocol, especially CO2, methane (CH4), and nitrous oxide (N2O), are expected to significantly affect future stratospheric

ozone levels (see also Chapter 3). Under many IPCC future scenarios, it is projected that these gases will cause globally averaged ozone changes larger than those resulting from any of the ODS reduction cases explored in this chapter.
5.2

Chapter 5

•	 A nitrous oxide (N2O) ODP of 0.017 has been calculated. The anthropogenic ODP-weighted emission of N2O

is larger than that of any current halogenated ODS emission. The ODP of N2O is more uncertain than it is for 

halogenated substances, but it has been known since 1970 that N2O depletes stratospheric ozone. Reductions in N2O 

emissions would also reduce climate forcing.

•	 Since the previous Assessment, new fluorocarbons have been suggested as possible replacements for potent

HCFC and hydrofluorocarbon (HFC) greenhouse gases. For example, HFC-1234yf (CF3CF=CH2) (ODP = 0; 

100-year GWP = 4) is proposed to replace HFC-134a (CH2FCF3) (ODP = 0; 100-year GWP = 1370) in motor vehicle 

(mobile) air conditioning. Each new fluorocarbon proposed as a replacement will require an evaluation for ODP, 

GWP, atmospheric fate, safety, and toxicity for a thorough understanding of its potential environmental impact. 

Preliminary analyses of the atmospheric fate of HFC-1234yf indicate that global replacement of HFC-134a with 

HFC-1234yf at today’s level of use is not expected to contribute significantly to tropospheric ozone formation or 

harmful levels of the degradation product TFA (trifluoroacetic acid). It is well established that TFA is a ubiquitous 

natural component of the hydrosphere, but uncertainties remain regarding its natural and anthropogenic sources, longterm fate, and abundances.

Options for Policymakers

A new baseline scenario for ODSs is presented in Chapter 5 that reflects our current understanding of atmospheric mixing ratios, production levels, and bank sizes. Elimination of future emissions, production, and banks of various ODSs are applied to this scenario to evaluate the maximum impacts of various hypothetical phase-outs (see Table 

S5-1). The year EESC returns to 1980 levels, and integrated EESC changes, are two metrics used in the evaluation. The 

calculations of the years when EESC returns to the 1980 level in these hypothetical cases do not consider other effects 

such as changing atmospheric transport and lifetimes. An elimination of anthropogenic N2O emissions is also considered 

and compared to some ODS cases using globally averaged total ozone. In addition to the hypothetical cases discussed 

below, the impacts on stratospheric ozone of other activities, such as the use of automotive biofuels, commercial subsonic 

aircraft, and rocket launches, are considered in Chapter 5. These other activities are not expected to substantially affect 

stratospheric ozone now or in the near future.

•	 Projections suggest that unmitigated HFC growth could result in GWP-weighted emissions up to 8.8 GtCO2-eq

per year by 2050, comparable to the GWP-weighted emissions of chlorofluorocarbons (CFCs) at their peak in

1988. The highest of these projections assumes that developing countries use HFCs with GWPs comparable to those 

currently used in the same applications in developed countries. The projected radiative forcing in 2050 from these 

compounds (up to 0.4 W/m2) can be reduced by using compounds with lower GWPs.

•	 Options available for limiting future halocarbon emissions will have less impact on future ozone levels than

what has already been accomplished by the Montreal Protocol.

•	 Leakage of CFCs and leakage of halons from the banks are the largest sources of current ODP-weighted

emissions of ODSs. A delay of four years, from 2011 to 2015, in the capture and destruction of the estimated CFC 

banks is currently thought to reduce the potential ozone and climate benefits from these actions by about 30%. The 

percentage impact of a four-year delay in the capture and destruction of the halon banks is similar.

•	 Elimination of future CCl4 emissions is now projected to have a larger impact on integrated EESC than was

projected in the previous Assessment. Recent observed CCl4 mixing ratios have declined more slowly than previously projected. Extrapolation of this trend leads to larger future projected emissions in the baseline scenario and thus 

to the increased projected impact of the elimination of emissions. 

•	 The estimated impact on integrated EESC resulting from elimination of future HCFC production is slightly

smaller than in the previous Assessment. The recent growth in reported HCFC production in developing countries 

was larger than projected in the previous Assessment. This alone would have resulted in a larger projected HCFC
5.3

A Focus on Information for Policymakers

Table S5-1. Summary of hypothetical cases for accelerating the recovery of the ozone layer and reducing carbon-equivalent emissions. The table below shows the reductions in integrated EESC and integrated 

CO2-eq emissions relative to the baseline (A1) scenario that can be achieved in several hypothetical cases. 

The EESC excess above 1980 levels is integrated from 2011 until the time EESC returns to the 1980 level 

(before 2050). Any potential contribution from very short-lived substances is neglected.

Substance or Group of

Substances

Reductions (%) in

Integrated EESC

(equivalent effective

stratospheric chlorine)

Reduction in Cumulative

GWP-Weighted Emissions

from 2011 to 2050

(Gt of CO2-equivalent)

Bank capture and 

destruction in 2011 and 2015:

2011 2015 2011 2015

CFCs 11 7.0 7.9 5.5

Halons 14 9.1 0.4 0.3

HCFCs 4.8 5.3 1 4.9 5.5 1

Production elimination after 2010:

HCFCs 8.8 13.2

CH3Br for quarantine and pre-shipment 6.7 0.002

Total emissions elimination after 2010:

CCl4

 2 7.6 0.9

CH3CCl3 0.1 0.004

HFCs 0.0 Up to 170 3

1 The impact of a 2015 HCFC bank recovery is larger than a 2011 bank recovery because this calculation assumes destruction of the bank in only a 

single year, and because the bank in 2015 is larger than the bank in 2011 owing to continued annual production that is larger than the annual bank 

release. 2 Banks are assumed to be zero. Emissions include uncertain sources such as possible fugitive emissions and unintended by-product emissions. 3 Strongly dependent on future projections and does not consider HFC-23 emissions. HFCs are not controlled by the Montreal Protocol, but are included 

in the basket of gases of the Kyoto Protocol.

production in the new baseline scenario compared to the previous Assessment, but is projected to be more than compensated for by the accelerated HCFC phase-out agreed to by the Parties to the Montreal Protocol in 2007. Projections 

suggest that total emissions of HCFCs will begin to decline in the coming decade due to measures already agreed to 

under the Montreal Protocol. 

•	 The elimination of all emissions of chlorine- and bromine-containing ODSs after 2010 would shift the year

EESC reaches the 1980 level by about 13 years, from 2046 to 2033. In terms relevant to climate, this would reduce 

emissions of these substances by about 0.7 GtCO2-eq per year averaged over 2011 through 2050. Future production 

of HCFCs and the sum of the current banks of CFCs plus HCFCs contribute about equally to this number. In comparison, global anthropogenic emissions of CO2 were greater than 30 Gt per year in 2008.

•	 A phase-out of methyl bromide emissions from quarantine and pre-shipment (QPS) applications beginning in

2011 would shift the year EESC reaches the 1980 level earlier by 1.5 years compared to continued use at current levels. Continuing critical-use exemptions (CUEs) indefinitely at the approved 2011 level would delay the return 

of EESC to 1980 levels by 0.2 years.

•	 Elimination of anthropogenic emissions of very short-lived substances (VSLS) could shift the year EESC reaches the 1980 level earlier by almost 3 years, if anthropogenic VSLS contribute 40 parts per trillion of EESC to the 

stratosphere. It remains unclear, however, how VSLS emissions reductions at different surface locations would affect 

their contribution to stratospheric chlorine. VSLS are not controlled by the Montreal Protocol.

5.5

A Focus on Information for Policymakers

5.1 SUMMARY OF PREVIOUS ASSESSMENT 

AND KEY ISSUES TO BE ADDRESSED IN 

THE CURRENT ASSESSMENT

The benefits of the Montreal Protocol and its 

Amendments and Adjustments to both stratospheric ozone 

and climate have been well documented. Controls on the 

production and consumption of ozone-depleting substances have been so successful, in fact, that remaining options 

for further reducing future emissions of ozone-depleting 

substances (ODSs) are not expected to be as effective in 

reducing future ozone depletion and climate forcing as 

what has already been accomplished. As the ability to 

make further ODS reductions becomes more limited and 

ODS emissions continue to decline, other processes and 

anthropogenic activities are expected to become relatively 

more important in affecting future ozone evolution.

The majority of this chapter is devoted to assessing the impacts of various ozone-relevant processes and 

activities on ozone depletion and climate forcing. It serves 

primarily as an update to Chapter 8 of the previous Assessment (Daniel and Velders et al., 2007). In that chapter, 

future equivalent effective stratospheric chlorine (EESC) 

projections were updated with the goal of providing a more 

accurate estimate of the time when the impact of ODSs on 

stratospheric ozone depletion would return to its 1980 level. The most significant updates resulted from improved 

bank estimates of chlorofluorocarbon-11 (CFC-11) and 

CFC-12. For the first time, bank estimates were available 

from a bottom-up method, leading to more reliable future 

projections. Before the previous Assessment, banks were 

estimated from the cumulative difference between production and emissions, a technique thought to be characterized by large uncertainties. The value for the relative effectiveness of bromine compared with chlorine was also 

increased to be consistent with the latest literature, making 

methyl bromide (CH3Br) emissions and halon banks more 

important relative to the chorine-containing ODSs than 

they would have been otherwise. Using the updated EESC 

projections, the impacts of eliminating future production, 

bank release, and emissions of several compound groups 

were quantified. It was found that the elimination of future hydrochlorofluorocarbon (HCFC) emissions could reduce future EESC more than the elimination of emissions 

of any other compound group. In 2007, the Parties to the 

Montreal Protocol decided to accelerate the production 

and consumption phase-out of the HCFCs. The previous 

Assessment was also the first assessment that presented 

dates for the return of EESC to 1980 levels relevant to 

both the midlatitude and Antarctic stratosphere.

In the current chapter, Ozone Depletion Potentials 

(ODPs) and Global Warming Potentials (GWPs) for ODSs 

and their replacements are updated. New scenarios are 

generated to explore the potential impacts of some current 

hypothetical ODS and nitrous oxide (N2O) emissions reductions on future ozone depletion and climate forcing. 

These new scenarios incorporate updated bottom-up bank 

estimates and the latest ODS mixing ratio observations. 

The impact of the 2007 HCFC accelerated phase-out is 

also quantified. This chapter assesses some additional 

processes and activities that are expected to affect future 

ozone levels through mechanisms that do not necessarily 

involve the emission of chlorine- and bromine-containing 

source gases. Some of these processes could affect future ozone levels more than future emissions of controlled 

ODSs. We also discuss the impact of the Montreal Protocol on climate forcing where appropriate. For example, 

in addition to the climate forcing of ODSs, we assess the 

impact of future hydrofluorocarbon (HFC) abundances 

on climate because these chemicals are commonly used 

as replacement compounds for ODSs. Finally, since the 

previous Assessment, additional work has been published 

that investigates the impact of the Montreal Protocol on 

both ozone and climate. These studies support our understanding that the Montreal Protocol and its Amendments 

and Adjustments have averted many profound changes to 

Earth and its atmosphere.

Uncertainties remain in our ability to evaluate the 

effects of human activities on future ozone levels. Where 

appropriate, we identify gaps in our understanding that 

inhibit a precise quantification of ozone impacts.

5.2 METRICS USED TO QUANTIFY OZONE 

AND CLIMATE IMPACTS

5.2.1 Background

Halocarbons and other long-lived gases released 

from Earth’s surface become mixed in the lower atmosphere and are transported into the stratosphere by atmospheric dynamical processes. They are removed from 

the atmosphere by photolysis, reaction with excited-state 

oxygen atoms (O(1D)) and hydroxyl radicals (OH) (the 

latter typically only for unsaturated compounds or those 

containing C-H bonds), and for some compounds, uptake 

by the oceans and/or land (see Chapter 1). Halocarbons 

that are transported intact to the stratosphere can react or 

undergo photolysis and release their degradation products 

there directly. Some fraction of the degradation products 

from halocarbons that react before leaving the troposphere 

can also be transported to the stratosphere. The final degradation products are inorganic halogen species containing 

fluorine, chlorine, bromine, and iodine atoms. The fraction of the inorganic halogen present in the stratosphere as 

X atoms (X = Cl, Br, or I) and XO largely determines the 

efficiency of ozone destruction there. Fluorine atoms, for 
5.6

Chapter 5

example, exist as F and FO in very small relative quantities because fluorine species are rapidly converted into hydrogen fluoride (HF), a stable reservoir species that does 

not react with ozone. This prevents fluorine from contributing to ozone destruction to any significant degree (Ravishankara et al., 1994; Wallington et al., 1995). Iodine atoms participate in catalytic ozone destruction cycles, but 

rapid tropospheric loss of iodine-containing compounds 

reduces the amount of iodine reaching the stratosphere 

from surface emissions (see Chapter 1). Thus, it is primarily chlorine- and bromine-containing compounds that lead 

to ozone depletion. Halocarbons also absorb terrestrial radiation (long-wavelength infrared radiation emitted from 

Earth’s surface and by the atmosphere) and contribute to 

the radiative forcing of climate.

Simple metrics have been widely used to quantify 

the contribution of individual compounds to stratospheric 

ozone depletion and climate change. ODPs and GWPs are 

the most established metrics and have been used in past 

climate and ozone assessments (IPCC, 1990, 1995, 1996, 

2001, 2007; IPCC/TEAP, 2005; WMO, 1989, 1991, 1995, 

1999, 2003, 2007). They have qualities that are particularly useful in policy discussions. Specifically, they are 

simple and transparent concepts that are straightforward 

to estimate and communicate. They approximate the integrated impact of the emission of a given gas relative to that 

for the emission of the same mass of a reference compound 

(generally CFC-11 for ODPs and carbon dioxide (CO2) 

for GWPs). Some uncertainties in translating emissions 

into absolute environmental impacts tend to cancel, and 

the relative benefits of controlling emissions of different 

gases are highlighted when using such indices. ODPs and 

GWPs have found widespread use in international agreements such as the Montreal Protocol and Kyoto Protocol 

and in national regulatory discussions.

Both steady-state and time-dependent ODPs can 

provide valuable information about the potential for ozone 

destruction by a compound. Steady-state ODPs are de- Steady-state ODPs are defined as the change in global ozone for a sustained unit 

mass emission of a specific compound relative to the 

change in global ozone for the sustained unit mass emission of CFC-11 (CFCl3) (Fisher et al., 1990; Solomon et 

al., 1992; Wuebbles, 1983). 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; 2002). CFC-11 was a widely used industrial compound in the 1970s and 1980s and 

so has been chosen as a convenient reference gas (Fisher 

et al., 1990; Wuebbles, 1981; Wuebbles, 1983). Steadystate ODPs have no time dependence and are frequently 

calculated using chemical transport models. The accuracy 

of the calculation depends on the model’s ability to simulate the distribution of the considered compound and the 

associated ozone loss. However, because the ODPs are 

defined relative to the ozone loss caused by CFC-11, it 

is generally expected that for chlorocarbon compounds 

ODPs demonstrate less sensitivity to photochemical 

modeling errors than do absolute ozone loss calculations. 

Steady-state ODPs are normally derived relative to a fixed 

atmosphere; there would be differences in some ODPs if 

calculations were made for a future atmosphere with different background composition. ODPs with some time 

horizon, referred to as time-dependent ODPs (Solomon 

and Albritton, 1992), are more analogous to GWPs and 

provide information regarding the different timescales 

over which the compound and reference gas (CFC-11) 

liberate chlorine and bromine into the stratosphere. Compounds that have shorter (longer) atmospheric lifetimes 

than CFC-11 have ODPs that decrease (increase) with increasing integration time. Semi-empirical ODPs have also 

been developed (Solomon et al., 1992) so steady-state and 

time-dependent ODPs could have an observational basis. 

Some semi-empirical steady-state values have been revised in this chapter to reflect updates to fractional release 

values and to some lifetimes. In previous assessments, 

the semi-empirical approach and model calculations have 

been shown to yield similar values for most gases.

The advantages and disadvantages of the ODP met ODP metric have been discussed in previous WMO/UNEP reports 

(WMO, 1989, 1991, 1995, 1999, 2003, 2007). The pro The projected importance of non-halocarbon emissions to future 

ozone levels presents new challenges to the ODP concept 

and raises questions about their continued comprehensiveness. Emissions occurring in the stratosphere or upper troposphere from aviation and rockets present special challenges to the ODP concept, but may be able to be treated 

in a manner similar to how very-short-lived species are 

treated (as a function of where and when emissions occur). 

ODPs are discussed in more detail in Section 5.2.2 below.

Calculation of the GWP of a compound requires 

knowledge of its radiative efficiency and global lifetime. 

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 Watts per square meter per part per billion (W 

m−2 ppb−1) or Watts per square meter per kilogram (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 as well as atmospheric structure, surface temperature, and clouds. The 

Absolute Global Warming Potential (AGWP) for time horizon t' is defined as

 AGWPx(t') = ∫

0

Fx·x(t)dt (5-1)

where Fx is the radiative efficiency of species x, x(t) describes the decay with time of a unit pulse of compound x, 

and t' is the time horizon considered. Fx is given in terms 

t'
5.7

A Focus on Information for Policymakers

of W m−2 kg−1 or in W m−2 ppb−1. The AGWP usually has 

units of W m−2 kg−1 yr and quantifies the future integrated 

radiative forcing over the time horizon of a unit mass pulse 

emission of a greenhouse gas. To compare the relative 

integrated effect of various compounds on climate, the 

GWP metric was developed. The GWP for time horizon t'

(IPCC, 1990; 2001; 2007) can be defined as

(5-2)

where FCO2 is the radiative efficiency of 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 with a response time of τx. The 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 that decrease 

(increase) with increasing time horizon. As shown in equations (5-1) and (5-2), the most common definition of GWPs 

applies to pulsed emissions. However, GWPs have also 

been developed to evaluate the effect of sustained emissions (Berntsen et al., 2005; Johnson and Derwent, 1996).

We note that GWP is not the only metric available to compare the climatic impacts of different gases. 

The science of alternative metrics was considered by an 

Intergovernmental Panel on Climate Change (IPCC) expert panel recently (IPCC, 2009), which noted that metric 

design depends critically on the policy goal, and that “the 

GWP was not designed with a particular policy goal in 

mind.” Furthermore, the choice of time horizon used in 

calculating GWPs is not determined purely by climate science considerations. Rather, the choice often depends on 

what information is useful to decision makers, based in 

part on the time horizon of the impacts and on the values 

they consider most important. In an effort to account for 

the impact of the choice of time horizon, typically three 

time horizons have been considered (20, 100, and 500 

years) when reporting GWPs.

The GWP index has three major advantages over 

most other indices used to measure the contribution of 

greenhouse gases to global warming: transparency, simplicity, and widespread acceptance. However, it also has 

several disadvantages (see, e.g., IPCC, 2009 and references therein). There is growing recognition of the limitations of the GWP metric especially when the impacts 

of short- and long-lived pollutants need to be considered 

together (Johansson et al., 2008; Tanaka et al., 2009; 

Fuglestvedt et al., 2010) and it has been argued that it is 

time to consider whether other metrics might be more useful (Shine, 2009). For example, there has been interest 

in including the economics of emissions mitigation into a 

climate metric by applying cost-benefit and cost-effective 

approaches (e.g., Manne and Richels, 2001). Various 

alternatives have been presented to overcome some of the 

GWP limitations, but none has been widely accepted as a 

suitable replacement to date.

ODP and GWP metrics are often used to evaluate 

the relative integrated impacts arising from emissions, 

banks, production, etc., on ozone destruction and climate 

forcing. If time series of these potential impacts are desired, EESC and radiative forcing are used. These metrics 

are discussed in Sections 5.2.2 and 5.2.3.

5.2.2 Ozone Impacts: ODPs and EESC

5.2.2.1 Fractional Release Values and

Global Lifetimes

Fractional release values for various halocarbons 

are used both in the calculation of semi-empirical ODPs 

and in deriving EESC. The fractional release value at 

some location in the stratosphere quantifies the fraction 

of the source gas that has become photochemically degraded and has released its halogen atoms since it entered 

the stratosphere. The previous Assessment (Daniel and 

Velders et al., 2007) used fractional release values from 

earlier assessments, only updating CFC-114 to the value 

suggested by Schauffler et al. (2003). For this Assessment, we adopt most of the Newman et al. (2007) updated release values; these values are based on National 

Aeronautics and Space Administration (NASA) ER-2 

field campaign observations using an approach similar to 

Schauffler et al. (2003). Douglass et al. (2008) showed 

that the lower stratospheric relationships between the 

fractional release of chlorine from CFC-11 and CFC-12 

and the age of air (as discussed in Chapter 1) produced 

by simulations with realistic age-of-air values are in good 

agreement with relationships derived from aircraft observations (Schauffler et al., 2003). Another advantage of 

using the Newman et al. (2007) results is that they provide observationally based relationships between fractional release values and age-of-air values. This allows 

for more appropriate estimates of EESC that are relevant 

to various stratospheric locations, and in particular, polar 

regions. In the derivation of semi-empirical ODPs and of 

EESC presented in this chapter (see Table 5-1), we use 

the values of Newman et al. (2007) for all compounds 

in that study except for HCFC-141b and HCFC-142b. 

These two compounds were present in small abundances and had large temporal trends when the atmospheric 

measurements upon which these fractional release values are based were made, leading to large uncertainties. 

For this reason, we have retained the fractional release 

GWPx(t') = = AGWPx(t')

AGWPCO2

(t')

∫

∫

Fx exp

FCO2

 R(t)dt

dt

t'

t'

0

0

( ) −t

τx
5.8

Chapter 5

Table 5-1. Lifetimes, fractional release values, and Ozone Depletion Potentials (ODPs) for long-lived 

halocarbons. ODPs recommended in this Assessment and ODPs adopted in the Montreal Protocol are 

included a.

Halocarbon Lifetime

(years)

Fractional Release Value Semi-Empirical ODP ODP in

Montreal

WMO Protocol

(2007) b

This

Assessment c

WMO

(2007)

This

Assessment d

Annex A-I

CFC-11 45 0.47 0.47 1.0 1.0 1.0

CFC-12 100 0.28 0.23 1.0 0.82 1.0

CFC-113 85 0.35 0.29 1.0 0.85 0.8

CFC-114 190 0.13 0.12 1.0 0.58 1.0

CFC-115 1020 0.04 0.44 e 0.57 0.6

Annex A-II

Halon-1301 65 0.29 0.28 16 15.9 10.0

Halon-1211 16 0.55 0.62 7.1 7.9 3.0

Halon-2402 20 0.57 0.65 11.5 13.0 6.0

Annex B-II

CCl4 26 0.50 0.56 0.73 0.82 1.1

Annex B-III

CH3CCl3 5.0 0.51 0.67 0.12 0.16 0.1

Annex C-I

HCFC-22 11.9 0.16 0.13 0.05 0.04 0.055

HCFC-123 1.3 0.02 0.01 f 0.02

HCFC-124 5.9 0.02 0.022

HCFC-141b 9.2 0.34 0.34 g 0.12 0.12 0.11

HCFC-142b 17.2 0.17 0.17 g 0.07 0.06 0.065

HCFC-225ca 1.9 0.02 0.025

HCFC-225cb 5.9 0.03 0.033

Annex E

CH3Br 0.8 0.53 0.60 0.51 0.66 0.6

Others

Halon-1202 2.9 0.62 1.7

CH3Cl 1.0 0.38 0.44 0.02 0.02

a Ravishankara et al. (2009) have calculated an ODP for N2O of 0.017. b In the previous Assessment, fractional release values relative to CFC-11 were used, with CFC-11 assumed to be 0.84. For this table, WMO (2007) 

values are scaled to a CFC-11 value of 0.47 to allow easy comparison with the current values. c From Newman et al. (2007), values for 3-year-old age of air. d Semi-empirical ODP values are not updated in this Assessment for halocarbons whose fractional release values were not estimated in Newman et 

al. (2007). e Model-derived value, WMO (2003). f From Wuebbles and Patten (2009) MOZART 3-D model calculation. g Values relative to CFC-11 are retained from WMO (2007) because of large uncertainties associated with the Newman et al. (2007) estimates for 

these compounds.
5.9

A Focus on Information for Policymakers

values relative to CFC-11 assumed in WMO (2007) for 

these two HCFCs, and we scale them to the appropriate 

CFC-11 fractional release value depending on the age of 

air considered. Additional information regarding the use 

of fractional release values can be found in Chapter 1. In 

calculating the 2010 ODP-weighted anthropogenic emissions for the A1 baseline scenario (Section 5.3.1), the 

most significant fractional release changes are for, in order of decreasing importance, CFC-12, HCFC-22, carbon 

tetrachloride (CCl4), and halon-1211. A description of 

the sensitivity of EESC to fractional release and to other 

parameters can be found in Newman et al. (2007). We 

also use absolute fractional release values in this chapter, 

rather than normalizing all of the values as was done in 

previous assessments. We do this so the portion of EESC 

estimates attributable to chlorine and bromine source gases are representative of the actual amount of total inorganic chlorine (Cly) and total inorganic bromine (Bry), respectively, in the regions of the stratosphere considered.

Global lifetimes are also used in the calculation of 

ODPs and EESC. As discussed in Chapter 1, lifetimes 

of CFC-114 and CFC-115 have been substantially revised since the previous Assessment. There have also 

been smaller revisions to several of the HCFCs that are 

included in Table 5-1. Douglass et al. (2008) also find, 

based on results from two-dimensional (2-D) and threedimensional (3-D) models, that the range of atmospheric 

lifetime estimates for CFC-11 from models that best reproduce the observed relationship between age of air and 

fractional release is 56–64 years. The lifetime calculated 

for CFC-12 in this study was consistent with the value 

from the previous Assessment (100 years). Wuebbles et 

al. (2009) find consistent lifetimes for CFC-11, 54 years in 

the Model of Ozone and Related Tracers (MOZART) 3-D 

model (version 3.1) and 57 years in the current version of 

the University of Illinois Urbana-Champaign (UIUC) 2-D 

model, which includes an improved calculation of stratospheric age of air (Youn et al., 2010). In the previous three 

assessments (WMO, 1999, 2003, 2007), ODP derivations 

assumed an atmospheric lifetime for CFC-11 of 45 years 

based on analyses of observations (Cunnold et al., 1997; 

Volk et al., 1997) and older model results (Prinn and Zander, 1999). Because of the dearth of studies estimating 

ODS lifetimes with models that accurately calculate age 

of air, and because the Douglass et al. (2008) study only 

evaluated the lifetimes of CFC-11 and CFC-12, we continue to use a lifetime of 45 years for CFC-11; however, 

it should be recognized that in the future this lifetime may 

be assessed to be too small. The lack of comprehensive 

studies examining halocarbon lifetimes with models that 

accurately calculate age of air can be considered a gap in 

our current understanding. A full discussion of lifetime 

revisions can be found in Chapter 1.

5.2.2.2 Ozone Depletion Potentials

An updated list of ODPs is provided in Table 5-1 for 

a number of long-lived halocarbons. The primary change 

from the previous Assessment (Daniel and Velders et al., 

2007) is to update the semi-empirical ODPs by incorporating revised fractional release values and global lifetimes 

for some compounds. For ODP estimates in this chapter, 

fractional release values representative of air that has been 

in the stratosphere for 3 years are used. In the absence 

of new evidence, we continue to use a value of 60 for the 

relative effectiveness of bromine compared to chlorine for 

destroying ozone (α) at midlatitudes. For Antarctic calculations we continue to use a value of 65 (WMO, 2007). 

The semi-empirical ODP of CCl4 is 10% higher than in the 

previous Assessment owing to a larger fractional release 

value (see Table 5-1 and Section 5.2.2.1). Although there 

are still significant gaps in our understanding of the CCl4

budget, the recommended lifetime remains 26 years (see 

Chapter 1).

The ODP for HCFC-123 has been updated based 

on results from a three-dimensional model (Wuebbles and 

Patten, 2009). The calculated value of 0.0098 is similar 

to but smaller than the previous model-derived value of 

0.014 based on 2-D model results (WMO, 2003). The 

ODPs for several short-lived compounds have also been 

discussed in Chapter 1.

Looking beyond the analyses of ODPs for halocarbons, Ravishankara et al. (2009) have calculated an ODP 

for nitrous oxide (N2O) of 0.017 using a 2-D model. Although N2O has a relatively small ODP, future changes in 

emissions and atmospheric concentrations of N2O could 

have a significant effect on ozone compared with emissions of controlled ODSs because of the larger magnitude 

of N2O’s anthropogenic emissions. The impact of N2O 

emissions will be quantified in Section 5.4. The magnitude of past and future N2O ODP-weighted emissions 

leads to concerns that include influences on the timing of 

the recovery of ozone, the “background” ozone level, the 

distribution of stratospheric ozone depletion, and the possibility of future decreases in ozone due to increasing N2O 

(Ravishankara et al., 2009; Wuebbles, 2009).

5.2.2.3 Equivalent Effective Stratospheric

Chlorine

EESC has historically been used to relate measured 

surface mixing ratios of chlorine- and bromine-containing 

source gases to the stratospheric inorganic chlorine and 

bromine arising from these gases in key regions of the 

stratosphere and thus to the amount of ozone they will 

destroy (Daniel et al., 1995; WMO, 1995, 1999, 2003, 

2007). It accounts for the fact that bromine is more 
5.10

Chapter 5

efficient than chlorine (on a per-atom basis) at destroying stratospheric ozone and that different source gases 

release their chlorine and bromine at different rates and 

locations. Both integrated EESC changes and the date 

when EESC returns to 1980 levels have been used in previous assessments to quantify the relative impacts of various hypothetical ODS emission cases. This approach has 

worked well because all the major anthropogenic sources 

of stratospheric chlorine and bromine are thought to be 

known, and they all have reasonably well-understood 

global lifetimes and stratospheric dissociation rates. 

Recently, EESC was reformulated so that it accounts for 

the age-of-air spectrum and the age-of-air dependent fractional release values (Newman et al., 2007). This reformulation allows EESC to represent the total Cly and Bry

in various regions of the stratosphere.

The concept of EESC does have limitations because of uncertainties in the transport time, uncertainties 

in the spatial and temporal dependencies of the bromine 

efficiency for ozone destruction versus chlorine (α) (generally considered a constant), the efficiency of stratospheric halogen release from the source gas, and possible temporal changes in transport times and source gas lifetimes 

(some effects described in Newman et al., 2006, 2007). It 

should be noted that the EESC concept also does not explicitly account for changing atmospheric emissions and 

concentration of other relevant constituents that can also 

affect ozone, such as CO2 and methane (CH4). Daniel et 

al. (2010) have recently provided a method to incorporate 

N2O abundances into EESC. We will not adopt this method for the calculation of EESC in this Assessment because 

its publication was so recent, but we do show the impact 

of a hypothetical N2O emission phase-out on globally averaged total column ozone that can be compared with the 

impact of other ODS phase-outs on ozone. By comparing 

2-D model results with EESC, Daniel et al. (2010) showed 

that integrated EESC, despite its limitations, is proportional to the calculated integrated total ozone depletion 

reductions arising from various hypothetical ODS policy 

actions to within 30%.

The year 1980 is often taken as roughly representative of the time before the major stratospheric ozone 

losses occurred due to halocarbons. As a result, analyses 

are often based on the return of EESC to 1980 levels. If 

stratospheric ozone levels were affected by only halocarbons, the ozone layer would be expected to recover from 

human activities as anthropogenic ODSs are removed 

from the atmosphere. However, the actual picture of future levels of stratospheric ozone and the recovery from 

the past emissions of halocarbons is more complicated, 

making the use of a single date as a metric for total ozone 

recovery less meaningful. There are a number of activities and processes that are likely to substantially affect the 

future distribution and integrated amounts of ozone (see 

Section 5.4 and Chapter 3). Some of these potential activities and processes include rocket launches, aviation, 

climate-related geoengineering actions, and emissions of 

CO2, CH4, and N2O. So while the evolution of EESC in 

this chapter provides important information regarding the 

impact of certain human activities on stratospheric ozone, 

ozone is expected to recover at a different rate than this 

simple metric might imply.

The relative importance of the various future chlorine- and bromine-containing emissions, production, 

and banks are compared as in previous assessments using EESC in Section 5.4. New in this Assessment is that 

the impacts of some of these ODS cases are compared in 

terms of ozone depletion using the results of the Daniel 

et al. (2010) study (see Section 5.4.2). The intent of this 

approach is to provide information that can be used for 

determining which options are available and what their 

potential effects are on the future ozone layer. It will 

also allow us to compare the expected impacts of several 

potential ODS policy options with ozone changes from 

a broader range of activities, including expected future 

changes in greenhouse gases such as CO2, N2O, and CH4.

5.2.3 Climate Impacts: GWPs and 

Radiative Forcing

Direct GWPs are tabulated in this chapter’s Appendix 5A. The list of compounds evaluated in Appendix 

Table 5A-1 is intended to include potential replacements 

for the Montreal Protocol ODSs, but this list is not exhaustive. CO2 and N2O are also included in the table. 

Most source gases with atmospheric lifetimes of less than 

6 months are not believed to contribute significantly to 

radiative forcing and are not included in Appendix Table 

5A-1 (note that, in contrast to short-lived source gases, 

short-lived aerosols and tropospheric ozone are thought 

to contribute significantly to radiative forcing (IPCC, 

2007)). Indirect GWPs are discussed later in this section and are also tabulated in Appendix 5A (Appendix 

Table 5A-4). The uncertainty associated with the direct 

GWPs listed in Appendix Table 5A-1 is estimated to be 

±35% with 90% confidence (IPCC, 2007) reflecting uncertainties in the radiative efficiencies, lifetimes of the 

halocarbons, and uncertainties in our understanding of the 

carbon cycle (IPCC, 2001). It should be noted that because uncertainties in the carbon cycle make an important 

contribution to the total GWP uncertainty (IPCC, 2007), 

the relative climatic effects of different halocarbons (e.g., 

expressed as ratios of their GWPs) are likely accurate to 

much better than 35%.

We limit further discussion of GWPs to updates 

since the last Ozone Assessment. There are two reasons that updates have been made: (1) updates to the 
5.11

A Focus on Information for Policymakers

atmospheric lifetimes and (2) new radiative efficiency 

recommendations. Although the atmospheric concentration of CO2 continues to increase, its AGWPs were chosen to be the same as the IPCC (2007) Assessment so as 

to reduce confusion when comparing to values from that 

Assessment. The revisions to the lifetimes and radiative 

efficiencies discussed below affect GWP calculations and 

radiative forcing projections.

5.2.3.1 Lifetime Updates

The lifetimes of several HCFCs and of most 

HFCs, including hydrofluoroethers (HFEs), have been 

updated from the previous Assessment. The lifetime for 

CH3CH2OCF2CHF2 (HFE-374pc2) of 5 years given in 

the previous Assessment was erroneous; the lifetime for 

this compound is actually rather short (approximately 2 

months, see Chapter 1). Accordingly, this compound is 

now classified as a short-lived species. The lifetimes of 

CFC-114, CFC-115, and nitrogen trifluoride (NF3) were 

also updated in response to work by Prather and Hsu 

(2008). A discussion of these lifetime updates and others, 

and the distinction between short- and long-lived species, 

can be found in Chapter 1.

5.2.3.2 Radiative Efficiency Updates

The radiative efficiency values used in the previous Assessment, the currently recommended values, and 

the values presented in the relevant references for the 

new and revised compounds are presented in Table 5-2. 

Since the last Assessment (WMO, 2007), additional radiative efficiency data have become available for HFE43-10pccc124, HFC-365mfc, HFC-263fb, HFC-245eb, 

HFC-245ea, and HFC-329p (all calculated assuming 

constant vertical profiles). Rajakumar et al. (2006) have 

provided the first measurements of the lifetimes and the 

radiative efficiencies for HFC-263fb, HFC-245eb, and 

HFC-245ea. These radiative efficiencies (Table 5-2) 

have been adopted in the calculation of the GWPs in the 

appendix (Appendix Table 5A-1). Inoue et al. (2008) report a radiative efficiency of 0.23 W m−2 ppb−1 for HFC365mfc. This value is 10% larger than the value of 0.209 

W m−2 ppb−1 reported by Barry et al. (1997) that was used 

in the previous Assessment (WMO, 2007). Inoue et al. 

(2008) concluded that the approximately 10% difference 

in the intensity of the infrared (IR) absorption spectrum 

was the origin of the different radiative efficiencies in 

the two studies. There being no obvious reason to prefer 

either study, an average of the results from Barry et al. 

(1997) and Inoue et al. (2008) was adopted in the current 

Assessment.

Sulfuryl fluoride (SO2F2) is being used increasingly as a replacement for methyl bromide (see Chapter 1, Section 1.5.2.2). It has an atmospheric lifetime of 

about 36 years and a substantial GWP. Papadimitriou et 

al. (2008a) and Sulbaek Andersen et al. (2009) measured 

the IR spectrum of SO2F2 and reported radiative efficiencies of 0.222 and 0.196 W m−2 ppb−1, respectively, for 

this molecule. The integrated absorption cross sections 

reported by Papadimitriou et al. (2008a) and Sulbaek 

Andersen et al. (2009) over the range 800–1540 cm−1 are 

in excellent agreement (within approximately 1%) and 

both differ by only about 5% from that reported by Dillon 

et al. (2008). The approximately 10% difference in radiative efficiencies reported by Papadimitriou et al. (2008a) 

and Sulbaek Andersen et al. (2009) is probably due to the 

fact that Sulbaek Andersen et al. (2009) did not consider 

Table 5-2. Radiative efficiency estimates (W m−2 ppb−1) for seven compounds whose recommended 

values have either changed since, or were not included in, the previous Assessment.

Chemical Formula

Common or

Industrial

Name

WMO

(2007)

Recent

Published

Estimates

This

Assessment

CHF2OCF2OC2F4OCHF2 HFE-43-10pccc124 a 1.37 1.02 (Wallington et al., 2009) 1.02

CH3CF2CH2CF3 HFC-365mfc 0.21 0.23 (Inoue et al., 2008) 0.22

CHF2CF2CF2CF3 HFC-329p 0.45 0.31 (Young et al., 2009) 0.31

CH3CH2CF3 HFC-263fb 0.13 (Rajakumar et al., 2006) 0.13

CH2FCHFCF3 HFC-245eb 0.23 (Rajakumar et al., 2006) 0.23

CHF2CHFCHF2 HFC-245ea 0.18 (Rajakumar et al., 2006) 0.18

SO2F2 sulfuryl fluoride 0.222 (Papadimitriou et al., 2008a) 0.22

0.196 (Sulbaek Andersen et al., 2009)

a Also referred to as H-Galden 1040x.
5.12

Chapter 5

stratospheric adjustment (allowing the stratospheric temperatures to equilibrate to the greenhouse gas perturbation) in their radiative efficiency calculations. Sulbaek 

Andersen et al. (2009) used the estimation method described by Pinnock et al. (1995), which agrees to within 

±0.3% with the global annual and annual cloudy-sky instantaneous radiative forcing calculations using the Reading narrowband model (Shine, 1991). Papadimitriou et 

al. (2008a) applied a line-by-line radiative transfer code, 

developed at the National Oceanic and Atmospheric 

Administration (NOAA), to a global-average cloudy-sky 

profile taking into account stratospheric adjustment. The 

effect of including stratospheric adjustment for a given gas 

depends on the position and intensity of its IR absorption 

features. Forster et al. (2005) have shown that inclusion 

of stratospheric adjustment increases the instantaneous 

radiative forcing of HFC-134a by approximately 10%. 

Pinnock et al. (1995) studied 19 halogenated alkanes and 

found that inclusion of stratospheric adjustment typically 

increased the instantaneous radiative efficiency by 5–10% 

(although in the case of HFC-41 there was actually a 7% 

decrease, see Table 4 in Pinnock et al. (1995)). Radiative 

forcing as defined by IPCC is based on the globally and 

annually averaged net downward irradiance change at the 

tropopause after allowing for stratospheric temperatures 

to adjust to radiative equilibrium. Hence, we adopt a value of 0.22 W m−2 ppb−1 from Papadimitriou et al. (2008a), 

which is consistent with the value reported by Sulbaek 

Andersen et al. (2009).

Wallington et al. (2009) revisited the IR spectrum 

of HFE-43-10pccc124, previously also referred to as HGalden 1040x, and argued that the IR spectrum reported 

by Cavalli et al. (1998) used in previous WMO and IPCC 

reports is approximately 30% too strong. It was noted 

that use of the reference spectrum of Cavalli et al. (1998) 

would lead to carbon balances substantially greater than 

100% in laboratory experiments of HFE-43-10pccc124 

oxidation. The IR spectrum reported by Wallington et al. 

(2009) is in good agreement with that measured using the 

same experimental set-up 12 years earlier by Christidis et 

al. (1997). As discussed by Young et al. (2008), the radiative efficiency of HFE-43-10pccc124 in WMO (2007) is 

inconsistent with the expected trends based upon the database for other HFEs. The radiative efficiency reported by 

Wallington et al. (2009) is adopted here.

Young et al. (2009) reported the first measurements 

of the lifetime and radiative efficiency for CHF2CF2CF2CF3

(HFC-329p); these are used here.

As in all previous assessments, the atmospheric 

lifetime and radiative efficiency values used in the present Assessment can be traced to experimentally measured rate coefficients and IR spectra. Ab-initio quantum mechanical computational methods are available to 

calculate rate coefficients and IR spectra. Results from 

ab-initio calculations have provided valuable fundamental insight into reaction mechanisms and the underling 

processes giving rise to the IR spectra. However, such 

computational results have generally not been used directly to calculate GWPs. With advances in computational techniques, recent studies (e.g., Blowers and 

Hollingshead, 2009) suggest that atmospheric lifetimes 

and radiative efficiencies can now be calculated from 

first principles with accuracies comparable to those from 

experimental studies. This suggestion has yet to be fully 

investigated by the scientific community, and in the present Assessment we do not consider atmospheric lifetimes 

and radiative efficiencies that have been evaluated only 

by ab-initio methods.

5.2.3.3 Updates to Indirect GWPs from

Ozone Destruction

In addition to being greenhouse gases themselves, 

the ODSs play a significant role in the destruction of 

stratospheric ozone, another radiatively important gas. 

The distribution of ozone has important implications for 

Earth’s climate system, not only because ozone absorbs 

solar radiation but also because it is a greenhouse gas 

that absorbs some of the infrared radiation emitted from 

Earth’s surface and atmosphere. The change in ozone 

radiative forcing due to the addition of some ODS can 

be attributed to that ODS as an indirect radiative forcing. Stratospheric ozone losses are generally thought to 

cause a negative radiative forcing, canceling part of the 

increased radiative forcing arising from the direct influence of the halocarbons. The magnitude of the indirect 

effect is strongly dependent on the altitude profile of the 

halogen-induced ozone loss and will vary depending on 

the source gas considered. The latest estimate of radiative forcing from changes in stratospheric ozone since 

preindustrial times is −0.05 ± 0.10 W/m2 (IPCC, 2007). 

However, there have not been recent studies that clearly 

estimate the amount of the ozone forcing attributable to 

halocarbon changes only.

In spite of the uncertainty associated with the radiative forcing from ozone loss caused by halocarbons, the 

indirect GWPs for various halocarbons have been included 

in previous Ozone Assessments using an approach similar 

to that described in Daniel et al. (1995) (e.g., see WMO, 

2003, 2007). This approach is also used here to update 

indirect GWPs for selected halocarbons (Appendix Table 

5A-4) and is based on the assumption of an approximate 

linear relationship between the change in EESC arising 

from a particular source gas and radiative forcing arising 

from stratospheric ozone loss. A complication that is ignored by this approach is that some of the observed ozone 

depletion, and its associated radiative forcing, is due to 

processes not related to ODSs. The previously published 
5.13

A Focus on Information for Policymakers

indirect GWPs have primarily changed over time as a response to updates in the EESC scenarios, changes in the 

estimated efficiency of bromine relative to chlorine for destroying ozone, and changes in the estimates of the overall 

stratospheric ozone radiative forcing owing to halocarbon 

changes. These factors continue to represent important 

uncertainties in estimating the indirect GWPs using the 

EESC approach. This approach also cannot capture the 

source gas-dependent variations in the ozone loss profile 

and its resulting impact on radiative forcing. In past estimates of these indirect GWPs, the 1980 level of EESC 

has been assumed to represent a value below which no additional ozone loss occurs. This implied that the presence 

of an ODS after EESC fell below this level would lead to 

no additional ozone loss or negative radiative forcing due 

to that loss. No such threshold is adopted here, leading 

to slightly more negative indirect GWPs for the longerlived ODSs in the absence of other changes. However, 

the change in stratospheric ozone from 1979 to 1998 is 

now estimated to be responsible for −0.05 W/m2 forcing 

(IPCC, 2007). The adoption of this lower revised IPCC 

radiative forcing due to stratospheric ozone, compared 

with −0.15 ± 0.10 W/m2 in WMO (2007), dominates the 

changes compared to the previous Assessment and makes 

the indirect effects less negative. The current calculations 

also incorporate the updated fractional release values for 

midlatitudes and polar regions (Newman et al., 2007). 

We continue to assume that radiative forcing due to polar 

ozone loss is responsible for 40% of the −0.05 W/m2 and 

that the polar depletion saturates at 1990 EESC levels 

(Daniel et al., 1995).

As a step toward obtaining indirect GWPs through 

a more fundamental approach, Youn et al. (2009) have 

explicitly evaluated the indirect radiative forcing for two 

of the halons, halon-1211 and halon-1301, using 2-D 

and 3-D atmospheric models. In Table 5-3, these values 

are compared with the direct and indirect GWPs found 

in WMO, 2007 (Daniel and Velders et al., 2007) and 

those updated for this Assessment using the approach 

discussed in the previous paragraph. The indirect GWP 

for halon-1211 derived by Youn et al. (2009) is much 

smaller (about 60%) than the WMO (2007) value, but 

does fall within a single standard deviation of it. The 

Youn et al. (2009) indirect GWP is larger in magnitude 

by about a factor of ten compared to its direct effect. 

The indirect effect of halon-1301 is about 25% smaller 

than previously reported but also agrees with the WMO 

(2007) value to within a single standard deviation. The 

updated indirect GWPs reported here for halon-1211 calculated using the EESC scaling approach is in much better agreement with the Youn et al. (2009) values than the 

WMO (2007) central estimate was. While the Youn et 

al. (2009) work represents a single study, these analyses 

suggest that more comprehensive atmospheric models 

have the potential to reduce our uncertainties in the halocarbon indirect GWPs.

Table 5-3. Direct and indirect Global Warming Potentials (mass basis) of halons. Youn et al. (2009) lifetimes are e-folding times based on an exponential curve fitted to the simulated atmospheric decay.

Lifetimes (years)

Global Warming Potentials

for 100-years Time Horizon

Direct GWP Indirect GWP Net GWPa

Halon1211Halon1301Halon1211Halon1301Halon1211Halon1301Halon1211Halon1301

This

Assessment

16 65 1890

± 660

7140

± 2500

−11,720

± 23,430

−27,060

± 54,130

WMO (2007) 16 65 1890

± 660

7140

± 2500

−40,280

± 27,120

−49,090

± 34,280

Youn et al. (2009)

2-D Model

14.4 72.4 1796 7122 −16,294 −36,247 −14,498 −29,127

Youn et al. (2009)

3-D Model

10.9 70.1 1699 6903 −17,050 −37,252 −15,351 −30,349

a Concerns remain about the quantitative accuracy of adding the indirect forcing due to ozone changes to the direct forcing for the well-mixed halons. 

This addition was not performed in WMO (2007) or in this chapter, except for the results taken from Youn et al. (2009).
5.14

Chapter 5

5.3 FUTURE BASELINE SCENARIOS

5.3.1 Chlorine- and Bromine-Containing 

Ozone-Depleting Substances

A new baseline scenario for the period 1950–2100 is 

developed for all ODSs along the same lines as in the previous Assessment (WMO, 2007), with future projections 

consistent with current controls imposed by the Montreal 

Protocol. Observed global average mixing ratios through 

the beginning of 2009 are used as a starting point for the 

projections, together with the production of ODSs through 

2008 reported by the countries to the United Nations Environment Programme (UNEP, 2009), estimates of the bank 

sizes of ODSs for 2008 from the Technology and Economic Assessment Panel (TEAP, 2009), approved essential use exemptions for CFCs, critical-use exemptions for 

methyl bromide for 2008–2010, and production estimates 

of methyl bromide for quarantine and pre-shipment (QPS) 

use. Details of the baseline scenario are given in Appendix Table 5A-2. Calculated mixing ratios are tabulated in 

Appendix Table 5A-3 for each of the considered halocarbons from 1955 through 2100 (See also Figure 5-1). The 

years when observations are used are indicated in the table 

by the shaded regions. For those years, scenario mixing 

ratios are equal to the observations.

The mixing ratios in the new baseline scenario 

are similar to those in the previous Assessment (WMO, 

2007) for most species. The larger future mixing ratios (2010–2050) for CFC-11 (+5 to 9 parts per trillion 

(ppt)) and CFC-12 (up to +9 ppt) are the result of slightly 

larger fractions of the bank emitted annually, based on 

1999–2008 averages. The new baseline scenario has significantly larger future mixing ratios for CCl4 (up to +12 

ppt) than in WMO (2007) because of a different assumption regarding the decrease in future emissions (see Section 5.4.2.1). Emissions of each of the three considered 

HCFCs begin to decline at some time in the next decade 

in the baseline scenario due to existing Montreal Protocol 

controls. The projected future mixing ratios of HCFC-22 

after about 2025 are lower than projected in the previous 

Assessment as a direct result of the accelerated HCFC 

phase-out (Montreal Protocol Adjustment of 2007, see 

Section 5.4.5). The initially larger mixing ratios in the 

period 2010–2020 are the result of an increase in reported 

production and a larger fraction of the bank emitted annually. The mixing ratios of HCFC-141b and HCFC-142b 

are larger than in the previous Assessment, by up to 8 ppt 

and 20 ppt, respectively, because of an increase in reported production and a different assumption regarding the 

future distribution of the total HCFC consumption over 

the three main HCFCs (see Appendix Table 5A-2). The 

mixing ratios of halon-1301 in the new baseline scenario 

are somewhat different from those in the previous Assessment’s baseline scenario. The older ones were based on a 

bank release calculation after 1995 because of large differences in observational records that existed at that time for 

the 1995–2005 period. Revised calibrations and measurements with new instruments using mass spectrometry (at 

NOAA) have led to an improved consistency between the 

labs (NOAA and Advanced Global Atmospheric Gases 

Experiment (AGAGE)) for the global mean mixing ratios 

and the rates of change in recent years (see Figure 1-1 of 

Chapter 1). Therefore, the mixing ratios here are taken as 

an average of the observations from those two networks 

in recent years and are derived using a consistent scaling to the NOAA data in years before AGAGE measured 

halon-1301 (before 2004).

The EESC corresponding to the baseline scenario is 

shown in Figure 5-2. The absolute values of EESC (righthand axis) are substantially lower than presented in the 

previous Assessment because different fractional release 

values have been adopted and they are no longer scaled so 

the fractional release value of CFC-11 is 0.84. As stated 

in Section 5.2.2.1, this approach is taken because it allows 

the values now to provide meaningful estimates of Cly and 

Bry abundances for 3-year-old air, a benefit that would be 

lost if the former scaling method were applied. Furthermore, since EESC is used in this chapter only as a relative 

measure, the absolute values are not important. If current 

EESC values are scaled so the peak value is the same as 

in the previous Assessment, the figure shows that the time 

series are quite similar. Differences appear because of the 

revised relative fractional release values and because of 

slightly revised emissions. In previous assessments, the 

year 1980 was taken as a benchmark year, representing a 

time before the major stratospheric ozone losses occurred 

due to halocarbons. Some ozone losses caused by human 

activities likely had already occurred due to, for example, 

halocarbon and N2O emissions (see Chapter 3). Nevertheless, the year EESC returns to 1980 levels has been used 

as a measure for some degree of ozone layer recovery and 

to compare scenarios. Although we continue to recognize 

the limitations of this metric (see Section 5.2.2) it is used 

here to compare the new baseline scenario with the one 

in the previous Assessment and to compare the effects of 

hypothetical ODS emission reductions of chlorine- and 

bromine-containing ODSs (Table 5-4). In the new baseline scenario, including only Montreal Protocol-controlled 

ODSs, EESC returns to its 1980 levels in 2046 for midlatitude conditions. This is 2–3 years earlier than in the 

previous Assessment (WMO, 2007) and is the result of 

two partially offsetting changes. First, future emissions of 

ODSs give a delay of 2–3 years, mainly as the net result of 

larger future emissions of CCl4, smaller future emissions 

of HCFCs owing to the accelerated HCFC phase-out (2007 

Adjustment of the Montreal Protocol, see Section 5.4.5), 
5.15

A Focus on Information for Policymakers

CFC-11

1980 2000 2020 2040

Year

140

160

180

200

220

240

260

280

ppt

CFC-12

1980 2000 2020 2040

Year

300

350

400

450

500

550

ppt

CFC-113

1980 2000 2020 2040

Year

20

40

60

80

100

ppt

CFC-114

1980 2000 2020 2040

Year

10

12

14

16

18

ppt

CCl4

1980 2000 2020 2040

Year

20

40

60

80

100

120

ppt

CH3CCl3

1980 2000 2020 2040

Year

0

50

100

150

ppt

HCFC-22

1980 2000 2020 2040

Year

0

50

100

150

200

250

300

350

ppt

HCFC-141b

1980 2000 2020 2040

Year

0

10

20

30

40

ppt

HCFC-142b

1980 2000 2020 2040

Year

0

10

20

30

40

ppt

halon-1211

1980 2000 2020 2040

Year

0

1

2

3

4

5

ppt

halon-1301

1980 2000 2020 2040

Year

0

1

2

3

4

ppt

CH3Br

1980 2000 2020 2040

Year

6.5

7.0

7.5

8.0

8.5

9.0

9.5

ppt

Figure 5-1. Historical and projected mixing ratios (in parts per trillion, ppt) of selected ODSs for the new baseline (A1) scenario (black) compared with the baseline scenarios of the previous Assessment (WMO, 2007) 

(solid red) and for WMO (2003) (dashed red). Shaded regions indicate years when mixing ratios are taken 

directly from observations (see Chapter 1).
5.16

Chapter 5

and a smaller 1980 mixing ratio of CH3Br (see Chapter 1, 

Section 1.2.1.6). Second, the use of revised fractional release values based on recent measurements brings forward 

the return to 1980 levels by 5–6 years. For Antarctic conditions EESC returns to its 1980 levels in the baseline in 

2073. This is 7–8 years later than in the previous Assessment (WMO, 2007), mainly due to the use of fractional 

release values representative for Antarctic conditions (age 

of air = 5.5 years) and of a smaller 1980 mixing ratio of 

CH3Br (Chapter 1, Section 1.2.1.6), with smaller contributions from the revisions in CCl4 and HCFCs emissions.

5.3.2 CO2, CH4, and N2O

The importance of non-ODS greenhouse gases to 

future ozone levels has been quantified in several published articles and is discussed in Chapter 3. In this chapter, IPCC Special Report on Emissions Scenarios (SRES) 

(IPCC, 2000) scenario A1B is used to prescribe the future 

evolution of these three gases in the two 2-D model runs 

described in Daniel et al. (2010). The results of that study 

are used here to evaluate the impact of hypothetical future 

controls involving N2O and the ODSs on globally averaged total column ozone against a backdrop of CO2 and 

CH4 changes and their associated impacts on ozone.

5.3.3 ODP- and GWP-Weighted Emissions, 

EESC, and Radiative Forcing

The contributions of ODSs to metrics relevant 

to ozone depletion and to climate change are shown in 

Figures 5-3 and 5-6. The panels of Figure 5-3 show the 

ODP-weighted emissions, GWP-weighted (100-year) 

emissions, EESC, and radiative forcing (RF) for the A1 

baseline scenario from 1980–2100. In terms of EESC 

Figure 5-2. Midlatitude EESC projections (in parts per trillion, ppt) calculated for the current baseline (A1) 

scenario (solid black) and three test 

cases: zero future production (dotted), full recovery and destruction of 

the 2010 bank (dashed), and zero 

future emission (dot-dashed) (all 

right-hand axis). EESC values for 

the baseline scenario of the previous 

Assessment are also shown (red, lefthand axis). For ease of comparison, 

the two ordinate axes are relatively 

scaled so the EESC peak of the current baseline scenario falls on top of 

the previous baseline scenario peak. 

Absolute EESC values are different in 

the two Assessments because in the 

previous Assessment, different relative halogen fractional release values 

were used, and the values were scaled so the CFC-11 absolute fractional release value was 0.84. No scaling 

was performed here so that EESC is more directly representative of Cly and Bry.

Table 5-4 (at right). Comparison of scenarios and cases a: the year when EESC drops below the 1980 

value for both midlatitude and Antarctic vortex cases, and integrated EESC differences (midlatitude 

case) relative to the baseline (A1) scenario b. Also shown are changes in integrated ODP- and GWPweighted emissions and, for selected cases, changes in integrated ozone depletion from 2011−2050. Future 

projected changes in CH4 and CO2 are also expected to significantly alter ozone levels, perhaps by amounts 

larger than any of the cases considered in this table. However, their effects are not included here because 

changes in CH4 and CO2 that would decrease globally averaged ozone depletion would increase climate forcing. Values cited in chapter text calculated as differences between table entries may appear slightly inconsistent with table entries because of rounding.

Midlatitude EESC

1980 2000 2020 2040 2060 2080 2100

Year

0

1000

2000

3000

4000

EESC in Previous Assessment (ppt)

0

500

1000

1500

2000

EESC in Current Assessment (ppt)

Baseline, this Assessment

No Production

Bank Recovery

No Emission

Baseline, previous Assessment
5.17

A Focus on Information for Policymakers

Scenario and Cases a

Percent Difference in

Integrated EESC

Relative to Baseline

Scenario for the

Midlatitude Case

Year (x) When EESC is

Expected to Drop Below

1980 Value b

Change in

ODPWeighted d

Emission:

2011−2050

Change

in GWPWeighted e

Emission:

2011−2050

Percent

Difference in

Integrated O3

Depletion f:

2011−2050

Midlatitude c Antarctic

Vortex c

x

∫ EESC dt

x

∫ EESC dt (Million tons 

CFC-11-eq)

(Billion tons

CO2-eq)

Scenarios

A1: Baseline scenario - - 2046.5 2072.7 - - -

Cases of zero production from 2011 onward of:

P0: All ODSs −5.4 −15.0 2043.0 2069.7 −0.70 −13.2

CFCs 0.0 0.0 2043.0 2069.7 0 0

Halons 0.0 0.0 2043.0 2069.7 0 0

HCFCs −3.2 −8.8 2044.5 2071.8 −0.45 −13.2 −0.15

CH3Br for QPS −2.4 −6.7 2044.9 2070.8 −0.26 −0.002

Cases of zero emissions from 2011 onward of:

E0: All ODSs (does not 

include N2O)

−16.0 −43.0 2033.5 2059.4 −3.82 −27.2 −0.67

CFCs −3.9 −11.0 2043.3 2068.5 −1.27 −7.9

Halons −5.0 −14.0 2043.3 2069.4 −1.09 −0.4

HCFCs −4.9 −13.0 2043.8 2071.4 −0.66 −18.1

CCl4

g −2.8 −7.6 2044.6 2071.6 −0.54 −0.9

CH3CCl3 0.0 −0.1 2046.5 2072.7 −0.004 −0.004

CH3Br for QPS −2.4 −6.7 2044.9 2070.8 −0.26 −0.002 −0.09

Anthropogenic N2O h −6.0 −103 −0.35

Cases of full recovery of the 2011 banks of:

B0: All ODSs −10.0 −27.0 2039.3 2064.6 −2.57 −13.1

CFCs −3.9 −11.0 2043.3 2068.5 −1.27 −7.9 −0.13

Halons −5.0 −14.0 2043.3 2069.4 −1.09 −0.4 −0.15

HCFCs −1.8 −4.8 2045.8 2072.4 −0.22 −4.9 −0.07

Cases of full recovery of the 2015 banks of:

B0: All ODSs −7.1 20.0 2040.3 2065.7 −2.03 −11.3

CFCs −2.6 −7.0 2044.0 2069.4 −0.95 −5.5

Halons −3.3 −9.1 2043.8 2069.8 −0.83 −0.3

HCFCs −1.9 −5.3 2045.5 2072.2 −0.26 −5.5

CH3Br sensitivity:

Same as A1, but 

critical-use exemptions 

continue at 2011 levels

+0.3 +0.8 2046.6 2073.0 +0.03 +0.0003

a Significance of ozone-depleting substances for future EESC were calculated in the hypothetical “cases” by setting production or emission to zero in 

2011 and subsequent years or the bank of the ODSs to zero in the year 2011 or 2015. b When comparing to WMO (2007), revised fractional halogen release values have contributed to an earlier EESC recovery year for midlatitudes and a 

later one for the Antarctic vortex. c For midlatitude conditions, an age-of-air value of 3 years, corresponding fractional release values, and a value for α of 60 are used. For Antarctic 

vortex conditions, an age-of-air value of 5.5 years with corresponding fractional release values and a value for α of 65 are used. d Using semi-empirical ODPs from Table 5-1. e Using direct GWPs with 100-year time horizon (see Appendix Table 5A-1). f Integrated globally averaged total column ozone changes are taken from Daniel et al. (2010). g Banks are assumed to be zero. Emissions include uncertain sources such as possible fugitive emissions and unintended by-product emissions. h The integrated ODP- and GWP-weighted emissions correspond to the elimination of all anthropogenic N2O emissions from 2011 onward in the A1B 

SRES scenario.

1980 2011
5.18

Chapter 5

Figure 5-4 (at right). Banks of CFCs, HCFCs, and halons in 1995, 2008, and 2020 in megatonnes (Mt) and 

weighted by their ODPs and 100-year GWPs (direct term only). The 2008 bank is from TEAP (2009) for almost 

all compounds. The 1995 bank is derived by a backwards calculation using the historic emissions, UNEP production data, and the 2008 bank. The 2020 bank is calculated in the baseline scenario from estimates of future 

production and bank release rates. The pie slices show the sizes relative to those of the maximum values (1995 

for the ODP- and GWP-weighted banks and 2008 for the bank in Mt) and therefore don’t fill a full circle. The 

total bank size is given below each pie chart. For comparison, the HFC bank size has been estimated to have 

been 1.1 GtCO2-eq in 2002 and 2.4 GtCO2-eq in 2008, and has been projected to be 4.6 GtCO2-eq in 2015 in 

a business-as-usual scenario (Velders et al., 2009).

1980 2000 2020 2040 2060 2080 2100

Year

0. 0

0. 5

1. 0

1. 5

2. 0

ODP-Weighted Emissions

Natural CH3Br and CH3Cl

CFCs

Halons

CCl4 + CH3CCl3

HCFCs

Anthropogenic

CH3Br

1980 2000 2020 2040 2060 2080 2100

Year

0. 0

0. 5

1. 0

1. 5

2. 0

2. 5

EESC (ppb)

EESC

Natural CH3Br and CH3Cl

CFCs

Halons

CCl4

CH3CCl3

HCFCs

Anthropogenic

CH3Br

1980 2000 2020 2040 2060 2080 2100

Year

0

2

4

6

8

10

GWP-Weighted Emissions

(100-yr)

CFCs HCFCs

Other ODSs

1980 2000 2020 2040 2060 2080 2100

Year

0.00

0.05

0.10

0.15

0.20

0.25

0.30

0.35

0.40

Radiative Forcing

CFCs

HCFCs

Other ODSs

RF (W

m−2)

Emissions (GtC

O2-eq yr−1) Emissions (MtCFC-11-eq yr−1)

Figure 5-3. ODP- and GWP-weighted emissions, EESC, and radiative forcing of the chlorine- and brominecontaining ODSs of the baseline (A1) scenario. Only the direct GWP-weighted emissions and radiative forcing 

are considered here. Semi-empirical steady-state ODPs (Table 5-1) and 100-year GWPs (Appendix Table 

5A-1) are used to weight emissions. The emissions and radiative forcing of the HFCs are not included in these 

graphs because of the large range in HFC scenarios. They are shown in Figure 5-5.
5.19

A Focus on Information for Policymakers

contribution, the largest contribution in the past and future comes from the CFCs and natural CH3Cl and CH3Br 

emissions. In terms of radiative forcing the CFCs are also 

the most important ODSs, but the HCFCs are projected to 

contribute more than 0.05 W/m2 at their peak.

CFC, HCFC, and halon bank size estimates for 

1995 and 2008 and projections for 2020 are shown in Figure 5-4. The contributions of the ODSs to the bank are 

compared by mass and when weighted by their ODPs and 

GWPs (100-year). The total size of the bank by mass remains more or less constant over time, but decreases significantly when considering their potential effect on depletion of the ozone layer (ODP-weighted) and on climate 

forcing (GWP-weighted). The decreasing contribution 

of the CFCs and increasing contribution of the HCFCs is 

particularly apparent when examining the banks by mass 

and when weighted by GWPs. The relative contributions 

of both the CFCs and halons remain the largest when considering the ODP-weighted banks.

5.4 IMPACTS OF HUMAN ACTIVITIES 

RELEVANT TO OZONE POLICY

5.4.1 Background

In this section, we assess the importance of various 

human activities that are relevant to ozone policy formulation. The focus is primarily on the extent to which future ozone depletion may be affected. We also discuss the 

potential effect of HFCs on climate forcing because the 

widespread use of HFCs has occurred directly as a result 

of ODS controls of the Montreal Protocol. We discuss 

geoengineering using atmospheric sulfur injections, and 

CO2 and CH4 emissions not because they necessarily have 

relevance for ozone policy decisions, but to assess their 

potential for inadvertent ozone impacts as a result of climate change policies or the absence of such policies.

1995

Bank in Mt

CFCs

(79%)

Halons

(5%)

HCFCs

(16%)

4.5 Mt

2008

CFCs

(39%)

Halons

(3%)

HCFCs

(59%)

4.8 Mt

2020

CFCs

(18%)

Halons

(1%)

HCFCs

(81%)

4.7 Mt

Bank ODP-weighted

CFCs

(58%)

Halons (41%)

HCFCs (1%)

5.6 MtCFC-11-eq

CFCs

(52%)

HCFCs Halons (43%) (6%)

3.4 MtCFC-11-eq

CFCs

(44%)

Halons

(40%)

HCFCs

(16%)

1.9 MtCFC-11-eq

Bank GWP-weighted

CFCs (93%)

Halons (3%) HCFCs (4%)

26 GtCO2-eq

CFCs

(71%)

HCFCs (26% Halons (3%) )

16 GtCO2-eq

CFCs

(44%)

Halons

(3%)

HCFCs

(53%)

 9 GtCO2-eq
5.20

Chapter 5

5.4.2 Ozone Impacts

The Montreal Protocol and its Amendments and 

Adjustments have effectively controlled the emission 

of the most important ozone-depletion gases containing chlorine and bromine. It is a testament to the success of the Montreal Protocol that current potential policy 

options and activities that do not involve chlorine- or 

bromine-containing compounds are expected to have 

some of the greatest effects on future ozone levels. However, to maintain consistency with past reports and to 

address the relative importance of continued emissions of 

halogenated source gases, we provide in Section 5.4.2.1 

an analysis of various hypothetical test cases using EESC 

as the comparison metric.

Chapter 8 of the previous Assessment (Daniel and 

Velders et al., 2007) focused strictly on the EESC index to 

quantify the potential effects of additional ODS controls 

on stratospheric ozone depletion. Here, results from 2-D 

model calculations in Daniel et al. (2010) are presented to 

augment the historical reliance on EESC and as a way to 

assess the ozone impact of several different assumptions of 

future emissions of halocarbons (e.g., CFCs, HCFCs, halons, CH3Br) and N2O (Section 5.4.2.2 and Chapter 3) (see 

Table 5-4). The phase-out and baseline scenarios assumed 

in Daniel et al. (2010) were similar to those in this chapter. However, that paper should be consulted for a consistent comparison of ozone changes with EESC changes. 

Estimates of ozone impacts from ODSs and N2O future 

emission changes also allow for comparison with other 

activities and emissions, such as CO2 and CH4. The computational efficiency of 2-D models makes it possible now 

to perform multiple long-term simulations in a reasonable 

amount of time to examine the effects of various scenarios. 

The residual circulation framework used in 2-D models has 

been shown to provide realistic simulations of atmospheric 

transport on long timescales (>30 days). Recent studies 

have demonstrated good overall model agreement with a 

variety of observations in reproducing transport-sensitive 

features in the meridional plane (e.g., Fleming et al., 2007). 

These capabilities make these models useful tools for 

studying long-term ozone changes.

5.4.2.1 Chlorine- and Bromine-Containing ODSs

Alternative test cases, based on the baseline scenario of ODSs (Section 5.3.1), are constructed and used 

to examine the relative effects of reducing future emissions of groups of halocarbons on EESC and stratospheric 

ozone levels. These alternative cases fall into four categories: (1) “No future emission” cases; (2) “No future 

production” cases; (3) cases assuming the full capture and 

destruction of the 2011 bank; and (4) cases assuming full 

capture and destruction of the 2015 bank. For the bank 

cases, the entire bank is considered, not just the portion 

that is estimated to be “accessible.” The 2015 bank recovery cases are run to quantify the impact of the implementation date of hypothetical bank emission controls; it 

should be recognized, however, that we do not have an 

estimate of the uncertainty in the future bank projections, 

on which these calculations are based. Full elimination 

of emissions, production, and banks are considered rather 

than some smaller fractional reduction because the intent 

is to provide policy-neutral cases that can be used to determine the impacts of other specific policy options. The 

impact of any option that involves a smaller reduction can 

be determined by simply scaling the results of Table 5-4. 

For example, if the “accessible” bank is estimated to be 

30% of the total bank, the results in the table for that compound can be simply multiplied by 0.3 to determine the 

impact of capture and destruction of the accessible bank. 

The elimination of all future emissions represents the 

greatest possible reduction in future abundances that can 

be imparted by direct controls. The elimination of future 

production 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 2011 bank complement the “No Production” scenarios in that the full 2011 

bank is destroyed in a manner that releases no ODSs to 

the atmosphere; future production and the resulting future 

buildup of banks is allowed to continue, however. The 

expected future production is virtually zero for the CFCs 

and halons, but still significant for the HCFCs and possibly for CCl4. The “No production” and “Full capture and 

destruction of the 2011 bank” cases are complementary 

in the sense that they may be combined to approximately 

estimate the “No emission” cases.

The effects of the alternative test cases are presented in Table 5-4 using metrics similar to those of previous assessments, i.e., the year EESC returns to its 1980 

level for midlatitude and Antarctic vortex conditions, and 

the change in integrated EESC relative to the baseline 

scenario. These integrated values are dependent on the 

choice of the 1980 level. If a lower level were chosen, 

for example the 1975 level, the tendency is for the CFC 

options to gain in importance relative to the other ODS 

policy options. This results from the longer atmospheric 

lifetimes of the CFCs. Also in this table is the change in 

cumulative ODP- and GWP-weighted (direct contribution 

only) emissions. The latter gives an indication of the potential effects of the cases for climate forcing. In general, 

the various cases presented here lead to smaller integrated 

EESC differences (when integrated from 1980) and return 

year differences from the baseline scenario than the same 

test cases from WMO (2007). This is expected, as banks 

of most ODSs are smaller due to continued emission for 

4 more years, and there are now fewer years to continue 
5.21

A Focus on Information for Policymakers

to produce the various ODSs before the mandated phaseout. CCl4 represents an exception to this. The increasing 

apparent importance of a CCl4 production or consumption phase-out arises from projected CCl4 emissions being larger in the current baseline scenario than in WMO 

(2007). The increased future emissions are based on an 

extrapolation of declining global emissions over the last 

5 years (6%/year). This decline in CCl4 emissions has 

not mirrored the strong decrease in reported production 

of CFC-11 and CFC-12 as it did in the 1980s and 1990s. 

This revised projection is thus not based on an improved 

understanding of the budget of CCl4, which still cannot be 

balanced nor can the interannual variability be explained 

with our current knowledge of CCl4 sources and sinks (see 

Chapter 1). The incomplete understanding of the CCl4

budget represents a significant gap.

Emissions of CFCs and halons from the banks are 

the largest sources of current ODP-weighted emissions 

of ODSs. An immediate phase-out of halon bank emission sources reduces future integrated EESC more than 

the elimination of emissions of any other compound or 

compound group, with slightly more of an impact than the 

elimination of future HCFC emissions. A delay of four 

years, from 2011 to 2015, in the capture and destruction 

of the estimated banks of CFCs and halons is currently 

thought to reduce the potential ozone and climate benefits 

by about 30% since ODSs continue to be emitted from 

these banks and production is thought to have essentially 

ceased.

A phase-out of methyl bromide emissions from 

quarantine and pre-shipment (QPS) applications beginning 

in 2011 would shift the year midlatitude EESC reaches 

the 1980 level earlier by 1.5 years compared to continued 

use at current levels. Continuing critical-use exemptions 

(CUE) at the approved 2011 level would delay the return 

of EESC to 1980 levels by 0.2 years.

The elimination of all emissions of chlorine- and 

bromine-containing ODSs after 2010 would shift the year 

midlatitude EESC reaches the 1980 level by about 13 

years, from 2046 to 2033. In terms relevant to climate, 

this would reduce emissions of these compounds by about 

0.7 gigatonnes (Gt) of CO2-equivalent (GtCO2-eq) per 

year averaged over 2011 through 2050. Future production of HCFCs and the sum of the current banks of CFCs 

plus HCFCs contribute about equally to this number. In 

comparison, global anthropogenic emissions of CO2 were 

greater than 30 Gt per year in 2008.

In past assessments, the potential impact of 

chlorine-containing VSLS has been neglected in the evaluation of hypothetical ODS emissions phase-outs. However, it is likely that their contribution is not negligible. 

As discussed in Chapter 1, anthropogenic VSLS are currently expected to contribute about 39 ppt chlorine to the 

stratosphere with fairly large uncertainty. Because these 

likely release their chlorine quickly, this amount could 

contribute directly to EESC with a unit fractional release 

value. If included in EESC in this way, at midlatitudes, 

a phase-out of these compounds is expected to advance 

the return of EESC to 1980 levels by almost 3 years. The 

magnitude of this effect is significant compared to the 

long-lived ODS cases presented in Table 5-4. This impact 

depends strongly on the region of the stratosphere considered, with greater impacts in regions where less chlorine 

has been liberated from long-lived source gases, but where 

chemical ozone loss is still important. It is also unclear 

how VSLS emission reductions in specific locations and 

during different times of the year would quantitatively affect the abundance of stratospheric chlorine.

The uncertainties of the effects of the test cases presented in Table 5-4 have been estimated by varying several model parameters for each class of compounds (CFC, 

HCFCs, etc.) in the calculation: the size of the banks in 

2008, the annual emission rate from the banks, the fractional release values, and the lifetimes of the compounds 

(Douglass et al., 2008; Martinerie et al., 2009). The lifetimes and annual emission rate from the banks are only 

varied for 2009–2100, since the data through 2008 is constrained by observations. A variation of 20% in the size of 

the 2008 banks of CFCs, HCFCs, and halons results in a 

variation of about 20% in the change in integrated EESC, 

for each class of compounds, in the case of a full recovery 

of the banks. The year EESC returns to the 1980 level varies by less than a year. Doubling the annual emission rate 

from the banks increases the change in integrated EESC 

by up to 15%, while reducing the annual emission rate by 

50% decreases the changes in integrated EESC by up to 

30%. Again, for most cases, the year EESC returns to the 

1980 level varies by less than a year. Variations of 20% in 

the fractional release values per class of compound gives 

similar variations. Variations of up to ±3 years occur in 

the year EESC returns to the 1980 level when varying the 

lifetimes of the CFCs by ±20%. The effects of such a 

variation on the change in integrated EESC is less than 

±20% for each compound group. These uncertainties have 

been applied to entire groups of compounds. These uncertainties are likely independent across compound groups 

for the size of the banks and annual emission rates from 

the banks, but perhaps not for the fractional release values 

and lifetimes.

The benefits of the Montreal Protocol that are calculated to have already been achieved are highlighted in 

Section 5.5. The cases considered here show that the options for limiting future halogenated ODS emissions are 

expected to have a much smaller impact on future ozone 

depletion and climate than what has already been accomplished by the Montreal Protocol.
5.22

Chapter 5

5.4.2.2 CO2, CH4, and N2O

The increasing concentration of carbon dioxide 

(CO2) in Earth’s atmosphere warms the troposphere but 

cools the stratosphere. This cooling leads to slower gasphase ozone loss reactions (Haigh and Pyle, 1979; Rosenfield et al., 2002), and thus to an increase in stratospheric 

ozone. Cooling of the polar lower stratosphere can also 

lead to enhanced polar stratospheric cloud (PSC) formation and more ozone destruction via more efficient heterogeneous chemistry (Chapter 3, Section 3.2.3). Increasing concentrations of methane and its effects on hydrogen 

oxides can enhance the destruction of ozone in the upper 

stratosphere. Increases in methane (CH4) can also reduce 

the ozone loss in the stratosphere by converting chlorine 

from active species to the reservoir HCl, and can lead to 

NOx-induced ozone production in the troposphere and 

lower stratosphere via the CH4 oxidation cycle. Increases 

in methane and the subsequent oxidation to water vapor 

will also lead to weak stratospheric cooling. Future levels 

of CO2 and CH4 may be higher for some period of time 

compared with today. If they are, they would be expected 

to lead to higher levels of column ozone at midlatitudes 

(e.g., Li et al., 2009; Waugh et al., 2009; Hsu and Prather, 

2010). Thus, there is the possibility of a “super-recovery,” 

where the total amount of atmospheric ozone exceeds 1960 

levels. Both 2-D (Rosenfield et al., 2002; Chipperfield 

and Feng, 2003; Portmann and Solomon, 2007; Daniel et 

al., 2010), and 3-D model (SPARC CCMVal, 2010) calculations suggest that increases in CO2 and CH4 will lead 

to total column ozone increases that are larger than what 

can be caused by decreases in future ODS or nitrous oxide 

(N2O) emissions. Because increasing levels of CO2 and 

CH4 are expected to continue to lead to increased total 

ozone, but also increased climate forcing, we present this 

information as relevant for inadvertent ozone impacts due 

to climate policy choices. Further complicating the matter 

is that, for ozone, changes in circulation due to greenhouse 

gases are expected to be the dominant factor in the tropics, and this will lead to lower column ozone levels there. 

Additional discussion of the impacts of CO2 and CH4 on 

ozone can be found in Chapter 3.

N2O emissions reductions, in contrast, would reduce climate forcing, and the resulting decrease in stratospheric nitrogen oxides would lead to higher ozone levels. 

Ravishankara et al. (2009) have shown that when ODS 

emissions are weighted by their steady-state ODPs, N2O 

emissions represent the single most important anthropogenic emissions today for ozone depletion. Unmitigated, 

these emissions are also expected to remain the most significant throughout the 21st century. In the current atmosphere, CFCs lead to much more ozone depletion than 

does N2O because of the large historic emissions and long 

lifetimes of the CFCs. The 2-D models used in Daniel et 

al. (2010) suggest that N2O causes an ozone depletion in 

the current atmosphere of about 3 Dobson units (DU) 

compared with the ODSs depleting about 18 DU. However, when the impact of phase-outs of future emissions 

are considered, an immediate phase-out of all anthropogenic N2O emissions leads to a greater reduction in ozone 

depletion integrated through 2100 than the integrated 

ozone depletion reduction arising from an elimination of 

all future ODS emissions (Daniel et al., 2010). The climate forcing and ozone impacts of a hypothetical emissions phase-out beginning in 2011 of several ODSs and 

anthropogenic N2O are compared in the final three columns of Table 5-4. These comparisons show the relative 

significance of these policy options using cumulative 

ODP- and GWP-weighted emissions and ozone integrated 

from 2011–2050. The significance of N2O compared to 

the ODSs is not surprising given the higher current ODPweighted emissions compared to the controlled ODSs. 

We emphasize that all the phase-outs considered are purely hypothetical with no consideration given to the economic or technological feasibility of implementing them.

5.4.2.3 N2O from Automotive Biofuels

Anthropogenic emissions of N2O total approximately 6.7 TgN yr−1 (IPCC, 2007). Agriculture is responsible for 2.8 TgN yr−1 and related emissions from rivers, 

estuaries and coastal zones add a further 1.7 TgN yr−1. 

Emissions associated with fossil fuel combustion and 

industrial activities, biomass burning, atmospheric deposition, and human excreta combined account for the remaining 2.2 TgN yr−1. The Greenhouse Gases, Regulated 

Emissions, and Energy Use in Transportation (GREET) 

model provides a life cycle assessment of biofuel production and uses, and includes allocation and differentiation 

of biofuels based on different agricultural practices (Wang, 

2010). The GREET model provides N2O emission factors 

for corn ethanol, sugarcane ethanol, and soy biodiesel in 

the United States of America (USA) of approximately 45, 

25, and 17 g/MMBtu of fuel produced (MMBtu = million 

British thermal units, 1 MMBtu = 1.055 gigajoules (GJ)). 

The USA is the largest biofuel producer, and corn ethanol 

is the dominant biofuel produced in the USA and has the 

highest N2O emissions per megajoule (MJ) of fuel produced. Combining 40 g N2O/MMBtu with a production 

of 17.5 million tonnes of oil equivalent in 2008 (British 

Petroleum, 2009; 1 tonne of oil equivalent = 42 GJ) we 

estimate an emission of 0.018 TgN yr−1 associated with 

production of corn ethanol in the USA. This is approximately 0.2% of the total anthropogenic emission of 6.7 

TgN yr−1 (IPCC, 2007). Crutzen et al. (2008) have argued 

that depending on N fertilizer uptake efficiency by plants, 

the formation of N2O during the production of commonly 

used automotive biofuels, such as biodiesel from rapeseed 
5.23

A Focus on Information for Policymakers

and bioethanol from corn (maize), can negate the CO2

mitigation from fossil fuel savings. However, the analysis 

by Crutzen et al. (2008) does not consider allocation of 

emissions to the coproducts in the process (e.g., animal 

feed produced during corn ethanol production).

The contribution of automotive biofuels to global 

N2O emissions is currently significantly below 1% and 

expected to remain so in the future. While global agricultural activities emit substantial amounts of N2O, biofuel 

production itself is currently not a major contributor to 

N2O emissions. Although current biofuel production represents a small fraction of total crude oil production, and 

increases in biofuel production appear likely, second generation biofuels (e.g., lignocellulosic ethanol) are anticipated to have substantially lower environmental impacts. 

Historical changes in the atmospheric concentrations 

of N2O and emissions sources are discussed in Section 

1.5.1.2 of Chapter 1.

5.4.2.4 Geoengineering: Enhancing Earth’s

Albedo by Stratospheric Injection

of Sulfur

There is increasing concern that efforts to reduce 

global greenhouse gas emissions will not be sufficient to 

prevent Earth’s climate from warming to levels that have 

never been experienced by modern societies. This has 

intensified the discussion about deliberate atmospheric 

modification schemes (“geoengineering”) to counteract 

enhanced greenhouse warming due to industrial emissions 

(for example, see the review panel report on geoengineering (Royal Society, 2009)). Although fraught with uncertainties and side effects that have yet to be studied in detail 

(Keith et al., 2010), the notion of enhancing Earth’s albedo by increasing particulate scattering of visible light in 

the stratosphere dates back many decades (e.g., Budyko, 

1977; Dyson and Marland, 1979). It has been estimated 

that injection of particles designed for optimum scattering of shortwave radiation (e.g., sulfate or aluminum oxide) might be economically viable and would likely be the 

most cost-effective method of a class of geoengineering 

options called “solar radiation management” (Teller et al., 

1997; Lenton et al., 2009; Royal Society, 2009). These 

studies also point out that the proposed solar radiation 

management techniques likely have the largest side effects of the geoengineering options considered. For example, changes in shortwave radiation would likely result 

in local temperature and precipitation changes, including 

increased prevalence of droughts (Hegerl and Solomon, 

2009). Various possible side effects of geoengineering 

(e.g., Robock, 2008) besides the impact on ozone are 

not further discussed in this Assessment. Other methods 

that have been suggested for managing solar radiation 

or taking up carbon dioxide (e.g., launching mirrors into 

space to harvest or block radiation; fertilizing the ocean to 

promote uptake of CO2 by phytoplankton) could also impact stratospheric ozone, but are not considered here due 

to a lack of scientific investigations that would allow for 

an assessment of the impacts of these approaches.

Part of the reason injection of sulfate into the stratosphere has received considerable attention in recent years 

from the scientific community as an option for solar radiation management (Crutzen, 2006) is due to observations 

of modest cooling of the globe following major volcanic 

eruptions, which may be viewed a natural analog to this 

particular approach to manage Earth’s radiation budget. 

The 1991 eruption of Mt. Pinatubo, for example, resulted in a decrease of global mean temperatures by 0.5°C 

for a few years (Soden et al., 2002). Observations from 

recent volcanic eruptions have shown that stratospheric 

ozone in middle and high latitudes is strongly influenced 

by the enhanced burden of liquid sulfate aerosols (e.g., 

Tabazadeh et al., 2002; Solomon et al.,1996; Portmann et 

al.,1996; Tilmes et al., 2008a). Recent model calculations 

have estimated a global temperature decrease of about one 

degree at Earth’s surface could be achieved within about 

five years after the start of a hypothetical sulfate aerosol 

injection experiment (Robock et al., 2008; Tilmes et al., 

2009). However, Heckendorn et al. (2009) have shown 

that the size distribution of aerosol particles is not fixed 

as assumed in earlier studies but strongly depends on 

different injection assumptions. The ability to achieve 

some amount of cooling that might be required to offset 

enhanced greenhouse-gas warming may be limited due 

to the coagulating and settling of aerosols. This process 

would also warm the tropical tropopause, perhaps resulting in enhanced abundances of stratospheric water vapor, 

which could accelerate ozone loss especially in the tropics 

and midlatitudes (Heckendorn et al., 2009). However, a 

significant increase of stratospheric water vapor was not 

observed after large volcanic eruptions.

Tilmes et al. (2008b, 2009) estimated the impact 

of an enhanced burden of stratospheric sulfur on the 

ozone layer during the period of time when stratospheric halogen loading is projected to slowly decline (Newman et al., 2007). Assuming an injection into the tropical stratosphere of a mass of volcanic-sized (submicron) 

sulfate aerosols large enough to counteract a doubling 

of CO2 with respect to preindustrial values (Rasch et al., 

2008a), ozone depletion in the Arctic polar vortex might 

be expected to double or triple (Tilmes et al., 2009). In 

addition, the expected recovery of the Antarctic ozone 

hole from the reduction in atmospheric halogen loading 

brought about by the Montreal Protocol might be delayed 

by between 30 and 70 years, depending on the assumed 

particle sizes (Tilmes et al., 2008b) (see Chapter 3). The 
5.24

Chapter 5

estimated impact on stratospheric ozone at low latitudes 

is expected to be smaller (about ±3%) for 2050 chlorine 

conditions (Tilmes et al., 2009).

Studies to quantify the impact of geoengineering 

on ozone have so far assumed basic scenarios and idealized conditions, and they have focused entirely on sulfate 

aerosols. Particle size distributions are usually described 

in simple terms, with large uncertainties in the evolution 

of those distributions (Rasch et al., 2008b). More comprehensive microphysical models, such as the one described 

in Heckendorn et al. (2009), are necessary to describe the 

complete size-distributions for geoengineering experiments. Dynamical and chemical changes in the stratosphere as a result of geoengineering are still uncertain 

given our current modeling capabilities (Chapter 3).

Uncertainties also exist in the amount of sulfur 

necessary to counteract global warming, which in turn depends on the amount of aerosol deposition (e.g., Heckendorn et al., 2009) and on the strength of the stratospheric 

circulation in the models, whether assumed for presentday conditions, or for a future accelerated stratospheric 

circulation (Rasch et al., 2008a). To date there have been 

no investigations of the impact on ozone caused by a gradual ramp-up of the amount of SO2 injected, with the purpose of keeping global average temperature nearly constant (Wigley, 2006). The amount of ozone that might be 

destroyed by deliberately increasing the aerosol loading of 

the stratosphere might also differ following a large volcanic eruption that further increases the amount of sulfur in 

the stratosphere (Tilmes et al., 2008b). Finally, there has 

been little study of an engineered stratospheric reflecting 

layer that uses materials other than sulfate (e.g., particles 

manufactured for maximum scattering of shortwave radiation and extended stratospheric lifetimes, as proposed by 

Teller et al., 1997).

These and other remaining gaps in our understanding of the full impacts of geoengineering on stratospheric 

ozone (and other aspects of the environment) are significant. The potential for significant risks to the ozone layer 

and the climate system, both known and unknown, from 

solar radiation management, as well as the fact that these 

approaches do not address other problems like ocean acidification as a result of increasing CO2 levels, have been 

acknowledged (Royal Society, 2009).

5.4.2.5 Emissions from Aviation and Rockets

Aviation

The importance of current aircraft emissions for 

ozone is well recognized (IPCC, 1999; WMO, 2003; 

Wuebbles et al., 2006; Maurice and Lee, 2009; Lee et 

al., 2009). Aircraft emissions, overall, are thought to 

increase globally averaged ozone columns by less than 

0.5%, primarily due to photochemical production of ozone 

catalyzed by nitrogen oxides (NOx), odd hydrogen (HOx) 

(from H2O), and hydrocarbon (HC) emissions. The production of ozone by emissions of NOx and HCs is offset 

by the destruction of ozone that occurs when emissions of 

aerosols and water vapor enhance abundances of ozonedestroying active halogens. As a consequence, little or no 

impact on stratospheric ozone column abundances is expected from aircraft emissions transported into the lower 

stratosphere from the troposphere (WMO, 2003; Brasseur, 

2008; Lee et al., 2009). New studies reinforce that there 

is a dual nature of aircraft emissions directly within the 

stratosphere or near the tropopause, with ozone production in some regions, and losses in others, and a strong dependence of the sign of net ozone changes on the latitude 

and altitude of those assumed emissions (Meilinger et al., 

2005; Søvde et al., 2007; Pitari et al., 2008; Köhler et al., 

2008; Cariolle et al., 2009).

Aviation-perturbed ozone can also affect the tropospheric oxidizing capacity, and thus levels of methane, 

a species that influences ozone chemistry in both the troposphere and stratosphere. Köhler et al. (2008) estimate 

that the increase in the oxidizing capacity of the troposphere due to aircraft emissions has reduced the lifetime 

of methane by 3.0%. Given that increases in methane 

are expected to increase stratospheric ozone (e.g., Section 5.4.2.2), this indirect change due to aircraft would 

be expected to have an offsetting effect on the increase 

in ozone columns.

Given that a large shift of worldwide aviation from 

subsonic to supersonic travel is unlikely in the foreseeable future, this Assessment assumes that aircraft emissions will be predominantly tropospheric. In this case, the 

net impact of those emissions will be ozone production, 

with some regional variation such as larger enhancements 

in heavily traveled flight corridors (e.g., Northern Hemisphere middle latitudes in summer). Ozone changes are 

expected to be smaller in the Southern Hemisphere, where 

there are fewer flights and emissions.

Rockets

A summary of the important processes that can 

contribute to ozone loss from rocket launches is found 

in Chapter 1. In the context of scientific and regulatory 

frameworks for protecting stratospheric ozone, the treatment of combustion emissions from rockets is somewhat 

complicated because ODP calculations involve some 

poorly understood processes, may be only regionally appropriate (not globally), and depend on many factors (e.g., 

propellant type, trajectory, latitude of launch site, etc.). 

In addition, the atmospheric lifetimes of rocket emissions (except for CO2) are less than a few years, which are 

short compared to the lifetimes of most halocarbon source 
5.25

A Focus on Information for Policymakers

gases. As is the case for aircraft, the impacts from rockets 

are due to primary and secondary products of combustion 

(e.g., H2O, NOx, hydrogen chloride (HCl), alumina, soot, 

and sulfate), and not to the fuels themselves. As with other processes presented in this chapter, the impact of rocket 

emissions on ozone may attract more attention in the coming decades if the space launch market grows (Ross et al., 

2009).

Several new classes of rocket propellant are of particular concern even though they do not contain chlorine, 

because they could still have significant effects on ozone. 

The first, composed of aluminum and H2O, has been proposed recently as an “environmentally friendly” solid 

rocket motor replacement. Because the emission from an 

Al/H2O rocket is mainly composed of alumina particles 

(compared to 30% for conventional solid rocket motors), 

ozone losses from such a rocket could exceed those from 

conventional solid rocket motors for equal emission rates 

if there is an ample reservoir of chlorine in the stratosphere, which will be the case for at least a few decades. 

The second, a “hybrid” that uses N2O and solid hydrocarbon as propellants, is being developed for widespread 

commercial spaceflight and could account for about onethird of all rocket emissions by 2020 (Seedhouse, 2008). 

Very little is known about the nature and magnitude of 

the emissions from this new hybrid rocket engine, though 

initial concerns for stratospheric ozone will focus on uncombusted N2O and soot.

Given these uncertainties and the long lead times in 

space launch hardware engineering and development, the 

lack of metrics for addressing the global impacts of rocket 

emissions can be particularly problematic (Ross et al., 

2009). Improved understanding of the processes involved 

in ozone depletion and development of suitable metrics 

can provide an important framework for launch industries 

to identify the factors responsible for the environmental 

impacts of rockets. This could allow, for example, for 

choice of propellant that best minimizes the environmental impacts.

5.4.2.6 Summary

The ozone impacts of the processes discussed in 

Section 5.4.2 are qualitatively compared in Table 5-5. Our 

understanding of the ozone impact of activities directly 

relating to production and emissions of halocarbons and 

CO2, CH4, and N2O is fairly advanced, while important 

gaps remain in our understanding of the impacts of other 

activities like aviation, space transport, and geoengineering. Chlorine- and bromine-containing ODSs play the 

largest role in current ozone depletion while the enhancement of CO2 and CH4 above background levels has likely 

led to higher globally averaged ozone levels than otherwise 

would have occurred. However, because of the controls 

already imposed by the Montreal Protocol on production 

and consumption of ODSs and the long atmospheric lifetimes of these compounds, the ability to further reduce 

future atmospheric ODS concentrations and thus ozone 

depletion through additional controls is more limited than 

before. Future ozone evolution is expected to be largely 

affected by activities and emissions not currently regulated 

by the Montreal Protocol, particularly by changes in CO2, 

CH4, and N2O. Some of these activities and emissions will 

likely increase ozone and some will decrease it. Our ability to project future stratospheric ozone is becoming more 

complicated, in that activities controlled by the Montreal 

Protocol are likely no longer the most important ones for 

the future evolution of stratospheric ozone.

5.4.3 Climate Impacts

5.4.3.1 Major HFCs Used as Replacements

for ODSs

Global production and use of CFCs and halons 

have decreased significantly because of the phase-outs under the Montreal Protocol and its subsequent Amendments 

and Adjustments. As a direct result, the use of HCFCs and 

HFCs as replacements has increased in Article 5 and nonArticle 5 countries (see Chapter 1). The HCFC increases 

occurred earlier and have been confirmed by long-term 

growth in observed atmospheric mixing ratios (Montzka et 

al., 2009; Stohl et al., 2009) (Chapter 1). Recent changes 

in northern-latitude mixing ratio observations are consistent with reduced use of HCFCs in non-Article 5 countries 

and increased use in Article 5 countries (Montzka et al., 

2009). HCFCs have been used as low-ODP substitutes 

in many applications for high-ODP substances and were 

classified under the Protocol as “transitional substitutes” 

to be used during the time it took to commercialize new 

ozone-safe alternatives and replacements. In 2007 the 

Parties to the Montreal Protocol decided to accelerate the 

HCFC phase-out, in part to protect future climate. HCFC 

production and consumption in Article 5 countries will 

be frozen in 2013 and stepwise reduced, with a virtually 

complete phase-out in 2030. Non-Article 5 countries have 

agreed to a virtually complete phase-out in 2020. In adopt- In adopting the accelerated HCFC phase-out, the Parties agreed to 

promote the use of HCFC alternatives that minimize the 

impact on climate.

With the global phase-out of HCFCs, much of the 

future application demand for refrigeration and air conditioning and heating is likely to met by HFCs, while the 

demand for thermal-insulating foam production is likely 

to be met by HFCs as well as hydrocarbons and not-inkind thermal insulation such as fiberglass and mineral 

wool (IPCC/TEAP, 2005; Velders et al., 2009). HFCs 
5.26

Chapter 5

Table 5-5. Qualitative summary of anthropogenic processes affecting stratospheric ozone.

Activity Mechanism

Effect on Ozone Layer

Ozone

Column

Increase or

Decrease a

Understanding b

Approximate

Maximum

Impact c

Destruction of CFC banks Ozone destruction by

chlorine chemistry

+ High Medium

Destruction of HCFC banks Ozone destruction by

chlorine chemistry

+ High Small

Destruction of halon banks Ozone destruction

by chlorine and

bromine chemistry

+ High Medium

Eliminate HCFC production Ozone destruction by

chlorine chemistry

+ High Medium

Eliminate CH3Br production Ozone destruction by

bromine chemistry

+ Medium Medium

Eliminate CCl4 production Ozone destruction by

chlorine chemistry

+ Medium-Low Medium

Reduction in N2O emissions d Ozone destruction by NOx + High Medium

Mitigation of ozone destruction

by chlorine/bromine chemistry

− Medium Small

Reduction in CH4 emissions d Ozone production − Medium Large

Ozone destruction by HOx + Medium Small

Mitigation of ozone destruction 

by chlorine chemistry

− Medium Small

Reduction in CO2 emissions d Middle- and upperstratospheric temperature 

changes

− High Large

Formation of PSCs + Medium Large

Circulation changes e + & − Low Large

Increased aviation emissions Aerosols: heterogeneous ozone 

destruction by halogens

− Medium Depends on # of 

flights, altitudes,

and locations

Production of ozone by 

HC/HOx/NOx chemistry

+ Medium

Contrail frequency ? Low

Activation of halogens − Low

Increased rocket / 

space-transport 

emissions

SRM f – ozone destruction by 

chlorine chemistry

− High Small at present 

launch rate

SRM f – aerosols: 

Heterogeneous ozone 

destruction by halogens

− Medium

Water vapor − Low

Soot ? Low

Ozone destruction by NOx − Low

Geoengineering by 

sulfur injection

Heterogeneous ozone 

destruction by halogens

− Medium-Low Potentially large

"

"

"

"

"

"

"
5.27

A Focus on Information for Policymakers

do not deplete the ozone layer but, along with CFCs and 

HCFCs, are greenhouse gases, which contribute to the radiative forcing of climate (IPCC/TEAP, 2005; Velders et 

al., 2009). Thus, the current transition away from CFCs 

in Article 5 countries and HCFCs in Article 5 and nonArticle 5 countries has implications not only for the ozone 

layer, but also for future climate. HFCs are not controlled 

by the Montreal Protocol, but they are included in the basket of gases of the Kyoto Protocol.

The HFC emissions scenarios discussed here can be 

divided in two classes:

• One class consists of non-intervention scenarios 

based on growth rates in gross domestic product and 

population (IPCC, 2000) for the period up to the middle or end of the 21st century. In these scenarios, 

historical and current technological developments 

are extrapolated without intervention from political 

processes. Such scenarios were first developed for 

the IPCC–Special Report on Emissions Scenarios 

(SRES) (IPCC, 2000) during the time when HFCs 

had still very limited usage and when the transition 

from CFCs and HCFCs had just begun. New nonintervention HFC scenarios have been formulated 

by Velders et al. (2009) and by the German Federal 

Environment Agency (UBA, 2009). The scenarios in 

Velders et al. (2009) are based on assumptions similar 

to those of IPCC-SRES with respect to growth rates 

in gross domestic product and population, but have 

incorporated new current information on (1) reported 

recent increases in consumption of HCFCs in Article 

5 countries of about 20% yr−1 (through 2007), (2) replacement patterns of HCFCs by HFCs as reported 

in non-Article 5 countries, and (3) accelerated phaseout schedules of HCFCs in Article 5 and non-Article 

5 countries (2007 Adjustment of the Montreal Protocol). The scenarios of UBA (2009) are based on 

more fundamental sector-specific growth rates and 

emissions factors for all relevant sectors in Article 5 

and non-Article 5 countries and take into account the 

accelerated HCFC phase-out schedules.

• The other class of scenarios is based on projections of 

technological developments affected by policy incentives and can therefore be seen as intervention scenarios. The discussions by the Parties of the Montreal 

Protocol to promote the use of HCFC alternatives that 

minimize the impact on climate can also be considered an intervention. An intervention scenario was 

developed for IPCC/TEAP (2005) and recently updated (TEAP, 2009), covering short time periods up to 

2015 and 2020, respectively. Others (Rao and Riahi, 

2006; Ottinger Schaefer et al., 2006; van Vuuren et al., 

2006) have reported HFC scenarios similar to those of 

SRES. Many of these latter scenarios have performed 

a relatively simple analysis to project future HFC 

emissions. Ottinger Schaefer et al. (2006) estimated 

HFC emissions through 2020 using a detailed sector 

analysis, but their analysis was performed before the 

2007 accelerated HCFC phase-out was adopted, and 

their 2008 HFC emissions do not agree with the emissions derived from observations.

The new scenarios constructed for IPCC, the Representative Concentration Pathways (RCPs), also contain 

HFC emissions, but use relatively simple or older analyses for the HFCs (Riahi and Nakicenovic, 2007; Clarke et 

al., 2007; Smith and Wigley, 2006; Wise et al., 2009; van 

Vuuren et al., 2006, 2007).

In Figure 5-5 global HFC emissions and radiative forcing are shown for six of the scenarios described 

above. The scenarios do not contain emissions from 

HFC-23, which is a by-product from the production of 

HCFC-22 and is not significantly used as a replacement 

for ODSs. HFC-23 emissions are estimated to be 0.20 ± 

0.03 GtCO2-eq yr−1 for 2006–2008 (Montzka et al., 2010), 

but are likely to decrease in the coming decade as a result 

of mitigation efforts for climate protection and because of 

projected decreases in HCFC-22 production. However, if 

mitigation is not successful, and unregulated production 

of HCFC-22 for use as a feedstock increases substantially, it is possible for HFC-23 emissions to increase in the 

future. The GWP-weighted HFC emissions are similar 

Table 5-5, continued (notes).

Notes:

a The “+” sign indicates more column ozone (i.e., less depletion) for the indicated policy option or activity; the “−” sign indicates less column ozone. b Understanding: “Low” indicates an inability to accurately quantify due to unknown or unquantifiable processes or emissions; “Medium” indicates 

remaining gaps in our understanding; “High” indicates that processes and values are believed to be well modeled and understood. c Approximate maximum impact: a relative scale that relates the various activities. This column is somewhat subjective and some of the entries depend 

strongly on the amount of future activity. Rocket emissions and aviation emissions have the potential to be important, given our current understanding, but specific information regarding flight altitude, number of flights, type of fuel, aerosol emissions, etc., needs to be specified before they can be 

usefully compared to the other ozone impacts. d The impacts of a reduction in N2O, CH4, and/or CO2 emissions will depend on the sizes of the reductions. Therefore, the “Approximate maximum 

impact” is somewhat more subjective than for the ODS phase-outs. e Circulation changes are expected to decrease (increase) total column ozone in the tropics (extratropics). See discussion in Chapter 3. f SRM stands for solid rocket motor.
5.28

Chapter 5

through about 2015 for all six scenarios. However, while 

the emissions of TEAP (2009) increase linearly through 

2020, the emissions of Velders et al. (2009) and UBA 

(2009) increase faster as a consequence of the growth 

rates in gross domestic product and the limits on consumption of HCFCs in Article 5 countries. By 2050, the 

total GWP-weighted HFC emissions are 5.5–8.8 GtCO2-

eq yr−1 in Velders et al. (2009) and 3.5 GtCO2-eq yr−1 in 

UBA (2009); both are significantly larger than those of 

SRES (1.3–2.3 GtCO2-eq yr−1) (IPCC, 2000). The HFC 

emissions in Velders et al. (2009) are about four times 

larger than those of SRES in 2050 for two principal reasons. First, the starting points (2008) for consumption of 

HFCs by Velders et al. (2009) are substantially higher than 

assumed in SRES, based on higher reported consumption 

of HCFCs from 2000 to 2007 (UNEP, 2009) than was 

assumed in SRES. Second, in the Velders et al. (2009) 

study, the HFCs assumed to meet future demand (HFC125 and HFC-143a) have larger GWPs than the HFCs 

assumed in SRES (mostly HFC-134a). This follows from 

assuming that the future use of HFCs in applications in 

developing countries will follow the same trends already 

observed in developed countries. The current consumption values for HFC-125 and HFC-143a are supported by 

their emissions as estimated from observed atmospheric 

mixing ratios and account for approximately 80% of the 

CO2-eq emission of HFCs in the year 2050. The lower 

emissions in the UBA (2009) scenarios compared with 

Velders et al. (2009) are the result of lower growth rates 

during 2020−2050 for the major HFC-consuming sectors 

and different replacement patterns for ODSs by HFCs and 

not-in-kind alternatives.

The current and expected future increased use of 

HFCs is largely the result of the phase-out of CFCs and 

HCFCs by the provisions of the Montreal Protocol, in 

combination with economic growth and an increase in 

living standards. As shown in Figure 5-6, total direct 

GWP-weighted emissions of ODSs peaked in 1988 at 

HFC Emissions

2000 2010 2020 2030 2040 2050

Year

0

2

4

6

8

10

Emissions (GtC

O2-eq yr

GWP-weighted (100-yr)

(excl. HFC-23)

Velders et al. (2009)

SRES (2000)

RCP (2009)

UBA (2009)

IPCC/TEAP(2005)

TEAP(2009)

HFC Radiative Forcing

2000 2010 2020 2030 2040 2050

Year

0.0

0.1

0.2

0.3

0.4

Radiative forcing (W 

m

Velders et al. (2009)

SRES (2000)

RCP (2009)

IPCC/TEAP(2005)

−1)

−2)

Figure 5-5. Global HFC emissions (left) and radiative forcing (right) in four long-term and two short-term scenarios. The long-term scenarios for the period 2000–2050 include the RCP scenarios and the non-intervention 

scenarios from Velders et al. (2009), UBA (2009), and SRES (IPCC, 2001). The short-term scenarios for the 

period 2000–2015/2020 include the IPCC-TEAP (2005) and TEAP (2009) scenarios. The RCP scenarios and 

the short-term scenarios assume intervention to mitigate climate forcing. Emissions of HFC-23, which is a 

by-product from the production of HCFC-22, are not included. HFC-23 emissions are estimated to be 0.20 ± 

0.03 GtCO2-eq yr−1 for 2006–2008 (Montzka et al., 2010). The emission values are multiplied by their GWPs 

(100-year time horizon) to obtain annual equivalent GtCO2 emissions. A GtCO2-eq equals 1000 million metric 

tons carbon dioxide equivalent (MMTCO2e). The color-shaded regions bound the upper and lower limits of the 

respective scenarios. Adapted from Velders et al. (2009).
5.29

A Focus on Information for Policymakers

9.4 GtCO2-eq yr−1 and decreased after that, whereas HFC 

emissions are projected to monotonically increase, primarily in Article 5 countries, exceeding those of ODSs 

after about 2020 in both Velders et al. (2009) and UBA 

(2009). In a business-as-usual scenario, starting in 1987, 

without Montreal Protocol regulations the GWP-weighted 

emissions of ODSs reach 15–18 GtCO2-eq yr−1 by 2010 

(Velders et al., 2007). So, growth in HFC use and emissions would offset at least part of the climate benefits calculated to have been achieved by the Montreal Protocol.

The HFC scenario results are further put into context by comparing to projected global CO2 emissions. 

Global annual HFC emissions in 2050 as projected by 

UBA (2009) and Velders et al. (2009) are equivalent 

(CO2-eq basis) to 6–8% and 9–19%, respectively, of pro- –19%, respectively, of pro- 19%, respectively, of projected global CO2 emissions in IPCC/SRES business-asusual scenarios (A1B, A2, B1, and B2) (Figure 5-6).

In these HFC scenarios, only the direct contribution to climate forcing due to ODS and HFC emissions 

is considered. Indirect climate forcings associated with 

0

1

2

3

4

Megatonnes per year

Mass-Weighted Emissions

Without Montreal Protocol

HFCs

ODSs

HCFCs

CFCs-only

Without accelerated

HCFC phaseout

high

low

0.0

0.5

1.0

1.5

2.0

Megatonnes CFC-11-eq per year

ODP-Weighted

Emissions

Without Montreal

Protocol

ODSs

HCFCs

CFCs, 

halons,

others

10

20

30

Gigatonnes CO2-eq per year

HFCs

ODSs

HCFCs

CFCs, halons, others

GWP-Weighted Emissions

Without Montreal Protocol

IPCC-SRES CO2 range

Magnitude of

Kyoto Protocol

reduction target

 2.0

high

low

1950 1970 1990 2010 2030 2050

Year

0

Figure 5-6. Emissions of ODSs (CFCs, halons, 

HCFCs, and others) and their non-ozone-depleting 

substitutes (HFCs) from 1950 to 2050 (adapted from 

Velders et al., 2007, 2009). Emissions are the total from developing and developed countries. The 

legends identify the specific compound classes included in each panel. The high and low HFC labels 

identify the upper and lower limits, respectively, in 

global baseline scenarios from Velders et al. (2009). 

The blue hatched regions indicate the emissions that 

would have occurred, in the absence of the Montreal 

Protocol, with 2–3% annual production increases 

in ODSs. Top panel: Global mass-weighted emissions expressed in megatonnes per year. The yellow 

dashed line shows HCFC emissions calculated without the provisions of the accelerated HCFC phaseout under the 2007 Adjustment of the Montreal Protocol. Middle panel: Global ODP-weighted emissions 

expressed in megatonnes of CFC-11-equivalent per 

year. The emissions of individual gases are multiplied 

by their respective ODPs (CFC-11 = 1) to obtain aggregate, equivalent CFC-11 emissions. The dashed 

line marks 1987, the year of the Montreal Protocol 

signing. Bottom panel: Global GWP-weighted emissions expressed in gigatonnes of CO2-equivalent per 

year. The emissions of individual gases are multiplied by their respective GWPs (direct, 100-year time 

horizon; CO2 = 1) to obtain aggregate, equivalent 

CO2 emissions. Shown for reference are emissions 

for the range of IPCC-SRES CO2 scenarios (IPCC, 

2000). The CO2 emissions for 1950–2007 are from 

global fossil fuel and cement production. Beyond 

2007, the shaded region for CO2 reflects the range 

bracketed by the A1B and B2 SRES scenarios. The 

dashed line denotes 2010, the middle year of the 

first commitment period of the Kyoto Protocol. Also 

shown is the magnitude of the reduction target of the 

first commitment period of the Kyoto Protocol, which 

is based on a 1990–2010 projection of global greenhouse gas emission increases and the reduction target for participating countries.
5.30

Chapter 5

halocarbon usage arise from the energy used or saved during the application or product lifetime and energy required 

to manufacture the product, including the halocarbon used. 

For example, thermal insulating products in buildings and 

appliances reduce energy consumption, and refrigeration 

and AC systems consume energy over their lifetimes. A 

full evaluation of the total climate forcing resulting from 

the global transition away from CFCs and HCFCs toward 

HFCs requires consideration of both direct and indirect 

impacts over all associated halocarbon, non-halocarbon, 

and not-in-kind application lifecycles.

Three primary options are available to reduce the 

future climate impact of HFCs as demand continues to 

increase for products currently using HCFCs and HFCs: 

first is an increasing implementation of not-in-kind technologies (non-fluorocarbon based); second, use of HFCs 

that result in less energy used, because many products 

have significantly higher indirect climate impacts due 

to the energy used over the lifetime of the product than 

they do from emissions of the chemicals contained in the 

product; and third, lower-GWP HFC or other replace third, lower-GWP HFC or other replacement chemicals could be used to satisfy demand. Figure 

5-7 shows steady-state radiative forcing resulting from a 

given annual emission rate for seven representative GWP 

limits on replacements. If the total 2050 HFC production projected in Velders et al. (2009) were to be met, for 

example, by HFCs with an average GWP less than 10 or 

100, and emissions are assumed to equal production, the 

steady-state globally averaged radiative forcing would be 

less than 0.003 or 0.03 W/m2, respectively. These values 

are substantially smaller than the 0.25–0.4 W/m2 range estimated in the unmitigated case in Velders et al. (2009). 

This analysis assumes that even very short-lived HFCs 

can be considered to be well mixed. In reality, the globalmean lifetimes and radiative efficiencies of substances 

that are not well mixed will depend on the location of 

emission. There is currently not sufficient discussion of 

impacts of such short-lived substances on radiative forcing 

and climate in the literature to assess the significance of 

their inhomogeneous abundances on this type of analysis.

5.4.3.2 Other Replacements for ODSs

Unsaturated HFCs (also known as hydrofluoroolefins, HFOs), hydrofluoroketones and perfluoroketones, 

hydrofluoroethers (HFEs), perfluoropolyethers, ammonia, 

hydrocarbons and CO2 are either under consideration or 

in use as replacements for ODSs. The potential for these 

compounds to contribute to radiative forcing of climate 

is considered in this section (other potential environmental impacts are discussed in Section 5.4.4). Most of these 

compounds are short-lived, which limits their radiative 

forcing. See Chapter 1 for additional discussion of these 

compounds, and Tables 1-10 and 1-11 for a list of alternatives for ozone-depleting substances.

The reaction of hydroxyl radicals (OH) with unsaturated HFCs proceeds rapidly, and as a consequence, 

these unsaturated HFCs have atmospheric lifetimes that 

are in the range of 1–20 days. With such short lifetimes 

the contribution of unsaturated HFCs to radiative forcing 

will be small. For example, HFC-1234yf (CF3CF=CH2) 

has a GWP of approximately 4 (Nielsen et al., 2007; 

Papadimitriou et al., 2008b). HFC emissions are projected to be dominated by HFC-125 and HFC-143a and 

reach 3.5–8.8 GtCO2-eq yr−1 in 2050 (see Section 5.4.3.1). 

While substitution by unsaturated HFCs would reduce the 

Steady-State Radiative Forcing

100 1000 10000

Annual Emission (106 kg/yr)

10-5

10-4

10-3

10-2

10-1

Radiative Forcing (W/m2

)

Range of 2050 HFC Radiative Forcing from Velders et al. [2009] Peak CFC Production (1988)

Annual Unmitigated HFC Production Projected in 2050 from Velders et al.

GWP=1000

300

100

30

10

3

1

Figure 5-7. Relationship between annual emission rates and 

steady-state radiative forcing for 

compounds having seven different GWPs. Relationships are valid 

for compounds having lifetimes 

less than 20 years or so. For example, if 1 Tg/yr (109

 kg/yr) were 

emitted of a gas with a GWP of 1, 

the steady-state radiative forcing 

would be slightly less than 10−4 W/

m2. For comparison, the 2008 estimated radiative forcing from CO2, 

CH4, and N2O was 1.74, 0.5, and 

0.17 W/m2, respectively (http://

www.esrl.noaa.gov/gmd/aggi/).
5.31

A Focus on Information for Policymakers

contribution of HFCs to radiative forcing, a full life cycle 

analysis (including emissions associated with the energy 

usage by appliances using the different compounds and 

including an analysis of the production of trifluoroacetic 

acid (TFA) and tropospheric ozone) would be needed to 

assess the total systemic environmental impact.

The atmospheric removal mechanism for HFEs is 

reaction with OH radicals. These rate coefficients decrease with increasing degree of fluorination of the ether. 

As seen from Appendix Table 5A-1, highly fluorinated 

HFEs (e.g., CHF2OCF3) can have substantial atmospheric 

lifetimes and large GWPs. Also, as evident from Appendix Table 5A-1 and discussed in Chapter 1, HFEs bearing several hydrogen atoms (e.g., C2H5OC4F9) have short 

atmospheric lifetimes and small GWPs. The GWPs for 

HFEs thus span a large range, and it is not possible to 

make any general statement regarding the benefit of the 

substitution of HFCs by HFEs. The benefits need to be 

assessed on a case-by-case basis. Perfluoropolyethers are 

not reactive toward OH radicals, or any other oxidant species in the atmosphere, have atmospheric lifetimes probably in excess of 800 years (Young et al., 2006), and will 

have large GWPs. The volatility and magnitude of possible emissions need to be considered when assessing the 

potential environmental impact of perfluoropolyethers. 

Finally, we note that the atmospheric oxidation products 

of unsaturated HFCs (HFOs), HFEs, hydrofluoroketones, 

perfluoroketones, and perfluoropolyethers are COF2, 

HC(O)F, HF, and fluorinated carboxylic acids (Calvert et 

al., 2008; D’Anna et al., 2005; Good and Francisco, 2003; 

Taniguchi et al., 2003; Wallington et al., 1994). These 

products have short atmospheric lifetimes and are not expected to contribute to radiative forcing.

As for the HFCs (Section 5.4.3.1), a full evaluation of the total environmental effects resulting from the 

global transition away from CFCs and HCFCs toward the 

compounds discussed in this section or other compounds 

requires consideration of both direct and indirect environmental impacts over all associated halocarbon, nonhalocarbon, and not-in-kind application lifecycles.

5.4.4 Other Environmental Impacts

The potential exists for certain HCFCs, HFCs, 

HFEs, and unsaturated HFCs (HFOs) to contribute to tropospheric ozone formation, and degrade to give toxic compounds. The atmospheric degradation of HCFCs, HFCs, 

HFEs, and unsaturated HFCs is initiated by reaction with 

OH radicals leading to the formation of halogenated carbonyl compounds, which undergo further oxidation to HF, 

HCl, CO2, and, in some cases, trifluoroacetic acid (TFA) 

(see for example, IPCC/TEAP 2005; Calvert et al., 2008; 

Hurley et al., 2008; Nielsen et al., 2007; Papadimitriou et 

al., 2008b).

Hayman and Derwent (1997) assessed the potential contribution of HCFCs and HFCs to the formation of 

tropospheric ozone and concluded that these compounds 

do not have a large potential to contribute to ground-level 

ozone formation. As indicated from the lifetimes of the 

compounds listed in Appendix Table 5A-1, the reactivity of HFEs toward OH radicals is comparable to those of 

analogous HCFCs and HFCs. As with HCFCs and HFCs, 

HFEs are not expected to make a significant contribution 

to ground-level ozone formation. OH radicals react rapidly with >C=C< double bonds, and unsaturated HFCs 

are typically much more reactive than HCFCs, saturated 

HFCs, and HFEs. The photochemical ozone creation potential (POCP) concept is a well-established method of 

ranking compounds by their ability to form ozone in the 

troposphere. POCP values are determined along an idealized straight-line trajectory using a photochemical trajectory model. Using the approach outlined by Derwent et al. 

(1998) and Jenkin (1998), Wallington et al. (2010) have 

estimated POCPs for representative unsaturated HFCs. 

The results are presented in Table 5-6 together with values 

for selected alkanes, alkenes, and HFCs. As seen from this 

table, the POCPs for unsaturated HFCs (HFOs) are much 

larger than those for longer-lived HFCs, much smaller than 

those for the parent alkenes, and for many compounds (including the commercially significant HFC-1234yf) lie between those of methane and ethane. Methane and ethane 

are oxidized sufficiently slowly that they do not contribute 

to any appreciable degree to local air quality issues and are 

generally exempt from air quality regulations. Luecken 

et al. (2010) conducted an atmospheric modeling study 

of the impact of replacing all HFC-134a currently used in 

vehicle air conditioning systems in the USA with HFC1234yf. They concluded that such large-scale use of HFC1234yf would result in a less than 0.01% increase in total 

ozone formed. Thus, this study provides further support 

for the expectation that unsaturated HFCs from mobile air 

conditioning will not make a significant contribution to 

tropospheric ozone formation in the near future.

The formation of hydrogen fluoride (HF) and hydrogen chloride (HCl) from the degradation of these 

source gases also should be addressed. Assuming that the 

combined global emissions of HCFCs, HFCs, HFEs, and 

unsaturated HFCs (HFOs) are of the order of 100 kilotonnes (Kt) per year, that they are uniformly distributed in 

the atmosphere, and an annual global precipitation of 4.9 

× 1017 liters (Erchel, 1975), the concentrations of HF and 

HCl in precipitation from degradation of HCFCs, HFCs, 

HFEs, and HFOs will be of the order of 3 × 10−9 molar. 

Although HF and HCl are strong acids, the concentration of fluoride and chloride and the additional acidity 

in precipitation resulting from the atmospheric oxidation 

of HCFCs, HFCs (saturated and unsaturated), and HFEs 

would be minor.
5.32

Chapter 5

TFA is a persistent, potentially toxic degradation 

product of some HCFCs (e.g., HCFC-123, -124), HFCs 

(e.g., HFC-134a, -227ea), and unsaturated HFCs (e.g., 

-1234yf, -1225ye (CF3CF=CHF)). 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). In WMO (2007) it 

was concluded that “TFA from the degradation of HCFCs 

and HFCs will not result in environmental concentrations 

capable of significant ecosystem damage.” The available 

data suggest that the same conclusion is applicable to unsaturated HFCs and HFEs. It has been shown that TFA is 

ubiquitous in precipitation and ocean water even in remote 

areas (Berg et al., 2000; Frank et al., 2002; Scott et al., 

2005; Scott et al., 2006; Von Sydow et al., 2000). Frank 

et al. (2002) estimated that the oceans contain 268 million 

tonnes of TFA. At a global level, the natural environmental loading of TFA greatly exceeds that expected from the 

atmospheric degradation of HCFCs, HFCs, and unsaturated HFCs (Kotamarthi et al., 1998). While Tromp et al. 

(1995) have argued that TFA will accumulate to high levels in seasonal wetlands, Boutonnet et al. (1999) showed 

that the assumptions made by Tromp et al. were highly improbable. Benesch et al. (2002) showed that TFA does not 

adversely affect the development of soil microbial communities and pool plant species in vernal ponds. Luecken 

et al. (2010) assessed the TFA concentrations following 

replacement of all HFC-134a currently used in vehicle air 

conditioning systems in the USA with HFC-1234yf. Their 

model predicted peak concentrations in rainfall of 1264 

nanograms per liter. This level is similar to peak concentrations currently observed (Scott et al., 2006) and is approximately two orders of magnitude lower than the level 

considered safe for the most sensitive aquatic organisms 

(Luecken et al., 2010). See also Section 1.3.6.3 of Chapter 

1 for a discussion on HFC-1234yf. Chapter 1 also discusses our current understanding of the sources of TFA. A 

current gap in our understanding is the lack of information 

concerning the natural sources and loss processes of TFA. 

This gap limits our ability to provide a precise assessment 

of the role of human activities in affecting local or regional 

abundances of TFA.

In summary, HCFCs, HFCs, HFEs, and HFOs are 

not expected to contribute significantly to the formation of 

tropospheric ozone. In the concentrations expected in the 

environment in the near future, the degradation products 

of HCFCs, HFCs (saturated and unsaturated), and HFEs 

are not expected to be toxic. However, for thorough understanding of the environmental impacts of replacements 

for CFCs, HCFCs, and long-lived HFCs, an evaluation of 

the ODPs, GWPs, atmospheric fate, safety, and toxicity is 

required for each replacement.

5.4.5 Impact of 2007 Accelerated 

HCFC Phase-Out

In 2007 the Parties to the Montreal Protocol adjusted the Protocol by accelerating the HCFC phase-out 

to protect the ozone layer and, in part, to further protect 

future climate. HCFC consumption in Article 5 countries 

will be frozen in 2013 and stepwise reduced, with a virtually complete phase-out in 2030. Before the 2007 Adjustment, a freeze in consumption was in place for Article 5 

countries starting in 2016 with a complete phase-out in 

2040. Non-Article 5 countries have agreed to a virtually 

Table 5-6. Photochemical ozone creation potentials (POCPs) for selected hydrofluoro-olefins 

and related alkanes, alkenes, and hydrofluorocarbons.

Compound Industrial or

Common Name POCP

CH2=CH2 ethylene; ethene 100 a

CH2=CF2 1,1 difluoroethylene;

HFO-1132a

18.0 b

CF2=CF2 Perfluoroethylene; 

HFO-1114

12.5 b

CH3CH=CH2 propylene; propene 117 c

CH2=CHCF3 HFO-1243zf 10.7 b

CH2=CFCF3 HFO-1234yf 7.0 b

(Z)-CHF=CFCF3 HFO-1225ye(Z) 5.6 b

CF2=CFCF3 Perfluoropropylene; 

HFO-1216

5.4 b

CH3CH2CH=CH2 1-butene 104 c

CH2=CHCF2CF3 HFO-1345czf 6.6 b

CH4 methane 0.6 d

CH2F2 HFC-32 0.2 e

C2H6 ethane 8 c

CH3CHF2 HFC-152a 1.0 e

CH3CF3 HFC-143a 0.0 e

CH2FCF3 HFC-134a 0.1 e

C3H8 propane 14 c

CH2FCHFCF3 HFC-245eb 0.2 b

CHF2CHFCF3 HFC-236ea 0.0 b

CF3CHFCF3 HFC-227ea 0.0 e

n-C4H10 butane 31 c

a By definition. b Wallington et al. (2010). c Derwent et al. (2007). d Derwent et al. (1998). e Hayman and Derwent (1997).
5.33

A Focus on Information for Policymakers

complete phase-out in 2020. The Adjustment also slightly 

changed the phase-out schedule for non-Article 5 countries, increasing the interim reduction during 2010 through 

2014 from 65% to 75% below the baseline level.

The accelerated HCFC phase-out has beneficial implications for the ozone layer and does have the potential 

to reduce climate forcing. It is expected to cause a reduction in cumulative HCFC emissions, relative to the current 

baseline scenario of this chapter, of 11−14 megatonnes 

(Mt) for the period 2011−2050, or an ODP-weighted 

equivalence of 0.6−0.8 MtCFC-11-eq (Figure 5-6). Following the same methods of calculating mixing ratios and 

EESC as described in Section 5.2, this emission reduction 

accelerates the projected return of EESC to 1980 levels by 

4−5 years. The resulting percentage reduction in EESC, 

integrated from 2011 until the 1980 level is reattained, is 

10−12%, which is larger than any of the hypothetical zeroproduction cases discussed in Section 5.4.3.1.

In adopting the accelerated HCFC phase-out, the 

Parties agreed to promote the use of HCFC alternatives 

that minimize the impact on climate and other environmental effects. The evaluation of the impact of a replacement of HCFCs with HFCs and other compounds requires 

consideration of both direct and indirect impacts, such 

as due to changes in energy usage associated with a particular replacement, over all associated halocarbon and 

not-in-kind application lifecycles. However, estimating 

all indirect impacts is extremely difficult. If only direct 

contributions are considered, the HCFC cumulative emissions reduction attributable to the accelerated phase-out 

is, in terms of climate forcing and relative to the current 

baseline, estimated to be 0.4−0.6 GtCO2-eq yr−1 averaged 

over 2011 through 2050. In comparison, global anthropogenic emissions of CO2 were greater than 30 Gt per year in 

2008. Not-in-kind alternative technologies (such as Stirling cycles, mineral wool, thermoacoustic refrigeration) 

are expected to only play a minor replacement role for the 

HCFCs. Therefore, to realize most of this forcing reduction, the avoided HCFC consumption and emissions must 

be replaced by alternative compounds with small GWPs 

(e.g., ammonia, hydrocarbons, CO2, low-GWP HFCs).

In addition to these potential benefits of the accelerated HCFC phase-out, there is also a climate benefit from 

the reduction in HFC-23 by-product emissions resulting 

from reduced HCFC-22 production. 

5.5 THE WORLD AVOIDED BY OZONE POLICY

The success of the Montreal Protocol and its subsequent Amendments and Adjustments (Andersen and 

Sarma, 2002; UNEP, 2009) is evident from the large 

decreases in the production of ODSs since their peak at the 

end of the 1980s, from the large decreases in emissions, and 

from decreases in midlatitude EESC since the middle of 

the 1990s that are expected to continue throughout the 21st

century. The success becomes even more striking when 

comparing current projections of production, emissions, 

and mixing ratios of ODSs and of the ozone layer with 

what might have happened without the Montreal Protocol. 

Scenarios generated to explore what might have occurred 

in the absence of ozone mitigation policies have been 

sometimes referred to as the “road not taken” or “world 

avoided” scenarios. The studies discussed in Section 5.5.3, 

made possible by the recent progress in chemistry-climate 

modeling (CCM), demonstrate the considerable damage 

mankind could have inflicted upon the Earth System in the 

space of just a few decades, and the benefit of the Montreal 

Protocol for both the ozone layer and climate.

ODSs affect the climate system both as greenhouse 

gases and indirectly through their effect on ozone. Both 

mechanisms are discussed below, starting with a discussion of ODS emission scenarios in a “World Avoided” by 

the Montreal Protocol.

5.5.1 ODS Production

An assessment of how ODS mitigation policies 

have affected climate and the ozone layer first requires 

establishing alternative scenarios of ODS production and 

emission. These scenarios are based on various reasonable assumptions about what might have happened in the 

absence of a policy response to scientific advances regarding the environmental impact of halocarbons, and about 

rates of emissions growth in such non-intervention scenarios. For example, two of the different scenarios that 

have been considered follow from the assumptions that (1) 

early warnings of dangers posed by CFCs to the ozone 

layer (Molina and Rowland, 1974) would have been ignored or would not have existed (referred to here as the 

MR74 scenario), and (2) that the Montreal Protocol was 

not implemented (NMP87 scenario) (Velders et al., 2007). 

Although negotiations about limiting the emissions of 

ODSs had started prior to the discovery of the ozone hole 

(Farman et al., 1985; Chubachi and Kajiwara, 1986), a 

failure to detect polar ozone depletion or to link it to CFCs 

would likely have led to the adoption of weaker controls, 

particularly in subsequent negotiations that tightened the 

initial measures following dissemination of the results 

attributing the Antarctic losses to halocarbons (Parson, 

2003). Prather et al. (1996) developed a scenario similar to MR74 assuming business-as-usual market-driven 

growth of CFC production after 1974. Under that scenario, although the ozone hole would have been detected, ozone science would have been delayed by around a 

decade. The mechanism of ozone depletion would have 

been established later, and emissions would have kept 

rising at 6% per year, with new uses for CFCs being introduced, until with a delay of 10 years ODS mitigation 
5.34

Chapter 5

policies would have been introduced. In this scenario, the 

tropospheric abundance of total chlorine would rise to 9 

parts per billion (ppb) by 2002 before mitigation policies 

took effect to reduce atmospheric abundances of chlorine. The Velders et al. (2007) MR74 scenario does not 

include mitigation. The second scenario from Velders et 

al. (2007), NMP87, is more conservative than MR74 and 

assumes CFC production would grow in proportion to the 

economy after 1987 (a characteristic of mature markets), 

with growth rates of 2−3%/year. Figure 5-8 summarizes the various World Avoided scenarios. Before 1987, 

NMP87 follows the observed chlorine abundance. Figure 

5-6 shows the global emissions of ODSs and HFCs from 

1950 to 2050 weighted by mass, ODPs, and GWPs for the 

baseline scenario and for the NMP87 scenario (Velders 

et al., 2007, 2009). Each of these scenarios results in a 

greatly enhanced stratospheric halogen loading compared 

to what has actually occurred.

5.5.2 Radiative Forcing

The radiative forcing (RF) potential of CFCs is well 

established, as was highlighted in WMO (1986). Velders 

et al. (2007) calculated the radiative forcing for the MR74 

and NMP87 scenarios introduced above, and for the actual 

and projected emissions of all ODSs (Table 5-7). For 

present-day abundances, the RF associated with halocarbons is around 0.3 W/m2

 or approximately 18% of that of 

CO2 (IPCC, 2007). Under the MR74 scenario, by 2010 the 

radiative forcing due to all ODS would contribute 44% to 

89% as much to global warming as CO2. Under NMP87, 

in 2010 the RF associated with ODSs would amount to 

about one-third of that of CO2. Even under this more conservative scenario, the RF increase due to ODSs between 

1988 and 2010 would be at least 0.34 W/m2, which is substantial compared to the increase in RF due to CO2 of 0.55 

W/m2 that has been experienced over that period.

The ODS forcing results should be adjusted for (a) 

the likely negative radiative forcings associated with ozone 

depletion caused by the ODS emissions (IPCC, 2007), and 

(b) the reduced emissions of replacement products that 

would not be needed without the mandated ODS reductions. 

HFCs are examples of replacement compounds; their usage 

is projected to grow sharply as a side effect of the Montreal 

Protocol (Figure 5-6), with potentially large increases in 

associated RF of up to 0.4 W/m2 in non-intervention scenarios by 2050 (Velders et al., 2009). Although the ODS 

emissions under a no-Montreal Protocol scenario (NMP87) 

would have a considerably larger radiative forcing, continuing unmitigated emissions of high-GWP HFCs could 

1960 1970 1980 1990 2000 2010 2020

Year

0.0

0.2

0.4

0.6

0.8

1.0

Mixing ratio (ppb)

Mixing ratio

CFC-11

CFC-12

MR74

NMP87

Baseline

1960 1980 2000 2020 2040 2060 2080 2100

Year

0

2

4

6

8

Tropospheric Chlorine (ppb)

Tropospheric Cl

MR74

NMP87

Baseline

Road-not-taken

Figure 5-8. Left panel: Total chlorine from the baseline scenario (black) compared with three “World Avoided” 

scenarios. The “Road-not-taken” scenario is from Prather et al. (1996) and includes only the portion of the “Free 

Market” curve (described in that study) before controls were implemented. The MR74 and NMP87 scenarios are 

derived from Velders et al. (2007). Right panel: Tropospheric mixing ratios of the most significant CFCs under 

the MR74 (green) (no identification of the link between CFCs and ozone in 1974 by Molina and Rowland (1974)) 

and the NMP87 (blue) (no Montreal Protocol restrictions) scenarios; and as observed and assuming compliance 

with the fully amended and adjusted Montreal Protocol (black). (From Velders et al., 2007).
5.35

A Focus on Information for Policymakers

nonetheless substantially offset the effects of other greenhouse gas mitigation efforts. See Section 5.4.2.1 for more 

details regarding future HFC emissions. The combined 

effects of ozone depletion and enhanced HFC emissions 

are estimated to offset about 30% of the radiative forcing 

avoided by the ODSs phased out under the Montreal Protocol by 2010 (Velders et al., 2007).

In comparing the benefits of the Montreal and Kyoto Protocols for climate, one also needs to assume that if 

the Montreal Protocol did not exist, the ODSs would almost certainly have been included in the Kyoto Protocol 

because of the ODSs’ considerable GWPs. However, this 

would have likely resulted in a delay in implementing controls similar to the one described by Prather et al. (1996), 

and could have led to unmitigated growth of CFCs lasting 

longer than what actually occurred.

Even allowing for these offsetting effects, Velders 

et al. (2007) note that actions taken under the Montreal 

Protocol have been of considerably greater benefit to climate than the reduction target of the first commitment 

period (2008−2012) of the Kyoto Protocol (Figure 5-6). 

The annual ODS emission reduction resulting from the 

Montreal Protocol is estimated to be 10 to 12 GtCO2-eq 

yr−1 by 2010, or 5 to 6 times the Kyoto Protocol reduction target. In making this comparison, it is acknowledged 

that the Kyoto Protocol does not include anthropogenic 

halocarbons because these substances had previously been 

regulated by the Montreal Protocol. Also unlike the Kyoto 

Protocol, the Montreal Protocol has gone through a series 

of Amendments and Adjustments that have led to further 

emission reductions. The Kyoto Protocol has not yet gone 

through this process. The first commitment period aims 

for an overall reduction, relative to 1990 levels, of 5.2% 

of CO2-equivalent greenhouse gas emissions in Annex 1 

countries. A proposed target for 2050 aims for 80% emission reductions, relative to 1990. Even compared to this 

much more ambitious target, the climate benefit of the 

Montreal Protocol calculated to date is significant.

5.5.3 Climate and Ozone Impacts Avoided 

by the Montreal Protocol

Prather et al. (1996) also estimated the consequences to ozone of unmitigated halocarbon emissions 

growth. Using a 2-D model, they calculated a 10% loss of 

extrapolar, annual-mean ozone by 1998, relative to 1980. 

More recently, Morgenstern et al. (2008) and Newman et 

al. (2009) assessed the impact of a continuing increase of 

chlorine in the atmosphere on both ozone depletion and 

climate. By using comprehensive CCMs, these are the 

first studies to examine the atmospheric structural and dynamical response to highly elevated chlorine. Both studies demonstrate the substantial changes that would likely 

have occurred if it were not for the Montreal Protocol.

Morgenstern et al. (2008) compare climate under a 

9 ppb total chlorine loading—the peak chlorine abundance 

projected in the Prather et al. (1996) scenario in about 

2002—to the actual peak chlorine loading of around 3.5 

ppb, which occurred in the late 1990s. In designing their 

calculations, they deliberately exclude the direct radiative 

effect of the CFC increases (the main topic of Velders et 

al., 2007) to focus exclusively on ozone changes and their 

climate impacts. The additional chlorine leads to a reduction of total ozone everywhere in their model (Morgenstern 

et al., 2009), with losses ranging from about 5% in the 

tropics to 20−30% in Arctic and over 50% in Antarctic 

spring; the average ozone loss between 60°S and 60°N in 

their model amounts to 8.3%. This is slightly less than 

estimated by Prather et al. (1996) for a somewhat lower 

chlorine abundance. The losses peak in three regions, 

namely in the chlorine layer around 40 km associated with 

gas-phase catalytic ozone destruction (Molina and Rowland, 1974), and at both high-latitude regions around 20 

km, associated with heterogeneous polar processes. The 

ozone loss has a considerable impact on stratospheric temperatures, with the upper stratosphere cooling by up to 6 K, 

and a few degrees cooling in the lower stratosphere with 

warming above it associated with a change in upwelling 

of 9.6 μm radiation due to the ozone loss. Morgenstern et 

al. (2008) also find substantial extratropical near-surface 

regional climate change. At southern high latitudes, they 

find a surface warming of up to 3 K in spring in the lee of 

the Antarctic Peninsula, and cooling in parts of Western 

and Eastern Antarctica (Figure 5-9); these calculations are 

qualitatively consistent with observations in recent decades 

(leading for example to the collapse of the Larsen B ice 

shelf; MacAyeal et al., 2003). Morgenstern et al. (2008) 

associate the climate change found in their model simulations with a further strengthening of the Southern Annular 

Table 5-7. Direct radiative forcing (W/m2) of controlled ODSs and CO2 (Velders et al., 2007; IPCC, 

2001; IPCC, 2007). For comparison: in 2008, CH4

and N2O were responsible for a direct radiative forcing of 0.50, and 0.17 W/m2

, respectively (www.esrl.

noaa.gov/gmd/aggi/).

Scenario 1975 1988 2010

ODS Baseline 0.12 0.26 0.32

ODS MR74 a 0.12 0.30–0.34 0.8–1.6

ODS NMP87 b 0.12 0.26 0.60–0.65

CO2 0.94 1.25 1.78–1.82

a Scenario in which the early warnings of dangers posed by CFCs to 

the ozone layer (Molina and Rowland, 1974) were ignored or did not 

exist.

b Scenario in which the Montreal Protocol was not implemented (“no 

Montreal Protocol”).
5.36

Chapter 5

Mode (SAM), the leading mode of variability of the southern extratropics, as has been observed during the period of 

historic ozone depletion (Thompson and Solomon, 2002; 

Fogt et al., 2009). For the Northern Hemisphere, their 

model produces substantial cooling in winter over Northern Asia and Western Europe, and warming over Northern 

North America and Greenland. Morgenstern et al. (2008) 

calculate a radiative forcing associated with the ozone depletion of –0.05 W/m2

, which, in absolute terms, is at least 

an order of magnitude less than the direct radiative forcing 

associated with the increased ODSs (Velders et al., 2007). 

However, they note that regional climate change due to 

the additional ozone depletion, over Antarctica and the 

northern continents, is large compared to recent observed 

climate changes.

In assessing these regional climate impacts, one 

needs to consider that the skill of climate models in representing regional climate trends on decadal timescales is 

often low, particularly in the Northern Hemisphere (e.g., 

Figure TS.24 of IPCC, 2007). In the Northern Hemisphere, 

atmosphere-ocean global circulation models (AOGCMs) 

typically disagree on basic patterns of climate change, both 

among each other and compared with observations (e.g., 

Gillett, 2005). At southern high latitudes, climate models 

are somewhat more successful in representing patterns of 

climate change (e.g., Gillett and Thompson, 2003). The 

projection by Morgenstern et al. (2008) appears plausible 

when comparing the patterns of the results to past climate 

changes in the Antarctic region. For the Northern Hemisphere, it is likely that ozone depletion would have led to 

substantial regional climate change, but the geographical 

pattern of change remains poorly understood.

Newman et al. (2009) perform and evaluate a 

transient simulation using the Goddard Earth Observing 

System Chemistry Climate Model (GEOSCCM) model 

(Pawson et al., 2008) in which EESC follows the MR74 

scenario of Velders et al. (2007) with an assumed growth 

rate of 3% per year, leading to 9 ppb in 2019. They terminate their simulation in 2065 when EESC reaches 45 ppb. 

Other greenhouse gases follow the IPCC (2000) A1B scenario. In their simulation, such drastic increases of ODSs 

completely alter the structure of the stratosphere. Total 

global ozone drops by two-thirds during the course of the 

simulation; the loss is approximately linear in EESC. The 

ozone loss produced by Newman et al. (2009) is similar to 

that of Prather et al. (1996) for similar chlorine loading, 

although a strict comparison is complicated by Newman et 

al. (2009) displaying ozone loss as a function of equivalent 

effective chlorine and Prather et al. (1996) using actual 

chlorine. The changes found by Newman et al. (2009) 

are initially largely driven by high-latitude ozone loss 

(Figure 5-10). In around 2020 in the simulation, Arctic 

ozone in April drops below 220 Dobson units (DU) and 

would thus satisfy a common definition of an Arctic ozone 

hole. By 2060, Arctic ozone in the simulation drops below 

100 DU in spring. There is a progressive deepening and 

lengthening of the lifetime of the Antarctic polar vortex in 

the simulation, eventually leading to a year-round ozone 

hole with westerlies in the southern mid- and high-latitude 

stratosphere. Newman et al. (2009) refer to this as the 

“perpetual winter.” At the end of the simulation, minimum 

ozone columns reach about 50 DU over Antarctica. After 

about 2045 in the simulation, enhanced upwelling cools 

the tropical lower stratosphere and triggers heterogeneous 

chemical processes similar to those currently observed in 

the Antarctic ozone hole. In the simulation, this leads to 

precipitous ozone depletion and nearly complete ozone 

loss in the tropical lower stratosphere by 2058. Newman 

et al. (2009) compare their results to complementary simulations assuming ODS abundances fixed at 1960 levels or 

following the A1 scenario (WMO, 2007) to demonstrate 

the effect of mitigating ODS emissions when the reference 

scenario has such dramatic chlorine increases.

Both Morgenstern et al. (2008) and Newman et al. 

(2009) exclude certain feedbacks that would have otherwise affected their results. In both studies, climate change 

due to ozone loss is not allowed to affect ocean conditions, 

which can have impacts that extend into the middle atmosphere (Kodama et al., 2007). In Newman et al. (2009), the 

direct RF due to the increasing ODSs (Velders et al., 2007) 

−3 −2 −1.5 −1 −0.5 0 0.5 1 1.5 2 3

SON

Temperature (°C)

Figure 5-9. Difference between the 9 and 3.5 ppb 

total chlorine cases in Antarctic spring (September–

November) temperature (°C) 20 m above the surface. (From Morgenstern et al., 2008).
5.37

A Focus on Information for Policymakers

is also not reflected in the ocean conditions; Morgenstern 

et al. (2008) exclude this effect by keeping the ODSs 

the same for radiation as in their reference simulation. 

Newman et al. (2009) prescribe tropospheric ozone and 

thus suppress feedback effects on tropospheric chemistry 

(e.g., tropospheric ozone production or loss caused by the 

increasing ultraviolet (UV) radiation). Morgenstern et al. 

(2008) use photolysis rates in the troposphere that do not 

reflect changes to the ozone column, thus also not allowing for such feedbacks. However, substantial decreases of 

tropospheric ozone do occur in Morgenstern et al. (2008) 

due to a decreased influx of stratospheric ozone. Further 

feedbacks not correctly treated in the simulations relate to a 

shortening of lifetimes of ODSs and long-lived greenhouse 

gases in the World Avoided simulations. Such shortening could occur from increased actinic fluxes and an acceleration of the Brewer-Dobson circulation, and would 

cause increasing emissions of ODSs to lead to less-thanproportional increases in their atmospheric abundances.

5.5.4 UV Impacts of the Avoided 

Ozone Depletion

The decreases in ozone abundances in the World 

Avoided simulations (Prather et al., 1996; Morgenstern 

et al., 2008; Newman et al., 2009) would have led to an 

increase in solar ultraviolet radiation at Earth’s surface. 

The magnitude of these increases varies with wavelength, 

depending on the absorption cross section of ozone, and 

with latitude, according to the latitudinal distribution 

of ozone depletion. Newman et al. (2009) calculate the 

spectral changes in UV occurring in their simulation. At 

the shortest UV-B wavelengths, in relative terms, the 

calculated changes are extremely large; however, in absolute terms the surface irradiance at these wavelengths 

is weak. Regarding biologically relevant radiation, they 

show that decreases in stratospheric ozone due to increasing CFCs would have led to marked increases in erythemally weighted irradiance (which is proportional to the 

UV Index; Chapter 2), up to three times more in the summer of northern midlatitudes by 2065 (Figure 5-11), and 

with UV Index values in excess of 30 at midlatitudes of 

both hemispheres. The excess radiation would have had 

large impacts on the biosphere and on human health; for 

example, Northern Hemisphere midlatitude ozone losses 

would have led to a reduction of the sunburn time, at local 

noon for clear-sky midsummer conditions, from 15 to 5 

minutes, and would increase DNA-damaging UV by approximately 550% between 1980 and 2065 (Newman et 

al., 2009).

5.5.5 Summary

“Road not taken” or “World Avoided” studies indicate that the Montreal Protocol has had considerable 

benefits for both the ozone layer and climate change. 

Total chlorine would have increased much beyond the 

observed peak of around 3.5 ppb. Consequences of uncontrolled growth in ODSs would have resulted both directly from the radiative forcing of the halocarbons, and 

indirectly from the ensuing ozone depletion. In 2010, the 

-90 -60 -30 0 30 60 90

Latitude

0

100

200

300

400

500

Total Ozone (DU)

2065

2040

2020

2000

fixed Cl (1970)

baseline (1970)

baseline (2000)

2052

WORLD AVOIDED

Figure 5-10. Zonal-mean 

total ozone in the World 

Avoided simulation (black), 

in 1970 for chlorine fixed at 

1960 levels (green), and for 

a baseline scenario simulation in 1970 (blue) and 2000 

(red). (Adapted from Newman et al., 2009).
5.38

Chapter 5

RF associated with unmitigated emissions of halocarbons 

could have reached 2 to 5 times what has actually occurred. 

The chemical consequences could have included a thinned 

ozone layer everywhere on the globe, the occurrence of an 

Arctic ozone hole, regional climate change in middle and 

high latitudes of both hemispheres, increased UV with associated biological and medical consequences, and changes in stratospheric circulation including, in an extreme 

case, year-round westerlies in the Southern-Hemisphere 

stratosphere (the “perpetual winter”). While many details 

associated with these calculated World Avoided impacts 

may have large uncertainties associated with them, there 

is no doubt that the Montreal Protocol and its Amendments and Adjustments have avoided additional profound 

changes to stratospheric ozone and climate.

REFERENCES

Andersen, S.O., and K.M. Sarma, Protecting the Ozone 

Layer — the Official UN History, Earthscan, ISBN 

1-85383-905-1, London, 2002.

Barry, J., G. Locke, D. Scollard, H. Sidebottom, J. 

Treacy, C. Clerbaux, R. Colin, and J. Franklin, 

1,1,1,3,3,-Pentafluorobutane (HFC-365mfc): Atmospheric degradation and contribution to radiative 

forcing, Int. J. Chem. Kinet., 29 (8), 607-617, doi: 

10.1002/(SICI)1097-4601(1997)29:8<607::AIDKIN6>3.0.CO;2-Y, 1997.

Benesch, J.A., M.S. Gustin, G.R. Cramer, and T.M. Cahill, 

Investigation of effects of trifluoroacetate on vernal 

pool ecosystems, Environ. Toxicol. Chem., 21 (3), 

640-647, doi: 10.1002/etc.5620210325, 2002.

Berg, M., S.R. Müller, J. Mühlemann, A. Wiedmer, and 

R.P. Schwarzenbach, Concentrations and mass 

fluxes of chloroacetic acids and trifluoroacetic 

acid in rain and natural waters in Switzerland, Environ. Sci. Technol., 34, 2675-2683, doi: 10.1021/

es990855f, 2000.

Berntsen, T.K., J.S. Fuglestvedt, M.M. Joshi, K.P. Shine, 

N. Stuber, M. Ponater, R. Sausen, D.A. Hauglustaine, and L. Li, Response of climate to regional emissions of ozone precursors: Sensitivities and 

warming potentials, Tellus, 57B (4), 283-304, 2005.

Blowers, P., and K. Hollingshead, Estimations of global 

warming potentials from computational chemistry 

calculations for CH2F2 and other fluorinated methyl 

species verified by comparison to experiment, J. 

Phys. Chem. A, 113 (20), 5942-5950, doi: 10.1021/

jp8114918, 2009.

Boutonnet, J.C., P. Bingham, D. Calamari, C. de Rooij, 

J. Franklin, T. Kawano, J.-M. Libre, A. McCulloch, G. Malinverno, J.M. Odom, G.M. Rusch, K. 

Smythe, I. Sobolev, R. Thompson, and J.M. Tiedje, 

Environmental risk assessment of trifluoroacetic 

acid, Hum. Ecol. Risk Assess., 5 (1), 59-124, 1999.

Brasseur, G.P., A Report on the Way Forward Based on the 

Review of Research Gaps and Priorities, Aviation 

Climate Change Research Initiative, Federal Aviation Administration, 53 pp., available: http://www.

faa.gov/about/office_org/headquarters_offices/aep/

aviation_climate/media/ACCRI_Report_final.pdf, 

Moderate

1980 2000 2020 2040 2060

5

10

15

20

25

30

35

UV Index

fixed Cl

WORLD AVOIDED

baseline

High

Very High

Extreme

Midlatitudes (30°–50°N)

2 July, Noon Conditions

Year

Figure 5-11. Erythemally weighted UV radiation (expressed in 

terms of the UV Index) for noon 

conditions on 2 July at northern 

midlatitudes. Black: World Avoided. Red: A1 baseline scenario 

(WMO, 2007) for ODSs. Green: 

Cl fixed at 1960 levels. Gray: 

1975–1985 average of the fixedCl simulation. The conventional 

color scale used in reporting the 

UV Index to the public is superimposed for reference showing the 

highest four divisions. (Adapted 

from Newman et al., 2009).
5.39

A Focus on Information for Policymakers

2008.

British Petroleum, BP Statistical Review of World Energy June 2009, Beacon Press, available: http://

www.bp.com/liveassets/bp_internet/globalbp/

globalbp_uk_english/reports_and_publications/

statistical_energy_review_2008/STAGING/local_assets/2009_downloads/statistical_review_of_

world_energy_full_report_2009.pdf, 2009.

Budyko, M.I., Present-day climatic changes, Tellus, 29 

(3), 193-204, doi: 10.1111/j.2153-3490.1977.

tb00725.x, 1977.

Calvert, J.G., R.G. Derwent, J.J. Orlando, G.S. Tyndall, 

and T.J. Wallington, Mechanisms of Atmospheric 

Oxidation of the Alkanes, 1008 pp., Oxford University Press, New York, NY, U.S.A., 2008.

Cariolle D., D. Caro, R. Paoli, D.A. Hauglustaine, 

B. Cuénot, A. Cozic, and R. Paugam, Parameterization of plume chemistry into large-scale 

atmospheric models: Application to aircraft NOx

emissions, J. Geophys. Res., 114, D19302, doi: 

10.1029/2009JD011873, 2009.

Cavalli, F., M. Glasius, J. Hjorth, B. Rindone, and N.R. 

Jensen, Atmospheric lifetimes, infrared spectra and 

degradation products of a series of hydrofluoroethers, Atmos. Environ., 32 (21), 3767-3773, doi: 

10.1016/S1352-2310(98)00106-X, 1998.

Chipperfield, M.P., and W. Feng, Comment on: Stratospheric ozone depletion at northern mid-latitudes 

in the 21st century: The importance of future concentrations of greenhouse gases nitrous oxide and 

methane, Geophys. Res. Lett., 30 (7), 1389, doi: 

10.1029/2002GL016353, 2003.

Christidis, N., M.D. Hurley, S. Pinnock, K.P. Shine, 

and T.J. Wallington, Radiative forcing of climate 

change by CFC-11 and possible CFC-replacements, J. Geophys. Res., 102 (D16), 19597-19609, 

doi: 10.1029/97JD01137, 1997.

Chubachi, S., and R. Kajiwara, Total ozone variations at 

Syowa, Antarctica, Geophys. Res. Lett., 13 (12), 

1197-1198, doi: 10.1029/GL013i012p01197, 1986.

Clarke, L.E., J.A. Edmonds, H.D. Jacoby, H.M. Pitcher, 

J.M. Reilly, and R.G. Richels, Scenarios of Greenhouse Gas Emissions and Atmospheric Concentrations, Sub-report 2.1a of Synthesis and Assessment 

Product 2.1 by the U.S. Climate Change Science 

Program and the Subcommittee on Global Change 

Research, Department of Energy, Office of Biological & Environmental Research, Washington, D.C., 

U.S., 154 pp., 2007.

Crutzen, P.J., Albedo enhancement by stratospheric sulfur injections: A contribution to resolve a policy 

dilemma?, Clim. Change, 77 (3-4), 211-220, doi: 

10.1007/s10584-006-9101-y, 2006.

Crutzen, P.J., A.R. Mosier, K.A. Smith, and W. Winiwarter, 

N2O release from agro-biofuel production negates 

global warming reduction by replacing fossil fuels, 

Atmos. Chem. Phys., 8 (2), 389-395, doi: 10.5194/

acp-8-389-2008, 2008.

Cunnold, D.M., R.F. Weiss, R.G. Prinn, D. Hartley, P.G. 

Simmonds, P.J. Fraser, B. Miller, F.N. Alyea, and 

L. Porter, GAGE/AGAGE measurements indicating reductions in global emissions of CCl3F and 

CCl2F2 in 1992-1994, J. Geophys. Res., 102 (D1), 

1259-1269, 1997.

Daniel, J.S, and G.J.M. Velders (Lead Authors), A.R. 

Douglass, P.M.D. Forster, D.A. Hauglustaine, 

I.S.A. Isaksen, L.J.M. Kuijpers, A. McCulloch, and 

T.J. Wallington, Halocarbon scenarios, ozone depletion potentials, and global warming potentials, 

Chapter 8 in Scientific Assessment of Ozone Depletion: 2006, Global Ozone Research and Monitoring Project—Report No. 50, World Meteorological 

Organization, Geneva, Switzerland, 2007.

Daniel, J.S., S. Solomon, and D.L. Albritton, On the evaluation of halocarbon radiative forcing and global 

warming potentials, J. Geophys. Res., 100 (D1), 

1271-1285, 1995.

Daniel, J.S., E.L. Fleming, R.W. Portmann, G.J.M. 

Velders, C.H. Jackman, and A.R. Ravishankara, 

Options to accelerate ozone recovery: Ozone and 

climate benefits, Atmos. Chem. Phys., 10 (16), 

7697-7707, doi: 10.5194/acp-10-7697-2010, 2010.

D’Anna, B., S.R. Sellevåg, K. Wirtz, and C.J. Nielsen, 

Photolysis study of perfluoro-2-methyl-3-pentanone under natural sunlight conditions, Environ. 

Sci. Technol., 39 (22), 8708-8711, doi: 10.1021/

es048088u, 2005.

Derwent, R.G., M.E. Jenkin, S.M. Saunders, and M.J. 

Pilling, Photochemical ozone creation potentials 

for organic compounds in Northwest Europe calculated with a master chemical mechanism, Atmos. 

Environ., 32 (14-15), 2429-2441, 1998.

Derwent, R.G., M.E. Jenkin, N.R. Passant, and M.J. Pilling, Reactivity-based strategies for photochemical 

ozone control in Europe, Environ. Sci. Policy, 10

(5), 445-453, doi: 10.1016/j.envsci.2007.01.005, 

2007.

Dillon, T.J., A. Horowitz, and J.N. Crowley, The atmospheric chemistry of sulphuryl flouride, SO2F2, Atmos. Chem. Phys., 8 (6), 1547-1557, doi: 10.5194/

acp-8-1547-2008, 2008.

Douglass, A.R., R.S. Stolarski, M.R. Schoeberl, C.H. 

Jackman, M.L. Gupta, P.A. Newman, J.E. Nielsen, 

and E.L. Fleming, Relationship of loss, mean age 

of air and the distribution of CFCs to stratospheric circulation and implications for atmospheric 

lifetimes, J. Geophys. Res., 113, D14309, doi: 

10.1029/2007JD009575, 2008.
5.40

Chapter 5

Dyson, F.J., and G. Marland, Technical fixes for the climate effects of CO2, in Workshop on the Global 

Effects of Carbon Dioxide from Fossil Fuels, edited by W.P. Elliott and L. Machta, CONF-770385, 

U.S. Department of Energy, Washington, D.C., 

U.S., pp. 111-118, 1979.

Erchel, E., World Water Balance, Elsevier, New York, 

NY, U.S.A., 1975.

Farman, J.C., B.G. Gardiner, and J.D. Shanklin, Large 

losses of total ozone in Antarctica reveal seasonal 

ClOx/NOx interaction, Nature, 315, 207-210, doi: 

10.1038/315207a0, 1985.

Fisher, D.A., C.H. Hales, D.L. Filkin, M.K.W. Ko, N.D. 

Sze, P.S. Connell, D.J. Wuebbles, I.S.A. Isaksen, 

and F. Stordal, Model calculations of the relative effects of CFCs and their replacements on 

stratospheric ozone, Nature, 344, 508-512, doi: 

10.1038/344508a0, 1990.

Fleming, E.L., C.H. Jackman, D.K. Weisenstein, and 

M.K.W. Ko, The impact of interannual variability 

on multidecadal total ozone simulations, J. Geophys. 

Res., 112, D10310, doi: 10.1029/2006JD007953, 

2007.

Fogt, R.L., J. Perlwitz, S. Pawson, and M.A. Olsen, Inter-annual relationships between polar ozone and 

the SAM, Geophys. Res. Lett., 36, L04707, doi: 

10.1029/2008GL036627, 2009.

Forster, P.M. de F., J.B. Burkholder, C. Clerbaux, P.F. 

Coheur, M. Dutta, L.K. Gohar, M.D. Hurley, G. 

Myhre, R.W. Portmann, A. Ravishankara, K.P. 

Shine, T.J. Wallington, and D. Wuebbles, Resolution of the uncertainties in the radiative forcing of 

HFC-134a, J. Quant. Spectrosc. Radiat. Transfer, 

93 (4), 447-460, doi: 10.1016/j.jqsrt.2004.08.038, 

2005.

Frank, H., E.H. Christoph, O. Holm-Hansen, and J.L. Bullister, Trifluoroacetate in ocean waters, Environ. 

Sci. Technol., 36, 12-15, doi: 10.1021/es0101532, 

2002.

Fuglestvedt, J.S., K.P. Shine, T. Berntsen, J. Cook, D.S. 

Lee, A. Stenke, R.B. Skeie, G.J.M. Velders, and 

I.A. Waitz, Transport impacts on atmosphere and 

climate: Metrics, Atmos. Environ., 44 (37), 4648-

4677, doi: 10.1016/j.atmosenv.2009.04.044, 2010.

Gillett, N.P., Northern Hemisphere circulation, Nature, 

437 (7058), 496, doi: 10.1038/437496a, 2005.

Gillett, N.P., and D.W.J. Thompson, Simulation of recent 

Southern Hemisphere climate change, Science, 302

(5643), 273-275, 2003.

Good, D.A., and J.S. Francisco, Atmospheric chemistry 

of alternative fuels and alternative chlorofluorocarbons, Chem. Rev., 103 (12), 4999-5024, doi: 

10.1021/cr020654l, 2003.

Haigh, J.D., and J.A. Pyle, A two-dimensional 

calculation including atmospheric carbon dioxide 

and stratospheric ozone, Nature, 279, 222-224, doi: 

10.1038/279222a0, 1979.

Hayman, G.D., and R.G. Derwent, Atmospheric chemical 

reactivity and ozone-forming potentials of potential 

CFC replacements, Environ. Sci. Technol., 31 (2), 

327-336, 1997.

Heckendorn, P., D. Weisenstein, S. Fueglistaler, B.P. Luo, 

E. Rozanov, M. Schraner, L.W. Thomason, and T. 

Peter, The impact of geoengineering aerosols on 

stratospheric temperature and ozone, Env. Res. Lett.,

4 (4), doi: 10.1088/1748-9326/4/4/045108, 2009.

Hegerl, G.C., and S. Solomon, Risks of climate engineering, Science, 325 (5943), 955-956, doi: 10.1126/

science.1178530, 2009.

Hsu, J., and M.J. Prather, Global long-lived chemical modes excited in a 3-D chemistry transport 

model: Stratospheric N2O, NOy, O3, and CH4

chemistry, Geophys. Res. Lett., 37, L07805, doi: 

10.1029/2009GL042243, 2010.

Hurley, M.D., T.J. Wallington, M.S. Javadi, and O.J. 

Nielsen, Atmospheric chemistry of CF3CF=CH2: 

Products and mechanisms of Cl atom and OH radical initiated oxidation, Chem. Phys. Lett., 450, 263-

267, doi: 10.1016/j.cplett.2007.11.051, 2008.

Inoue, Y., M. Kawasaki, T.J. Wallington, and M.D. Hurley, Atmospheric chemistry of CF3CH2CF2CH3

(HFC-365mfc): Kinetics and mechanism of chlorine atom initiated oxidation, infrared spectrum, and 

global warming potential, Chem. Phys. Lett., 462, 

164-168, doi: 10.1016/j.cplett.2008.07.054, 2008.

IPCC (Intergovernmental Panel on Climate Change), Climate Change: The IPCC Scientific Assessment, 

edited by J.T. Houghton, G.J. Jenkins, and J.J. 

Ephraums, 364 pp., Cambridge University Press, 

Cambridge, U.K., 1990.

IPCC (Intergovernmental Panel on Climate Change), 

Climate Change, 1994: Radiative Forcing of Climate Change and an Evaluation of the IPCC IS92 

Emission Scenarios, edited by J.T. Houghton, L.G. 

Meira Filho, J. Bruce, H. Lee, B.A. Callander, E. 

Haites, N. Harris, and K. Maskell, 339 pp., Cambridge University Press, Cambridge, U.K., 1995.

IPCC (Intergovernmental Panel on Climate Change), 

Climate Change 1995: The Science of Climate 

Change, edited by J.T. Houghton, L.G. Meira Filho, B.A. Callander, N. Harris, A. Kattenberg, and 

K. Maskell, 572 pp., Cambridge University Press, 

Cambridge, U.K., 1996.

IPCC (Intergovernmental Panel on Climate Change), 

Aviation and the Global Atmosphere, A Special 

Report of the Intergovernmental Panel on Climate Change, edited by J.E. Penner, D. Lister, D.J. 

Griggs, D.J. Dokken, and M. McFarland, 373 pp., 
5.41

A Focus on Information for Policymakers

Cambridge University Press, Cambridge, U.K., doi: 

10.2277/0521664047, 1999.

IPCC (Intergovernmental Panel on Climate Change), 

Special Report on Emissions Scenarios, Special 

Report of Working Group III of the Intergovernmental Panel on Climate Change, edited by N. 

Nakicenovic and R. Swart, 570 pp., Cambridge 

University Press, Cambridge, U.K. and New York, 

NY, U.S.A., 2000.

IPCC (Intergovernmental Panel on Climate Change), Climate Change 2001: The Scientific Basis, Contribution of Working Group I to the Third Assessment 

Report of the Intergovernmental Panel on Climate 

Change, edited by J.T. Houghton, Y. Ding, D.J. 

Griggs, M. Noguer, P.J. van der Linden, X. Dai, 

K. Maskell, and C.A. Johnson, 881 pp., Cambridge 

University Press, Cambridge, U.K., 2001.

IPCC (Intergovernmental Panel on Climate Change), Climate Change 2007: The Physical Science Basis, 

Contribution of Working Group I to the Fourth Assessment Report of the Intergovernmental Panel on 

Climate Change, edited by S. Solomon, D. Qin, M. 

Manning, Z. Chen, M. Marquis, K.B. Averyt, M. 

Tignor, and H.L. Miller, 996 pp., Cambridge University Press, Cambridge, U.K., and New York, 

NY, U.S.A., 2007.

IPCC (Intergovernmental Panel on Climate Change), Summary Report of the IPCC Expert Meeting on the 

Science of Alternative Metrics, 18–20 March 2009, 

Oslo, Norway. IPCC-XXX/Doc.13 (31.III.2009), 

available: www.ipcc.ch/meetings/session30/doc13.

pdf, 2009.

IPCC/TEAP (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, prepared by Working Groups I 

and III of the Intergovernmental Panel on Climate 

Change, and the Technical and Economic Assessment Panel, Cambridge University Press, Cambridge, U.K., and New York, NY, U.S.A., 2005.

Jenkin, M.E., Photochemical ozone and PAN creation 

potentials: Rationalisation and methods of estimation, AEA Technology plc, Report AEAT4182/20150/003, 1998.

Johansson, D.J.A., U.M. Persson, and C. Azar, Uncertainty and learning: Implications for the trade-off 

between short-lived and long-lived greenhouse gases, Clim. Change, 88 (3-4), 293-308, doi: 10.1007/

s10584-007-9381-x, 2008.

Johnson, C.E., and R.G. Derwent, Relative radiative forcing consequences of global emissions of hydrocarbons, carbon monoxide and NOx from human 

activities estimated with a zonally-averaged twodimensional model, Clim. Change, 34 (3-4), 439-

462, doi: 10.1007/BF00139301, 1996.

Keith, D.W., E. Parson, and M.G. Morgan, Research on 

global sun block needed now, Nature, 463, 426-

427, doi: 10.1038/463426a, 2010.

Kodama, C., T. Iwasaki, K. Shibata, and S. Yukimoto, 

Changes in the stratospheric mean meridional circulation due to increased CO2: Radiation- and sea 

surface temperature-induced effects, J. Geophys. 

Res., 112, D16103, doi: 10.1029/2006JD008219, 

2007.

Köhler, M.O., G. Rädel, O. Dessens, K.P. Shine, H.L. 

Rogers, O. Wild, and J.A. Pyle, Impact of perturbations to nitrogen oxide emissions from global 

aviation, J. Geophys. Res., 113, D11305, doi: 

10.1029/2007JD009140, 2008.

Kotamarthi, V.R., J.M. Rodriguez, M.K.W. Ko, T.K. 

Tromp, N.D. Sze, and M.J. Prather, Trifluoroacetic acid from the degradation of HCFCs and HFCs: 

A three-dimensional modeling study, J. Geophys. 

Res., 103 (D5), 5747-5758, 1998.

Lee, D.S., G. Pitari, V. Grewe, K. Gierens, J.E. Penner, 

A. Petzold, M.J. Prather, U. Schumann, A. Bais, 

T. Berntsen, D. Iachetti, L.L. Lim, and R. Sausen, 

Transport impacts on atmosphere and climate: 

Aviation, Atmos. Environ., 44 (37), doi: 10.1016/j.

atmosenv.2009.06.005, 2009.

Lenton, T.M., and N.E. Vaughan, The radiative forcing 

potential of different climate geoengineering options, Atmos. Chem. Phys., 9 (15), 5539-5561, doi: 

10.5194/acp-9-5539-2009, 2009.

Li, F., R.S. Stolarski, and P.A. Newman, Stratospheric 

ozone in the post-CFC era, Atmos. Chem. Phys., 

9 (6), 2207-2213, doi: 10.5194/acp-9-2207-2009, 

2009.

Luecken, D.J., R.L. Waterland, S. Papasavva, K. Taddonio, W.T. Hutzell, J.P. Rugh, and S.O. Andersen, Ozone and TFA impacts in North America 

from degradation of 2,3,3,3-tetrafluoropropene 

(HFO-1234yf), a potential greenhouse gas replacement, Environ. Sci. Technol., 44 (1), 343-348, doi: 

10.1021/es902481f, 2010.

MacAyeal, D.R., T.A. Scambos, C.L. Hulbe, and M.A. 

Fahnestock, Catastrophic ice-shelf break-up by an 

ice-shelf-fragment-capsize mechanism, J. Glaciol.,

49 (164), 22-36, 2003.

Manne, A.S., and R.G. Richels, An alternative approach 

to establishing trade-offs among greenhouse gases, 

Nature, 410, 675-677, doi: 10.1038/35070541, 

2001.

Martinerie, P., E. Nourtier-Mazauric, J.-M. Barnola, W.T. 

Sturges, D.R. Worton, E. Atlas, L.K. Gohar, K.P. 

Shine, and G.P. Brasseur, Long-lived halocarbon 
5.42

Chapter 5

trends and budgets from atmospheric chemistry 

modelling constrained with measurements in polar 

firn, Atmos. Chem. Phys., 9 (12), 3911-3934, doi: 

10.5194/acp-9-3911-2009, 2009.

Maurice, L.Q., and D.S. Lee (eds.), Assessing Current 

Scientific Knowledge, Uncertainties and Gaps in 

Quantifying Climate Change, Noise and Air Quality Aviation Impacts, L.Q. Maurice, D.S. Lee, D.J. 

Wuebbles, I. Isaksen, L. Finegold, M. Vallet, M. 

Pilling, and J. Spengler, Final Report of the International Civil Aviation Organization (ICAO) Committee on Aviation and Environmental Protection 

(CAEP) Workshop, U.S. Federal Aviation Administration and Manchester Metropolitan University, 

Washington D.C. and Manchester, available: http://

web.mit.edu/aeroastro/partner/reports/caepimpactreport.pdf, 2009.

Meilinger, S.K., B. Kärcher, and Th. Peter, Microphysics 

and heterogeneous chemistry in aircraft plumes – 

high sensitivity on local meteorology and atmospheric composition, Atmos. Chem. Phys., 5 (2),

533-545, doi: 10.5194/acp-5-533-2005, 2005.

Molina, M.J., and F.S. Rowland, Stratospheric sink for 

chlorofluoromethanes: Chlorine atomc-atalysed 

destruction of ozone, Nature, 249, 5460, 810-812, 

doi: 10.1038/249810a0, 1974.

Montzka, S.A., J.H. Butler, B.D. Hall, D.J. Mondeel, 

and J.W. Elkins, A decline in tropospheric organic bromine, Geophys. Res. Lett., 30 (15), doi: 

10.1029/2003GL017745, 2003.

Montzka, S.A., B.D. Hall, and J.W. Elkins, Accelerated 

increases observed for hydrochlorofluorocarbons 

since 2004 in the global atmosphere, Geophys. Res. 

Lett., 36, L03804, doi: 10.1029/2008GL036475, 

2009.

Montzka, S.A., L. Kuijpers, M.O. Battle, M. Aydin, K.R. 

Verhulst, E.S. Saltzman, and D.W. Fahey, Recent 

increase in global HFC-23 emissions, Geophys. Res. 

Lett., 37, L02808, doi: 10.1029/2009GL041195, 

2010.

Morgenstern, O., P. Braesicke, M.M. Hurwitz, F.M. 

O’Connor, A.C. Bushell, C.E. Johnson, and J.A. 

Pyle, The world avoided by the Montreal Protocol, Geophys. Res. Lett., 35, L16811, doi: 

10.1029/2008GL034590, 2008.

Morgenstern, O., P. Braesicke, F.M. O’Connor, A.C. 

Bushell, C.E. Johnson, S.M. Osprey, and J.A. Pyle, 

Evaluation of the new UKCA climate-composition 

model – part 1: The stratosphere, Geosci. Model 

Dev., 2 (1), 43-57, doi: 10.5194/gmd-2-43-2009,

2009.

Newman, P.A., E.R. Nash, S.R. Kawa, S.A. Montzka, and 

S.M. Schauffler, When will the Antarctic ozone 

hole recover?, Geophys. Res. Lett., 33, L12814, 

doi: 10.1029/2005GL025232, 2006.

Newman, P.A., J.S. Daniel, D.W. Waugh, and E.R. Nash, 

A new formulation of equivalent effective stratospheric chlorine (EESC), Atmos. Chem. Phys., 7

(17), 4537-4552, doi: 10.5194/acp-7-4537-2007, 

2007.

Newman, P.A., L.D. Oman, A.R. Douglass, E.L. Fleming, S.M. Frith, M.M. Hurwitz, S.R. Kawa, C.H. 

Jackman, N.A. Krotkov, E.R. Nash, J.E. Nielsen, S. 

Pawson, R.S. Stolarski, and G.J.M. Velders, What 

would have happened to the ozone layer if chlorofluorocarbons (CFCs) had not been regulated?, Atmos. Chem. Phys., 9 (6), 2113-2128, doi: 10.5194/

acp-9-2113-2009, 2009.

Nielsen, O.J., M.S. Javadi, M.P. Sulbaek Andersen, M.D. 

Hurley, T.J. Wallington, and R. Singh, Atmospher Atmospheric chemistry of CF3CF=CH2: Kinetics and mechanisms of gas-phase reactions with Cl atoms, OH 

radicals, and O3, Chem. Phys. Lett., 439, 18-22, doi: 

10.1016/j.cplett.2007.03.053, 2007.

Ottinger Schaefer, D., D. Godwin, and J. Harnisch, Estimating future emissions and potential reductions of 

HFCs, PFCs, and SF6, Energy J. Special Issue 3, 

27, 63-68, in Multi-Greenhouse Gas Mitigation and 

Climate Policy, edited by F.C. de la Chesnaye and 

J.P. Weyant, International Association for Energy 

Economics, 2006.

Papadimitriou, V.C., R.W. Portmann, D.W. Fahey, J. 

Mühle, R.F. Weiss, and J.B. Burkholder, Experimental and theoretical study of the atmospheric 

chemistry and global warming potential of SO2F2, 

J. Phys. Chem. A, 112, 12657-12666, doi: 10.1021/

jp806368u, 2008a.

Papadimitriou, V.C., R.K. Talukdar, R.W. Portmann, A.R. 

Ravishankara, and J.B. Burkholder, CF3CF=CH2

and (Z)-CF3CF=CHF: Temperature dependent 

OH rate coefficients and global warming potentials, Phys. Chem. Chem. Phys., 10, 808-820, doi: 

10.1039/b714382f, 2008b.

Parson, E.A., Protecting the Ozone Layer: Science and 

Strategy, 396 pp., Oxford University Press, New 

York, NY, U.S.A., 2003.

Pawson, S., R.S. Stolarski, A.R. Douglass, P.A. Newman, 

J.E. Nielsen, S.M. Frith, and M.L. Gupta, Goddard 

Earth Observing System chemistry-climate model 

simulations of stratospheric ozone-temperature 

coupling between 1950 and 2005, J. Geophys. Res.,

113, D12103, doi: 10.1029/2007JD009511, 2008.

Pinnock, S., M.D. Hurley, K.P. Shine, T.J. Wallington, 

and T.J. Smyth, Radiative forcing of climate by 

hydrochlorofluorocarbons and hydrofluorocarbons, 

J. Geophys. Res., 100 (D1), 23227-23238, doi: 

10.1029/95JD02323, 1995.

Pitari, G., D. Iachetti, E. Mancini, V. Montanaro, N. De 
5.43

A Focus on Information for Policymakers

Luca, C. Marizy, O. Dessens, H. Rogers, J. Pyle, V. 

Grewe, A. Stenke, and O.A. Søvde, Radiative forcing from particle emissions by future supersonic 

aircraft, Atmos. Chem. Phys., 8 (14), 4069-4084, 

doi: 10.5194/acp-8-4069-2008, 2008.

Portmann, R.W., and S. Solomon, Indirect radiative forcing of the ozone layer during the 21st

century, Geophys. Res. Lett., 34, L02813, doi: 

10.1029/2006GL028252, 2007.

Portmann, R.W., S. Solomon, R.R. Garcia, L.W. Thomason, L.R. Poole, and M.P. McCormick, Role of 

aerosol variations in anthropogenic ozone depletion 

in the polar regions, J. Geophys. Res., 101 (D17), 

22991-23006, 1996.

Prather, M.J., Time scales in atmospheric chemistry: Theory, GWPs for CH4 and CO, and runaway growth, 

Geophys. Res. Lett., 23 (19), 2597-2600, 1996.

Prather, M.J., Lifetimes of atmospheric species: Integrating environmental impacts, Geophys. Res. Lett., 29

(22), 2063, doi: 10.1029/2002GL016299, 2002.

Prather, M.J., and J. Hsu, NF3, the greenhouse gas missing 

from Kyoto, Geophys. Res. Lett., 35, L12810, doi: 

10.1029/2008GL034542, 2008.

Prather, M., P. Midgley, F. Sherwood Rowland, and R. 

Stolarski, The ozone layer: The road not taken, Nature, 381, 551-554, doi: 10.1038/381551a0, 1996.

Prinn, R.G. and R. Zander (Lead Authors), D.M. Cunnold, 

J.W. Elkins, A. Engel, P.J. Fraser, M.R. Gunson, 

M.K.W. Ko, E. Mahieu, P.M. Midgley, J.M. Russell III, C.M. Volk, and R.F. Weiss, Long-lived 

ozone-related compounds, Chapter 1 in Scientific 

Assessment of Ozone Depletion: 1998, Global 

Ozone Research and Monitoring Project—Report 

No. 44, World Meteorological Organization, Geneva, Switzerland, 1999.

Rajakumar, B., R.W. Portmann, J.B. Burkholder, and A.R. 

Ravishankara, Rate coefficients for the reactions of 

OH with CF3CH2CH3 (HFC-263fb), CF3CHFCH2F 

(HFC-245eb), and CHF2CHFCHF2 (HFC-245ea) 

between 238 and 375 K, J. Phys. Chem. A, 110, 

6724-6731, doi: 10.1021/jp056248y, 2006.

Rao, S., and K. Riahi, The role of non-CO2 greenhouse 

gases in climate change mitigation: Long-term 

scenarios for the 21st century, Energy J. Special 

Issue 3, 27, 177-200, in Multi-Greenhouse Gas 

Mitigation and Climate Policy, edited by F.C. de la 

Chesnaye, and J.P. Weyant, International Association for Energy Economics, 2006.

Rasch, P.J., P.J. Crutzen, and D.B. Coleman, Exploring the geoengineering of climate using stratospheric sulfate aerosols: The role of particle size, 

Geophys. Res. Lett., 35, L02809, doi: 10.1029/ 

2007GL032179, 2008a.

Rasch, P.J., S. Tilmes, R.P. Turco, A. Robock, L. Oman, 

C.-C. Chen, G.L. Stenchikov, and R.R. Garcia, 

An overview of geoengineering of climate using 

stratospheric sulphate aerosols, Phil. Trans. Royal Soc. A., 366 (1882), 4007-4037, doi: 10.1098/

rsta.2008.0131, 2008b.

Ravishankara, A.R., A.A. Turnipseed, N.R. Jensen, S. Barone, M. Mills, C.J. Howard, and S. Solomon, Do 

hydrofluorocarbons destroy stratospheric ozone?, 

Science, 263, 71-75, 1994.

Ravishankara, A.R., J.S. Daniel, and R.W. Portmann, Nitrous oxide (N2O): The dominant ozone-depleting 

substance emitted in the 21st century, Science, 326

(5949), 123-125, doi: 10.1126/science.1176985, 

2009.

Riahi, K., and N. Nakicenovic (eds.), Greenhouse Gases– 

Integrated Assessment, Technol. Forecast. Soc. 

Change, Special Issue, 74 (7), September 2007, 

234 pp., ISSN 0040-1625, doi: 10.1016/S0040-

1625(07)00140-0, 2007.

Robock, A., 20 reasons why geoengineering may be 

a bad idea, Bull. Atom. Sci., 64 (2), 14-18, doi: 

10.2968/064002006, 2008.

Robock, A., L. Oman, and G. Stenchikov, Regional climate responses to geoengineering with tropical 

and Arctic SO2 injections, J. Geophys. Res., 113, 

D16101, doi: 10.1029/2008JD010050, 2008.

Rosenfield, J.E., A.R. Douglass, and D.B. Considine, 

The impact of increasing carbon dioxide on ozone 

recovery, J. Geophys. Res., 107 (D6), 4049, doi: 

10.1029/2001JD000824, 2002.

Ross, M., D. Toohey, M. Peinemann, and P. Ross, Limits 

on the space launch market related to stratospheric ozone depletion, Astropolitics, 7, 50-82, doi: 

10.1080/14777620902768867, 2009.

Royal Society, Geoengineering the Climate: Science,

Governance and Uncertainty, Report 10/09, London GB, The Royal Society, available: http://royalsociety.org/geoengineering-the-climate/, 2009.

Schauffler, S.M., E.L. Atlas, S.G. Donnelly, A. Andrews, 

S.A. Montzka, J.W. Elkins, D.F. Hurst, P.A. Romashkin, G.S. Dutton, and V. Stroud, Chlorine 

budget and partitioning during the Stratospheric 

Aerosol and Gas Experiment (SAGE) III Ozone 

Loss and Validation Experiment (SOLVE), J. 

Geophys. Res., 108 (D5), 4173, doi: 10.1029/ 

2001JD002040, 2003.

Scott, B.F., R.W. Macdonald, K. Kannan, A. Fisk, A. 

Witter, N. Yamashita, L. Durham, C. Spencer, 

and D.C.G. Muir, Trifluoroacetate profiles in the 

Arctic, Atlantic, and Pacific Oceans, Environ. 

Sci. Technol., 39 (17), 6555-6560, doi: 10.1021/ 

es047975u, 2005.

Scott, B.F., C.A. Moody, C. Spencer, J.M. Small, 

D.C.G. Muir, and S.A. Mabury, Analysis for 
5.44

Chapter 5

perfluorocarboxylic acids/anions in surface waters and precipitation using GC-MS and analysis 

of PFOA from large-volume samples, Environ. 

Sci. Technol., 40 (20), 6405-6410, doi: 10.1021/

es061131o, 2006.

Seedhouse, E., Tourists in Space: A Practical Guide, 

Springer Praxis Books, Chichester, U.K., doi: 

10.1007/978-0-387-74644-9, 2008.

Shine, K.P., On the cause of the relative greenhouse strength of gases such as the halocarbons, J. Atmos. Sci., 48, 1513-1518, doi: 

10.1175/1520-0469(1991)048<1513:OTCOTR>2.

0.CO;2, 1991.

Shine K.P., The global warming potential—the need for 

an interdisciplinary retrial, Clim. Change, 96, 467-

472, doi: 10.1007/s10584-009-9647-6, 2009.

Smith, S.J., and T.M.L. Wigley, Multi-gas forcing stabilization with Minicam, Energy J. Special Issue 3, 27, 

373-392, in Multi-Greenhouse Gas Mitigation and 

Climate Policy, edited by F.C. de la Chesnaye and 

J.P. Weyant, International Association for Energy 

Economics, 2006.

Soden, B.J., R.T. Wetherald, G.L. Stenchikov, and A. Robock, Global cooling after the eruption of Mount 

Pinatubo: A test of climate feedback by water vapor, Science, 296 (5568), 727-730, doi: 10.1126/

science.296.5568.727, 2002.

Solomon, S., and D.L. Albritton, Time-dependent ozone 

depletion potentials for short- and long-term forecasts, Nature, 357, 33-37, doi: 10.1038/357033a0, 

1992.

Solomon, S., M. Mills, L.E. Heidt, W.H. Pollock, and A.F. 

Tuck, On the evaluation of ozone depletion potentials, J. Geophys. Res., 97 (D1), 825-842, 1992.

Solomon, S., R.W. Portmann, R.R. Garcia, L.W. Thomason, L.R. Poole, and M.P. McCormick, The role of 

aerosol variations in anthropogenic ozone depletion at northern midlatitudes, J. Geophys. Res., 101 

(D3), 6713-6727, 1996.

Solomon, K.R., X. Tang, S.R. Wilson, P. Zanis, and A.F. 

Bais, Changes in tropospheric composition and air 

quality due to stratospheric ozone depletion, Photochem. Photobiol. Sci., 2, 62-67, doi: 10.1039/

b211086e, 2003.

Søvde, O.A., M. Gauss, I.S.A. Isaksen, G. Pitari, and C. 

Marizy, Aircraft pollution — a futuristic view, Atmos. Chem. Phys., 7 (13), 3621-3632, doi: 10.5194/

acp-7-3621-2007, 2007.

SPARC CCMVal (Stratospheric Processes And their Role 

in Climate), SPARC CCMVal Report on the Evaluation of Chemistry-Climate Models, edited by V. 

Eyring, T.G. Shepherd, and D.W. Waugh, SPARC 

Report No. 5, WCRP-132, WMO/TD-No. 1526, 

available: http://www.atmosp.physics.utoronto.ca/

SPARC/ccmval_final/index.php, 2010.

Stohl, A., P. Seibert, J. Arduini, S. Eckhardt, P. Fraser, 

B.R. Greally, C. Lunder, M. Maione, J. Mühle, S. 

O’Doherty, R.G. Prinn, S. Reimann, T. Saito, N. 

Schmidbauer, P.G. Simmonds, M.K. Vollmer, R.F. 

Weiss, and Y. Yokouchi, An analytical inversion 

method for determining regional and global emissions of greenhouse gases: Sensitivity studies and 

application to halocarbons, Atmos. Chem. Phys.,

9 (5), 1597-1620, doi: 10.5194/acp-9-1597-2009, 

2009.

Sulbaek Andersen, M.P., D.R. Blake, F.S. Rowland, 

M.D. Hurley, and T.J. Wallington, Atmospheric 

chemistry of sulfuryl flouride: Reaction with OH 

radicals, Cl atoms and O3, atmospheric lifetime, IR 

spectrum, and global warming potential, Environ. 

Sci. Technol., 43 (4), 1067-1070, doi: 10.1021/

es802439f, 2009.

Tabazadeh, A., K. Drdla, M.R. Schoeberl, P. Hamill, and 

O.B. Toon, Arctic “ozone hole” in a cold volcanic 

stratosphere, Proc. Natl. Acad. Sci., 99 (5), 2609-

2612, doi: 10.1073/pnas.052518199, 2002.

Tanaka, K., B.C. O’Neill, D. Rokityanskiy, M. Obersteiner, and R.S.J. Tol, Evaluating global warming potentials with historical temperature, Clim. Change,

96, 443-466, doi: 10.1007/s10584-009-9566-6, 

2009.

Tang, X., S. Madronich, T. Wallington, and D. Calamari, 

Changes in tropospheric composition and air quality, J. Photochem. Photobiol. B: Biol., 46 (1-3), 83-

95, doi: 10.1016/S1011-1344(98)00187-0, 1998.

Taniguchi, N., T.J. Wallington, M.D. Hurley, A.G. Guschin, L.T. Molina, and M.J. Molina, Atmospheric 

chemistry of C2F5C(O)CF(CF3)2: Photolysis and 

reaction with Cl atoms, OH radicals, and ozone, J. 

Phys. Chem. A, 107 (15), 2674-2679, doi: 10.1021/

jp0220332, 2003.

TEAP (Technology and Economic Assessment Panel), 

Task Force Decision XX/8 Report, Assessment of 

Alternatives to HCFCs and HFCs and Update of the 

TEAP 2005 Supplement Report Data, coordinated 

by L. Kuijpers, and D. Verdonik, United Nations 

Environment Programme, Nairobi, Kenya, available: http://ozone.unep.org/teap/Reports/TEAP_

Reports/teap-may-2009-decisionXX-8-task-forcereport.pdf, 2009.

Teller, E., L. Wood, and R. Hyde, Global warming and Ice 

Ages: I. Prospects for physics-based modulation of 

global change, UCRL-JC-128715, Lawrence Livermore National Laboratory, Livermore, California, 

U.S., 20 pp., 1997.

Thompson, D.W.J., and S. Solomon, Interpretation of 

recent Southern Hemisphere climate change, 

Science, 296 (5569), 895-899, doi: 10.1126/
5.45

A Focus on Information for Policymakers

science.1069270, 2002.

Tilmes, S., R. Müller, R.J. Salawitch, U. Schmidt, C.R. 

Webster, H. Oelhaf, C.C. Camy-Peyret, and J.M. 

Russell III, Chemical ozone loss in the Arctic win Chemical ozone loss in the Arctic winter 1991–1992, Atmos. Chem. Phys., 8 (7), 1897-

1910, doi: 10.5194/acp-8-1897-2008, 2008a.

Tilmes, S., R. Müller, and R. Salawitch, The sensitivity 

of polar ozone depletion to proposed geoengineering schemes, Science, 320 (5880), 1201-1204, doi: 

10.1126/science.1153966, 2008b.

Tilmes, S., R.R. Garcia, D.E. Kinnison, A. Gettelman, 

and P.J. Rasch, Impact of geo-engineered aerosols 

on troposphere and stratosphere, J. Geophys. Res., 

114, D12305, doi: 10.1029/2008JD011420, 2009.

Tromp, T.K., M.K.W. Ko, J.M. Rodriguez, and N.D. Sze, 

Potential accumulation of a CFC-replacement degradation product in seasonal wetlands, Nature, 376, 

327-330, doi: 10.1038/376327a0, 1995.

UBA (Umweltbundesamt), Projections of Global Emissions of Fluorinated Greenhouse Gases in 2050, B. 

Gschrey and W. Schwarz, German Federal Environment Agency (Umweltbundesamt), Report-no. 

UBA-FB 001318, available: http://www.umweltdaten.de/publikationen/fpdf-l/3866.pdf, 2009.

UNEP (United Nations Environment Programme), data 

on production and consumption of ozone depleting substances under the Montreal Protocol, downloaded from http://www.unep.org/ozone, Nairobi, 

Kenya, 2009.

van Vuuren, D.P., B. Eickhout, P.L. Lucas, and M.D.G. 

Den Elzen, Long-term multi-gas scenarios to stabilise radiative forcing, Energy J. Special Issue 3, 27, 

201-233, in Multi-Greenhouse Gas Mitigation and 

Climate Policy, edited by F.C. de la Chesnaye and 

J.P. Weyant, International Association for Energy 

Economics, 2006.

van Vuuren, D.P., M.G.J. den Elzen, P.L. Lucas, B. Eickhout, B.J. Strengers, B. van Ruijven, S. Wonink, 

and R. van Houdt, Stabilizing greenhouse gas concentrations at low levels: An assessment of reduction strategies and costs, Clim. Change, 81 (2), 119-

159, doi: 10.1007/s10584-006-9172-9, 2007.

Velders, G.J.M., S.O. Andersen, J.S. Daniel, D.W. Fahey, 

and M. McFarland, The importance of the Montreal Protocol in protecting climate, Proc. Natl. 

Acad. Sci., 104 (12), 4814-4819, doi: 10.1073/

pnas.0610328104, 2007.

Velders, G.J.M., D.W. Fahey, J.S. Daniel, M. McFarland, 

and S.O. Andersen, The large contribution of projected HFC emissions to future climate forcing, 

Proc. Natl. Acad. Sci., 106 (27), 10949-10954, doi: 

10.1073/pnas.0902817106, 2009.

Volk, C.M., J.W. Elkins, D.W. Fahey, G.S. Dutton, J.M. 

Gilligan, M. Loewenstein, J.R. Podolske, K.R. 

Chan, and M.R. Gunson, Evaluation of source gas 

lifetimes from stratospheric observations, J. Geophys. Res., 102 (D21), 25543-25564, 1997.

Von Sydow, L.M., A.B. Grimvall, H.B. Borén, K. 

Laniewski, and A.T. Nielsen, Natural background 

levels of trifluoroacetate in rain and snow, Environ. Sci. Technol., 34, 3115-3118, doi: 10.1021/

es9913683, 2000.

Wallington, T.J., W.F. Schneider, D.R. Worsnop, O.J. 

Nielsen, J. Sehested, W. DeBruyn, and J.A. Shorter, Environmental impact of CFC replacements — 

HFCs and HCFCs, Environ. Sci. Technol., 28 (7), 

320A-326A, 1994.

Wallington, T.J., W.F. Schneider, J. Sehested, and O.J. 

Nielsen, Hydrofluorocarbons and stratospheric 

ozone, Faraday Discuss., 100, 55-64, 1995.

Wallington, T.J., M.D. Hurley, and O.J. Nielsen, The radiative efficiency of HCF2OCF2OCF2CF2OCF2H 

(H-Galden 1040x) revisited, Atmos. Environ., 43, 

4247-4249, doi: 10.1016/j.atmosenv.2009.05.046, 

2009.

Wallington, T.J., M.P. Sulbaek Andersen, and O.J. 

Nielsen, Estimated photochemical ozone creation 

potentials (POCPs) of CF3CF=CH2 (HFO-1234yf) 

and related hydrofluoroolefins (HFOs), Atmos. 

Environ., 44 (11), 1478-1481, doi: 10.1016/j.atmosenv.2010.01.040, 2010.

Wang, M., Greenhouse Gases, Regulated Emissions, and 

Energy Use in Transportation (GREET) Model, 

Version 1.8d.1, downloaded August 27, 2010, 

available: http://www.transportation.anl.gov/modeling_simulation/GREET/, 2010.

Waugh, D.W., L. Oman, S.R. Kawa, R.S. Stolarski, S. 

Pawson, A.R. Douglass, P.A. Newman, and J.E. 

Nielsen, Impacts of climate change on stratospheric 

ozone recovery, Geophys. Res. Lett., 36, L03805, 

doi: 10.1029/2008GL036223, 2009.

Wigley, T.M.L., A combined mitigation/geoengineering approach to climate stabilization, Science, 314

(5798), 452- 454, doi: 10.1126/science.1131728, 

2006.

Wise, M., K. Calvin, A. Thomson, L. Clarke, B. BondLamberty, R. Sands, S.J. Smith, and A. Janetos, J. 

Edmonds, Implications of limiting CO2 concentrations for land use and energy, Science, 324 (5931), 

1183-1186, doi: 10.1126/science.1168475, 2009.

WMO (World Meteorological Organization), Atmospheric 

Ozone 1985: Assessment of Our Understanding of 

the Processes Controlling its Present Distribution 

and Change (3 volumes), Global Ozone Research 

and Monitoring Project–Report No. 16, Geneva, 

Switzerland, 1986.

WMO (World Meteorological Organization), Scientific 

Assessment of Stratospheric Ozone: 1989, Volume 
5.46

Chapter 5

I, Global Ozone Research and Monitoring Project–

Report No. 20, Geneva, Switzerland, 1989.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 1991, Global 

Ozone Research and Monitoring Project–Report 

No. 25, Geneva, Switzerland, 1991.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 1994, Global 

Ozone Research and Monitoring Project–Report 

No. 37, Geneva, Switzerland, 1995.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 1998, Global 

Ozone Research and Monitoring Project–Report 

No. 44, Geneva, Switzerland, 1999.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 2002, Global 

Ozone Research and Monitoring Project–Report 

No. 47, Geneva, Switzerland, 2003.

WMO (World Meteorological Organization), Scientific 

Assessment of Ozone Depletion: 2006, Global 

Ozone Research and Monitoring Project–Report 

50, Geneva, Switzerland, 2007.

Wuebbles, D.J., The Relative Efficiency of a Number of 

Halocarbons for Destroying Stratospheric Ozone, 

11 pp., Lawrence Livermore National Laboratory, 

Livermore, CA, 1981.

Wuebbles, D.J., Chlorocarbon emission scenarios: Potential impact on stratospheric ozone, Geophys. Res. 

Lett., 88 (C2), 1433-1443, 1983.

Wuebbles, D.J., Nitrous oxide: No laughing matter, 

Science, 326 (5949), 56-57, doi: 10.1126/science.1179571, 2009.

Wuebbles, D.J., and K.O. Patten, Three-dimensional modeling of HCFC-123 in the atmosphere: Assessing 

its potential environmental impacts and rationale 

for continued use, Environ. Sci. Technol., 43 (9), 

3208-3213, doi: 10.1021/es802308m, 2009.

Wuebbles, D.J., A. Douglass, B. Kärcher, W.-C. Wang, 

S. Baughcum, R.P. d’Entremont, A. Dessler, P. 

Ginoux, R. Herman, A. Heymsfield, I. Isaksen, D. 

Jacob, S. Lele, J. Logan, J. McConnell, R. MiakeLye, P. Minnis, D. Mitchell, D. Murphy, L. Pan, 

J. Penner, M. Prather, J. Rodriguez, K. Rosenlof, 

K. Sassen, R. Sausen, K. Shine, A. Tabazadeh, I. 

Waitz, P. Yang, and F. Yu, Workshop on the Impacts of Aviation on Climate Change: A Report of 

Findings and Recommendations, Partnership for 

Air Transportation Noise and Emissions Reduction, Report No. PARTNER-COE-2006-004, available: http://web.mit.edu/aeroastro/partner/reports/

climatewrksp-rpt-0806.pdf, 2006.

Wuebbles, D.J., D. Youn, K. Patten, D. Wang, and M. 

Martínez-Avilés, Metrics for ozone and climate: 

Three-dimensional modeling studies of ozone 

depletion potentials and indirect global warming 

potentials, in Twenty Years of Ozone Decline. Proceedings of the Symposium for the 20th Anniversary 

of the Montreal Protocol, edited by C. Zerefos, G. 

Contopoulos, and G. Skalkeas, 297-326, Springer Publishing, Dordrecht, The Netherlands, doi: 

10.1007/978-90-481-2469-5, 2009.

Youn, D., K.O. Patten, J.-T. Lin, and D.J. Wuebbles, Explicit calculation of indirect global warming potentials for halons using atmospheric models, Atmos. 

Chem. Phys., 9 (22), 8719-8733, doi: 10.5194/acp9-8719-2009, 2009.

Youn, D., K.O. Patten, D.J. Wuebbles, H. Lee, and C.-

W. So, Potential impact of iodinated replacement 

compounds CF3I and CH3I on atmospheric ozone: 

A three-dimensional modeling study, Atmos. Chem. 

Phys., 10 (20), 10129-10144, doi: 10.5194/acp-10-

10129-2010, 2010.

Young, C.J., M.D. Hurley, T.J. Wallington, and S.A. Mabury, Atmospheric lifetime and global warming potential of a perfluoropolyether, Environ. Sci. Technol., 40, 2242-2246, doi: 10.1021/es052077z, 2006.

Young, C.J., M.D. Hurley, T.J. Wallington, and S.A. 

Mabury, Molecular structure and radiative efficiency of fluorinated ethers: A structure-activity 

relationship, J. Geophys. Res., 113, D24301 doi: 

10.1029/2008JD010178, 2008.

Young, C.J., M.D. Hurley, T.J. Wallington, and S.A. 

Mabury, Atmospheric chemistry of CF3CF2H and 

CF3CF2CF2CF2H: Kinetics and products of gasphase reactions with Cl atoms and OH radicals, 

infrared spectra, and formation of perfluorocarboxylic acids, Chem. Phys. Lett., 473, 251-256, doi: 

10.1016/j.cplett.2009.04.001, 2009.

Yvon-Lewis, S.A., E.S. Saltzman, and S.A. Montzka, Recent trends in atmospheric methyl bromide: Analysis of post-Montreal Protocol variability, Atmos. 

Chem. Phys., 9 (16), 5963-5974, doi: 10.5194/acp9-5963-2009, 2009.
5.47

A Focus on Information for Policymakers

Table 5A-1. Direct Global Warming Potentials for selected gases.

Industrial

Designation or

Common Name

Chemical Formula Radiative

Efficiency a

(W m−2 ppb−1)

Lifetime b

(years)

Global Warming Potential for

Given Time Horizon

20 years 100 years 500 years

Carbon dioxide CO2 1.38×10−5 c 1 1 1

Nitrous oxide N2O 3.03×10−3 114 289 298 153

Chlorofluorocarbons

CFC-11 CCl3F 0.25 45 6,730 4,750 1,620

CFC-12 CCl2F2 0.32 100 11,000 10,900 5,200

CFC-13 CClF3 0.25 640 10,800 14,400 16,400

CFC-113 CCl2FCClF2 0.30 85 6,540 6,130 2,690

CFC-114 CClF2CClF2 0.31 190 7,890 9,180 6,330

CFC-115 CClF2CF3 0.18 1,020 5,290 7,230 9,120

Hydrochlorofluorocarbons

HCFC-21 CHCl2F 0.14 1.7 530 151 46

HCFC-22 CHClF2 0.20 11.9 5,130 1,790 545

HCFC-123 CHCl2CF3 0.14 1.3 273 77 24

HCFC-124 CHClFCF3 0.22 5.9 2,110 619 188

HCFC-141b CH3CCl2F 0.14 9.2 2,240 717 218

HCFC-142b CH3CClF2 0.20 17.2 5,390 2,220 678

HCFC-225ca CHCl2CF2CF3 0.20 1.9 429 122 37

HCFC-225cb CHClFCF2CClF2 0.32 5.9 2,060 606 184

Hydrofluorocarbons

HFC-23 CHF3 0.19 222.0 11,900 14,200 10,700

HFC-32 CH2F2 0.11 5.2 2,470 716 218

HFC-41 CH3F 0.02 2.8 377 107 33

HFC-125 CHF2CF3 0.23 28.2 6,290 3,420 1,070

HFC-134 CHF2CHF2 0.18 9.7 3,420 1,110 339

HFC-134a CH2FCF3 0.16 13.4 3,730 1,370 416

HFC-143 CH2FCHF2 0.13 3.5 1,240 352 107

HFC-143a CH3CF3 0.13 47.1 5,780 4,180 1,440

HFC-152a CH3CHF2 0.09 1.5 468 133 40

HFC-227ea CF3CHFCF3 0.26 38.9 5,480 3,580 1,180

HFC-236cb CH2FCF2CF3 0.23 13.1 3,560 1,290 392

APPENDIX 5A

TABLES
5.48

Chapter 5

Industrial

Designation or

Common Name

Chemical Formula Radiative

Efficiency a

(W m−2 ppb−1)

Lifetime b

(years)

Global Warming Potential for

Given Time Horizon

20 years 100 years 500 years

HFC-236ea CHF2CHFCF3 0.30 11.0 4,170 1,410 430

HFC-236fa CF3CH2CF3 0.28 242.0 8,100 9,820 7,710

HFC-245ca CH2FCF2CHF2 0.23 6.5 2,440 726 221

HFC-245ea CHF2CHFCHF2 0.18 d 3.2 983 280 85

HFC-245eb CH2FCHFCF3 0.23 d 3.1 1,220 346 105

HFC-245fa CHF2CH2CF3 0.28 7.7 3,410 1,050 318

HFC-263fb CH3CH2CF3 0.13 d 1.2 365 104 31

HFC-329p CHF2CF2CF2CF3 0.31 d 28.4 4,640 2,530 792

HFC-365mfc CH3CF2CH2CF3 0.22 d 8.7 2,670 842 256

HFC-43-10mee CF3CHFCHFCF2CF3 0.40 16.1 4,170 1,660 506

Chlorocarbons

Methyl chloroform CH3CCl3 0.06 5.0 506 146 45

Carbon tetrachloride CCl4 0.13 26.0 2,700 1,400 435

Methyl chloride CH3Cl 0.01 1.0 45 13 4

Bromocarbons and Halons

Methyl bromide CH3Br 0.01 0.8 19 5 2

Bromodifluoromethane CHBrF2 0.14 5.2 1,250 362 110

Halon-1211 CBrClF2 0.30 16.0 4,750 1,890 575

Halon-1301 CBrF3 0.32 65.0 8,480 7,140 2,760

Halon-2402 CBrF2CBrF2 0.33 20.0 3,680 1,640 503

Fully Fluorinated Species

Nitrogen trifluoride NF3 0.21 500 13,300 17,500 18,500

Sulfur hexafluoride SF6 0.52 3,200 16,300 22,800 32,600

TrifluoromethylsulfurpentafluorideSF5CF3 0.57 800 e 13,200 17,800 21,400

Perfluoromethane CF4 0.10 50,000 5,210 7,390 11,200

Perfluoroethane C2F6 0.26 10,000 8,630 12,200 18,200

Perfluoropropane C3F8 0.26 2,600 6,310 8,830 12,500

Perfluorocyclopropane c-C3F6 0.42 ~3,000 12,800 17,900 25,500

Perfluorobutane C4F10 0.33 2,600 6,330 8,850 12,500

Perfluorocyclobutane c-C4F8 0.32 3,200 7,310 10,300 14,700

Perfluoropentane C5F12 0.41 4,100 6,510 9,150 13,300

Perfluorohexane C6F14 0.49 3,100 6,620 9,290 13,300

Perfluorodecalin C10F18 0.56 ~2,000 5,530 7,700 10,600

Table 5A-1, continued.
5.49

A Focus on Information for Policymakers

Table 5A-1, continued.

Industrial

Designation or

Common Name

Chemical Formula Radiative

Efficiency a

(W m−2 ppb−1)

Lifetime b

(years)

Global Warming Potential for

Given Time Horizon

20 years 100 years 500 years

Halogenated Alcohols and Ethers

HFE-125 CHF2OCF3 0.44 119.0 13,700 14,200 7,510

HFE-134 CHF2OCHF2 0.45 24.4 11,930 5,960 1,840

HFE-143a CH3OCF3 0.27 4.8 2,920 840 256

HFE-227ea CF3OCHFCF3 0.40 51.6 8,170 6,180 2,200

HFCE-235da2 

(isoflurane) CHF2OCHClCF3 0.38 3.5 1,650 470 143

HFE-236ea2

(desflurane)

CHF2OCHFCF3 0.44 10.8 5,460 1,840 560

HFE-236fa CF3OCH2CF3 0.34 7.5 3,240 988 300

HFE-245cb2 CH3OCF2CF3 0.32 4.9 2,350 680 207

HFE-245fa1 CF3OCH2CHF2 0.30 6.6 2,880 859 261

HFE-245fa2 CHF2OCH2CF3 0.31 5.5 2,540 740 225

HFE-254cb2 CH3OCF2CHF2 0.28 2.5 1,210 345 105

HFE-329mcc2 CF3CF2OCF2CHF2 0.49 22.5 6,300 3,000 925

HFE-338mcf2 CF3CF2OCH2CF3 0.43 7.5 3,150 963 293

HFE-347mcc3 CF3CF2CF2OCH3 0.34 5.0 1,910 553 168

HFE-347mcf2 CF3CF2OCH2CHF2 0.41 6.6 2,950 881 268

HFE-356mec3 CF3CHFCF2OCH3 0.30 3.8 1,430 408 124

HFE-356pcf2 CHF2CF2OCH2CHF2 0.37 5.7 2,580 754 229

HFE-356pcf3 CHF2CF2CH2OCHF2 0.39 3.5 1,710 488 148

HFE-356pcc3 CHF2CF2CF2OCH3 0.33 3.8 1,570 448 136

(CF3)2CHOCHF2 0.41 21.2 5,580 2,570 789

(CF3)2CFOCH3 0.31 3.7 1,310 373 114

HFE-43-10pccc124 CHF2OCF2OC2F4OCHF2 1.02 d 13.5 8,130 2,990 909

HFE-7100 (HFE-449 blend) C4F9OCH3 0.31 4.7 1,320 379 115

HFE-7200 (HFE-569 blend) C4F9OC2H5 0.30 0.8 208 59 18

HFE-236ca12 CHF2OCF2OCHF2 0.66 25.0 11,320 5,730 1,770

HFE-338pcc13 CHF2OCF2CF2OCHF2 0.87 12.9 8,660 3,120 949

(CF3)2CHOH 0.28 1.9 726 206 63

Miscellaneous

Sulfuryl fluoride SO2F2 0.22 d 36 7,580 4,740 1,540

a Taken from WMO (2007), unless stated otherwise. b Taken from Chapter 1, unless stated otherwise. c Assumes CO2 mixing ratio is 378 parts per million (ppm). d See Table 5-2. e Midpoint of range given in Chapter 1.
5.50

Chapter 5

Table 5A-2. 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) are used. Before this, WMO (2007) production values are 

generally used. In the future, annual figures are determined from the lesser of the Protocol limitations and the 

most recent annual estimates.

Emission: For the years when mixing ratio observations are available, emissions are calculated using the box model 

described in Daniel and Velders et al. (2007) with the lifetimes of Appendix Table 5A-1. Emissions before this 

are usually consistent with WMO (2007) but are also forced to yield mixing ratios that meld smoothly into the 

measurement record. Future emissions are determined using a fixed annual bank release fraction, calculated as 

the average ratio over the past 10 years between the annual emissions and estimated bank plus production.

Bank: The bank assumed to be in place at the start of the measurement record is set at such a value that the IPCC 

(2009) bank for 2008 is attained. The bank at the end of each year is equal to the bank at the beginning of the 

year plus production during that year minus emissions during that year.

Approach for Specific Species

Species Description

CFC-11 Production:

1950–1985: WMO (2007)

1986, 1989–2008: UNEP (2009) a, 1987–1988 interpolated from 1986 and 1989 values

2009–2010: Approved critical-use exemptions for non-Article 5(1) countries

2010: Approved critical-use exemptions for Article 5(1) countries, 2009 value interpolated from 

 2008 and 2010 values. All critical-use exemptions for CFCs are assumed to be CFC-11

After 2010: No production

Emission:

1979–2008: Emissions calculated from observed global abundance trends

2009–2100: Annual emissions are a constant fraction of 0.05 of the bank, extrapolated from the 

 previous years

Bank:

1979 bank set to 1,705 Kt to attain a 2008 bank of 1,420 Kt (TEAP, 2009). The bank projected 

 for 2008 in the previous Assessment was 1,380 Kt.

CFC-12 Same as CFC-11, except:

Annual emissions from 2009–2100 are a constant fraction of 0.15 of the bank

Bank:

1979 bank set to 941 Kt to attain a 2008 bank of 394 Kt (TEAP, 2009). The bank projected for 

 2008 in the previous Assessment was 499 Kt.

CFC-113 Same formalism as for CFC-11, except:

Annual emissions from 2009–2100 are a constant fraction of 0.5 of the estimated bank. (Although 

 our scenario has a non-zero bank in 2010, according to TEAP (2009) the bank has been zero 

since 2002.)

CFC-114 Same formalism as for CFC-11 except:

Bank calculated based on historic production and emission data

Annual emission from 2009–2100 is a constant fraction of 0.1 of the estimated bank

CFC-115 Same as CFC-11, except:

Emissions from 2009–2100 are a constant fraction of 0.02 of the bank

Bank:

2008 bank of 16 Kt (TEAP, 2009)
5.51

A Focus on Information for Policymakers

Approach for Specific Species

CCl4 Production and bank not considered due to gaps in our understanding of the origin of much of the 

global emission

Emission:

1951–1979: WMO (2003)

1980–2008: Emissions calculated from observed global abundance trends

2009–2050: Emissions decrease by 6% per year (average derived from observations and a 26-year 

 lifetime during 2004 to 2008)

2051–2100: No emission

CH3CCl3 Production not explicitly considered

Emission:

1950–1978: WMO (2003)

1979–2008: Emissions calculated from observed global abundance trends

2009–2014: Emission equal to the 2004–2008 average

2015–2100: No emission

HCFC-22 Production:

1950–1988: WMO (2003)

1989, 1992–2008: UNEP (2009)a, 1990–1991 interpolated from 1989 and 1992 values

Non-Article 5(1) countries: 2009 production is 2004–2008 average; 2010–2013 Montreal Protocol 

 phase-out schedule applied

The base level HCFC consumption in non-Article 5(1) countries is calculated as the consumption 

 in 1989 + 2.8% of the 1989 CFC consumption (ODP-weighted). In WMO (2007), the 

 2.8% of the CFC consumption was assigned completely to HCFC-22. In this Assessment it is 

 distributed over the three main HCFCs based on their historic production. This results in 

 larger future emissions of HCFC-141b and HCFC-142b.

Article 5(1) countries: 2009–2012 is linearly extrapolated from 2004–2008; 2013–2040 Montreal 

 Protocol phase-out schedule applied

Emission:

1950–1992: Emissions calculated to yield mixing ratios consistent with WMO (2007)

1993–2008: Emissions calculated from observed global abundance trends

2009–2100: Assume an annual bank release fraction of 0.18 (average of 2002–2008)

Bank:

1993 bank set to 302 Kt to attain a 2008 bank of 1,618 Kt (TEAP, 2009). The bank projected for 

 2008 in the previous Assessment was 2,083 Kt.

HCFC-141b Production as in HCFC-22

Emission:

1990–2008: Emissions calculated from observed global abundance trends

2009–2100: Annual emissions calculated assuming a bank release fraction of 0.05 (average of 

 2003–2008)

Bank:

1994 bank set to 115 Kt to attain a 2008 bank of 941 Kt (TEAP, 2009). The bank projected for 

 2008 in the previous Assessment was 961 Kt.

HCFC-142b Production as in HCFC-22

Emission:

1979–2008: Emissions calculated from observed global abundance trends

2009–2100: Annual emissions calculated assuming a bank release fraction of 0.11 (average of 

 1999–2008)

Bank:

1993 bank set to 95 Kt to attain a 2008 bank of 273 Kt (TEAP, 2009). The bank projected for 

 2008 in the previous Assessment was 211 Kt.

Table 5A-2, continued.
5.52

Chapter 5

Approach for Specific Species

Halon-1211 Production:

1986, 1989–2008: UNEP (2009), 1987–1988 interpolated from 1986 and 1989 values

non-Article 5(1) countries: zero production from 2009–2100

Article 5(1) countries: 2009 production is 2006–2008 average; zero from 2010–2100

Emission:

1950–1992: WMO (2007)

1993–2008: Emissions calculated from observed global abundance trends

2009–2100: Annual emissions calculated assuming a bank release fraction of 0.075 (average of 

 1999–2008)

Bank:

1993 bank is set to 128 Kt to attain a 2008 bank of 74 Kt (TEAP, 2009). The bank projected for 

 2008 in the previous Assessment was 86 Kt. 

Halon-1301 Production as in halon-1211

Emission:

1950–1992: WMO (2007)

1992–2008: Emissions calculated from observed global abundance trends

2009–2100: Annual emissions calculated assuming a bank release fraction of 0.04 (average of 

 1999–2008)

Bank:

1992 bank is set to 74 Kt to attain a 2008 bank of 47 Kt (TEAP, 2009). The bank projected for 

 2008 in the previous Assessment was 31 Kt.

Halon-1202 Same as WMO (2007)

Halon-2402 Production as in halon-1211

Emission:

1970–2003: based on WMO (2007), but adjusted to meld smoothly into the measurement record 

2004–2008: Emissions calculated from observed global abundance trends

2009–2100: Annual emissions calculated assume a bank release fraction of 0.08 (average of 

 2004–2008)

Bank:

1995 bank is set to 20 Kt to allow the emissions meld smoothly into the emissions derived from 

 observations. The corresponding 2008 bank of 9.1 Kt is much larger than from TEAP (2009) 

 of 1.3 Kt. A bank of 1.3 Kt would result in emissions a factor of 8 smaller than derived from 

 observations for 2004–2008. The bank projected for 2008 in the previous Assessment was 

 1.8 Kt.

CH3Br Production/emissions from Yvon-Lewis et al. (2009): natural production/emission (102 Kt), gasoline 

(5.7 Kt), biomass (11.3 Kt), biofuel (6.1 Kt), non QPS-fumigation. Parenthetical values for 

anthropogenic sources are emission estimates for 1996 and 2007 (Yvon-Lewis et al., 2009 and 

references therein) and are held constant into the future.

QPS production from 1995–2008 from UNEP (2009), held constant at 11.2 Kt (2004–2008 

 average) for 2009–2100

Approved critical-use exemptions (CUE) for 2009–2011 (no CUE after 2011)

Emissions:

1950–2008: Emissions calculated from surface observations and derived from South Pole firn 

 measurements (using a global lifetime of 0.75 years to be consistent with Yvon-Lewis et al., 

 2009)

2009–2100: Fumigation emissions equal 0.6 times production for fumigation. The combination 

 of anthropogenic and natural emissions, constrained by the total emissions derived from 

 observations of concentrations, leads to an anthropogenic fraction of total emissions of 0.30 

Table 5A-2, continued.
5.53

A Focus on Information for Policymakers

Approach for Specific Species

(CH3Br,

continued)

 in 1992, in agreement with Montzka et al. (2003) (see also Chapter 1).

Bank:

No bank considered

CH3Cl Emission:

1950–1995: Global emissions derived from firn measurements at the South Pole

1996–2100: Emissions held constant at 1995 levels, as in WMO (2003) and WMO (2007)

a Estimated in cases in which reporting is not complete or reporting is made in compound classes rather than individually. The production data for each 

species in each year is obtained from UNEP (2009) and is consistent with the totals for each group of species as usually reported by UNEP.

Table 5A-2, continued.
5.54

Chapter 5Table 5A-3. Mixing ratios (ppt) of the ODSs considered in the baseline (A1) scenario 1. 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 40–130 ppt of stratospheric chlorine (see Chapter 1) are not shown in the table.

CFC-11 CFC-113 CFC-115 CH CCl 3 3 HCFC-141b Halon-1211 Halon-1301 CH Br 3

CFC-12 CFC-114 CCl4 HCFC-22 HCFC-142b Halon-1202 Halon-2402 CH Cl 3

1955 3.3 14.3 1.2 2.6 0.0 42.3 0.1 1.0 0.0 0.0 0.00 0.00 0.00 0.00 6.3 491.31960 9.5 29.5 1.9 3.8 0.0 52.1 1.5 2.1 0.0 0.0 0.00 0.00 0.00 0.00 6.5 510.31965 23.5 58.8 3.1 5.0 0.0 64.4 4.7 4.9 0.0 0.0 0.00 0.00 0.00 0.00 6.7 528.11970 52.8 114.3 5.5 6.5 0.2 75.9 16.2 12.1 0.0 0.0 0.02 0.00 0.00 0.02 7.0 539.91975 106.1 203.1 10.4 8.3 0.6 85.5 40.0 23.8 0.0 0.2 0.12 0.01 0.04 0.06 7.4 545.81980 161.9 297.4 19.0 10.7 1.3 93.3 81.6 42.5 0.0 0.4 0.42 0.01 0.24 0.15 7.7 548.31981 170.3 313.9 21.5 11.1 1.5 94.8 88.3 46.6 0.0 0.5 0.52 0.01 0.31 0.17 7.8 548.61982 179.2 329.9 25.3 11.6 1.8 95.8 93.0 50.7 0.0 0.5 0.63 0.02 0.39 0.19 7.9 548.91983 187.4 346.7 29.2 12.0 2.1 97.1 97.2 54.8 0.0 0.6 0.75 0.02 0.49 0.21 8.0 549.11984 196.0 363.9 32.5 12.4 2.4 98.2 101.7 58.8 0.0 0.7 0.88 0.02 0.61 0.24 8.1 549.31985 205.4 381.2 37.3 12.9 2.8 99.6 106.1 62.7 0.0 0.7 1.04 0.02 0.74 0.26 8.2 549.41986 215.8 399.3 42.0 13.4 3.1 100.8 110.0 66.9 0.0 0.8 1.22 0.02 0.89 0.29 8.2 549.51987 227.0 419.2 47.4 14.0 3.5 102.5 112.3 71.5 0.0 0.8 1.44 0.02 1.06 0.32 8.3 549.61988 238.0 438.4 54.4 14.5 3.9 103.6 118.0 76.7 0.0 0.9 1.69 0.03 1.25 0.35 8.4 549.71989 247.7 459.6 61.3 15.0 4.3 104.9 122.5 82.5 0.0 1.0 1.98 0.03 1.46 0.38 8.5 549.81990 256.2 477.5 67.6 15.4 4.7 106.5 127.2 88.2 0.0 1.2 2.27 0.03 1.66 0.41 8.6 549.81991 262.2 491.0 73.8 15.7 5.1 106.3 130.0 93.7 0.0 1.8 2.54 0.03 1.84 0.44 8.7 549.91992 265.2 500.7 78.7 15.8 5.6 105.3 131.4 99.1 0.1 2.8 2.79 0.03 1.97 0.46 8.9 549.91993 267.8 510.1 81.6 15.9 6.0 104.7 130.1 103.9 0.3 3.9 3.02 0.04 2.20 0.49 9.0 549.91994 267.8 517.1 83.0 16.0 6.4 104.0 121.5 109.2 1.2 5.1 3.20 0.04 2.48 0.51 9.1 549.91995 267.4 523.8 83.6 16.1 6.8 103.2 110.3 113.6 2.7 6.3 3.34 0.04 2.63 0.52 9.2 550.01996 266.8 528.9 83.7 16.2 7.1 102.5 97.8 119.4 4.5 7.3 3.50 0.05 2.60 0.51 9.1 550.01997 266.0 533.5 83.6 16.3 7.4 101.3 83.4 124.2 6.4 8.4 3.66 0.05 2.67 0.51 9.1 550.01998 264.5 536.4 83.3 16.4 7.6 100.5 70.6 128.7 8.2 9.3 3.80 0.05 2.70 0.51 9.2 550.01999 263.3 539.0 82.9 16.4 7.7 99.8 59.3 134.7 10.1 10.3 3.92 0.05 2.73 0.51 9.3 550.02000 261.7 541.0 82.3 16.5 7.9 98.6 49.7 139.5 11.8 11.4 4.02 0.04 2.84 0.50 8.9 550.02001 260.1 542.8 81.8 16.5 8.0 97.7 41.6 144.7 13.5 12.4 4.11 0.04 2.85 0.50 8.5 550.02002 258.3 543.4 81.1 16.5 8.0 96.6 34.6 150.9 14.8 13.2 4.17 0.03 2.87 0.50 8.2 550.02003 255.9 543.5 80.3 16.5 8.1 95.6 28.9 155.4 16.1 13.9 4.22 0.03 2.88 0.49 8.2 550.02004 253.9 543.4 79.6 16.5 8.2 94.7 24.1 160.3 17.0 14.5 4.25 0.02 2.95 0.49 8.0 550.0
5.55

A Focus on Information for Policymakers

CFC-11 CFC-113 CFC-115 CH CCl 3 3 HCFC-141b Halon-1211 Halon-1301 CH Br 3

CFC-12 CFC-114 CCl4 HCFC-22 HCFC-142b Halon-1202 Halon-2402 CH Cl 3

2005 251.6 542.7 78.8 16.6 8.3 93.7 20.1 165.5 17.5 15.1 4.26 0.02 3.03 0.48 7.9 550.02006 249.4 541.8 78.2 16.5 8.3 92.7 16.7 172.2 17.8 15.8 4.27 0.01 3.06 0.48 7.7 550.02007 247.2 539.8 77.6 16.5 8.4 91.6 14.1 179.3 18.6 16.9 4.24 0.01 3.12 0.48 7.6 550.02008 245.1 537.7 76.8 16.5 8.4 90.4 11.7 187.3 19.0 18.2 4.19 0.01 3.15 0.47 7.5 550.02009 243.3 535.3 76.2 16.4 8.4 89.0 9.8 195.2 19.6 19.3 4.14 0.00 3.18 0.47 7.2 550.02010 240.9 532.5 75.6 16.4 8.4 87.6 8.3 206.8 20.3 20.5 4.07 0.00 3.20 0.46 7.2 550.02011 238.4 529.4 74.8 16.3 8.4 86.0 7.1 218.6 21.1 21.7 3.99 0.00 3.22 0.46 7.2 550.02012 235.8 525.9 74.0 16.3 8.4 84.5 6.1 230.9 22.1 22.9 3.91 0.00 3.24 0.45 7.2 550.02013 233.1 522.3 73.2 16.2 8.4 82.8 5.3 243.6 23.2 24.2 3.82 0.00 3.26 0.44 7.1 550.02014 230.4 518.4 72.3 16.2 8.4 81.2 4.6 255.7 24.4 25.4 3.72 0.00 3.27 0.44 7.1 550.02015 227.6 514.4 71.5 16.1 8.4 79.5 4.1 267.1 25.6 26.6 3.62 0.00 3.28 0.43 7.1 550.02020 213.0 492.8 67.4 15.8 8.4 70.9 1.5 301.8 30.9 30.9 3.08 0.00 3.29 0.38 7.1 550.02025 197.8 470.3 63.6 15.5 8.4 62.4 0.5 299.8 34.1 32.4 2.55 0.00 3.26 0.32 7.1 550.02030 182.6 448.0 59.9 15.1 8.4 54.4 0.2 265.4 34.4 31.2 2.06 0.00 3.19 0.27 7.1 550.02035 167.7 426.5 56.5 14.7 8.4 47.0 0.1 208.0 31.9 27.8 1.65 0.00 3.09 0.22 7.1 550.02040 153.5 405.8 53.3 14.4 8.4 40.3 0.0 151.0 27.9 23.3 1.30 0.00 2.97 0.18 7.1 550.02045 139.9 386.1 50.2 14.0 8.4 34.4 0.0 105.0 23.5 18.9 1.01 0.00 2.85 0.15 7.1 550.02050 127.2 367.3 47.4 13.6 8.4 29.2 0.0 71.1 19.3 14.9 0.78 0.00 2.71 0.12 7.1 550.02055 115.4 349.4 44.7 13.3 8.4 24.2 0.0 47.5 15.6 11.6 0.60 0.00 2.57 0.09 7.1 550.02060 104.4 332.4 42.1 12.9 8.3 20.0 0.0 31.5 12.4 9.0 0.46 0.00 2.43 0.07 7.1 550.02065 94.4 316.2 39.7 12.6 8.3 16.5 0.0 20.8 9.8 6.8 0.35 0.00 2.29 0.06 7.1 550.02070 85.2 300.7 37.4 12.3 8.3 13.6 0.0 13.7 7.7 5.2 0.26 0.00 2.16 0.05 7.1 550.02075 76.8 286.1 35.3 12.0 8.3 11.2 0.0 9.0 6.1 3.9 0.20 0.00 2.02 0.04 7.1 550.02080 69.1 272.1 33.3 11.6 8.2 9.3 0.0 5.9 4.7 3.0 0.15 0.00 1.90 0.03 7.1 550.02085 62.2 258.8 31.4 11.3 8.2 7.6 0.0 3.9 3.7 2.2 0.11 0.00 1.77 0.02 7.1 550.02090 55.9 246.2 29.6 11.1 8.2 6.3 0.0 2.6 2.9 1.7 0.08 0.00 1.66 0.02 7.1 550.02095 50.2 234.2 27.9 10.8 8.2 5.2 0.0 1.7 2.2 1.3 0.06 0.00 1.55 0.01 7.1 550.02100 45.1 222.8 26.3 10.5 8.1 4.3 0.0 1.1 1.7 0.9 0.05 0.00 1.44 0.01 7.1 550.01Note: Areas are shaded for compounds in years when mixing ratio values are forced to equal global estimates inferred from observations. The pre-1995 global mixing ratio trend for CH Br was derived 3

from SH firn data by including a time-varying hemispheric ratio that increased linearly from 1.0 in 1940 to 1.3 in 1995 (see Section 1.2.1.6 of Chapter 1). This is based on the concurrent decline in the measured hemispheric difference and industrially averaged emissions since 1998, and suggests a mean hemispheric mixing ratio difference close to 0 ppt for CH Br in preindustrial times. Previously, 3a constant hemispheric ratio had been used. CH Cl mixing ratios before 1995 were also derived from South Pole firn data (Butler et al., 1999). From 1995 onward, global means of CH 3 Br are based 3on surface observations, while global means of CH Cl are held constant (Chapter 1). All other shaded areas are from atmospheric observations from one or more measurement networks (Chapter 1). 3
5.56

Chapter 5

Table 5A-4. Halocarbon indirect GWPs from ozone depletion using the EESC-based method described in Daniel et al. (1995). Values have been updated for the new scenarios and fractional release 

values. Indirect forcing from ozone depletion is also now assumed to continue even after EESC drops below 

the 1980 level, consistent with our understanding that this level does not represent a threshold value for 

ozone depletion. Uncertainties in direct GWPs represent 35% of the direct value; uncertainties in the indirect GWPs reflect the uncertainty in stratospheric ozone radiative forcing quoted in IPCC (−0.05 ± 0.10 W/

m2) even though that forcing is not attributed to only halocarbons.

Gas Direct GWP Indirect GWP

CFC-11 4,750 ± 1,660 −1,630 ± 3,250

CFC-12 10,900 ± 3,810 −1,270 ± 2,540

CFC-113 6,130 ± 2,150 −1,310 ± 2,630

CFC-114 9,180 ± 3,210 −562 ± 1,120

CFC-115 7,230 ± 2,530 −137 ± 275

HCFC-22 1,790 ± 630 −58 ± 116

HCFC-123 77 ± 27 −25 ± 50

HCFC-124 619 ± 217 −30 ± 60

HCFC-141b 717 ± 252 −158 ± 316

HCFC-142b 2,220 ± 780 −97 ± 193

HCFC-225ca 122 ± 42 −27 ± 54

HCFC-225cb 606 ± 214 −39 ± 78

CH3CCl3 146 ± 53 −212 ± 424

CCl4 1,400 ± 490 −1,300 ± 2,600

CH3Br 5 ± 2 −851 ± 1,700

Halon-1211 1,890 ± 660 −11,720 ± 23,430

Halon-1301 7,140 ± 2,500 −27,060 ± 54,130

Halon-2402 1,640 ± 570 −19,780 ± 39,570
Twenty Questions and Answers

About the Ozone Layer: 2010 Update

Coordinating Lead Authors:

David W. Fahey

Michaela I. Hegglin

The update of this component of the Assessment was discussed by the 74 scientists who attended the Panel Review Meeting 

for the 2010 Ozone Assessment (Les Diablerets, Switzerland, 28 June–2 July 2010). In addition, subsequent contributions, 

reviews, or comments were provided by the following individuals: Ross J. Salawitch (Special Recognition), Stephen A. 

Montzka (Special Recognition), Stephen O. Andersen (Special Recognition), Pieter J. Aucamp, Alkiviadis F. Bais, Peter F. 

Bernath, Gregory E. Bodeker, Janet F. Bornman, Geir O. Braathen, Peter Braesicke, Irene Cionni, Martin Dameris, John 

S. Daniel, Susana B. Diaz, Ellsworth G. Dutton, James W. Elkins, Christine A. Ennis, Veronika Eyring, Vitali E. Fioletov, 

Marvin A. Geller, Sophie Godin-Beekmann, Malcolm K.W. Ko, Kirstin Krüger, Lambert Kuijpers, Michael J. Kurylo, 

Igor Larin, Gloria L. Manney, C. Thomas McElroy, Rolf Müller, Eric R. Nash, Paul A. Newman, Samuel J. Oltmans, Nigel 

D. Paul, Judith Perlwitz, Jean-Pierre Pommereau, Claire E. Reeves, Stefan Reimann, Alan Robock, Michelle L. Santee, 

Dian J. Seidel, Theodore G. Shepherd, Peter Simmonds, Anne K. Smith, Richard S. Stolarski, Matthew B. Tully, Guus J.M. 

Velders, Elizabeth C. Weatherhead, Ann R. Webb, Ray F. Weiss, and Durwood Zaelke.

Twenty Questions and Answers About the Ozone Layer: 2010 Update

Contents

INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . .................................................................. Q.1

Section I: OZONE IN OUR ATMOSPHERE

Q1: What is ozone and where is it in the atmosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.4

Q2: How is ozone formed in the atmosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.6

Q3: Why do we care about atmospheric ozone? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.8

Q4: How is total ozone distributed over the globe? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.10

Q5: How is ozone measured in the atmosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.12

Section II: THE OZONE DEPLETION PROCESS

Q6: What are the principal steps in stratospheric ozone depletion caused by human activities? . . . . . . . . . . . . . . . . . . Q.14

Q7: What emissions from human activities lead to ozone depletion? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.16

Q8: What are the reactive halogen gases that destroy stratospheric ozone? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.20

Q9: What are the chlorine and bromine reactions that destroy stratospheric ozone? . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.24

Q10: Why has an “ozone hole” appeared over Antarctica when ozone-depleting 

 substances are present throughout the stratosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.27

Section III: STRATOSPHERIC OZONE DEPLETION

Q11: How severe is the depletion of the Antarctic ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.31

Q12: Is there depletion of the Arctic ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.36

Q13: How large is the depletion of the global ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.40

Q14: Do changes in the Sun and volcanic eruptions affect the ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.42

Section IV: CONTROLLING OZONE-DEPLETING SUBSTANCES

Q15: Are there controls on the production of ozone-depleting substances? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.45

Q16: Has the Montreal Protocol been successful in reducing ozone-depleting 

 substances in the atmosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.48

Section V: IMPLICATIONS OF OZONE DEPLETION AND THE MONTREAL PROTOCOL

Q17: Does depletion of the ozone layer increase ground-level ultraviolet radiation? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.52

Q18: Is depletion of the ozone layer the principal cause of climate change? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.55

Q19: Have reductions of ozone-depleting substances under the Montreal Protocol 

 also protected Earth’s climate? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.60

Section VI: STRATOSPHERIC OZONE IN THE FUTURE

Q20: How is ozone expected to change in the coming decades? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.64

Additional Topics

•	 Global Ozone Network . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.13

•	 Understanding Stratospheric Ozone Depletion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.15

•	 Heavier-Than-Air CFCs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.19

•	 Replacing Lost Ozone in the Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.26

•	 The Discovery of the Antarctic Ozone Hole . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.30

•	 The 2002 Antarctic Ozone Hole . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.35

Q.1

Ozone is present only in small amounts in Earth’s atmosphere. Nevertheless, it is vital to human wellbeing and ecosystem health.

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.

It was discovered in the mid-1970s that some humanproduced chemicals could lead to depletion of the ozone 

layer. The resulting increase in ultraviolet radiation at Earth’s 

surface would likely increase the incidences of skin cancer 

and eye cataracts, and also adversely affect plants, crops, and 

ocean plankton.

Following the discovery of this environmental issue, 

researchers sought a better understanding of this threat to the 

ozone layer. Monitoring stations showed that the abundances 

of the ozone-depleting substances (ODSs) 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 in the atmosphere characterized the chemical 

reactions that were involved in ozone destruction. Computer 

models of the atmosphere employing this information were 

used to predict how much ozone depletion was occurring and 

how much more might 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 and southern midlatitudes.

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 indeed occurring 

and why. Most important, we know that if the most potent 

ODSs were to continue to be emitted and increase 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 an international 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 “ozone-friendly” 

substitutes for the now-controlled chemicals, the total global 

accumulation of ODSs has slowed and begun to decrease. In 

response, global ozone depletion is no longer increasing. Now, 

with continued compliance, we expect substantial recovery of 

the ozone layer by the late 21st century. The day the Montreal 

Protocol was agreed upon, 16 September, is now celebrated as 

the International Day for the Preservation of the Ozone Layer.

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, in which continuous dialogue has 

been a key ingredient. A timeline of milestones associated with 

stratospheric ozone depletion is illustrated in Figure Q0-1. The 

milestones relate to stratospheric ozone science, international 

scientific assessments, and the Montreal Protocol.

To help maintain a broad understanding of the relationship between ozone depletion, ODSs, and the Montreal Protocol, this component of the Scientific Assessment of Ozone 

Depletion: 2010 presents 20 questions and answers about the 

often-complex science of ozone depletion. Most questions 

and answers are updates of those presented in previous Ozone 

Assessments, while others have been added or expanded to 

address newly emerging issues. The questions address the 

nature of atmospheric ozone, the chemicals that cause ozone 

depletion, how global and polar ozone depletion occur, the 

success of the Montreal Protocol, and what could lie ahead for 

the ozone layer. Computer models project that the influence 

on global ozone of greenhouse gases and changes in climate 

will grow significantly in the coming decades and exceed 

the importance of ODSs in most atmospheric regions by the 

end of this century. Ozone and climate are indirectly linked 

because both ODSs and their substitutes contribute to climate 

change. 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 2010 and earlier Assessment 

reports as well as other international scientific assessments. 

These reports and the answers provided here were prepared 

and reviewed by a large international group of scientists1.

INTRODUCTION
Q.2

 20 Questions: 2010 Update Introduction

1960 1970 1980 1990 2000 2010 2020

Year

Emissions of halogen source gases (ODSs + natural sources) 

(Weighted by Ozone Depletion Potential (ODP))

0

19891991 1994 1998 2002 2006 2010

UNEP/WMO Scientific Assessments of Ozone Depletion

Vienna Convention for the

Protection of the Ozone Layer – 1985

Montreal Protocol on Substances

that Deplete the Ozone Layer – 1987

Montreal Protocol

Amendments

WMO Atmospheric

Ozone– 1985

WMO World Plan of Action

on the Ozone Layer–1977

International Scientific

Assessments

Montreal Protocol

Milestones

Scientific Milestones

Milestones in the History of Stratospheric Ozone Depletion 

2.0

0.5

1.0

1.5

(Megatonnes CFC-11-equivalent per year)

London – 1990

Copenhagen – 1992

Vienna –1995

Montreal – 1997

Beijing – 1999 Global production

of CFCs and halons

ends – 2010

2005 – IPCC/TEAP Special

Report on Safeguarding the

Ozone Layer and the

Global Climate System

Based on atmospheric

observations or industry data

Future projections

WMO The Stratosphere 1981:

Theory and Measurements

1995 Nobel Prize in Chemistry

to Crutzen, Molina, and Rowland 

Stratospheric abundance of EESC

in midlatitudes down 10%

from 1997 peak – 2008

Ozone hole reported 

over Antarctica – 1985 

CFCs observed in the

stratosphere & reported as

potent greenhouse gases – 1975

CFC ozone-depletion 

hypothesis – 1974 

Global surface observations

of CFCs reported – 1973

WMO World Ozone

Data Centre – 1961

Total Ozone Mapping

Spectrometer (TOMS) 

instrument launched –1978 

Peak stratospheric abundance of 

EESC in midlatitudes –1997 

Peak global emissions

of ozone-depleting 

substances– 1988 

Stratospheric Aerosol &

Gas Experiment I (SAGE-I)

instrument launched – 1979 Laboratory, field, and modeling studies

show chlorine and bromine deplete

polar stratospheric ozone – 1985–1989 

WMO Report of the International

Ozone Trends Panel – 1988

Montreal Protocol 

universal ratification 

(196 nations) –2009

Global consumption of ODSs

reduced by 98%– 1986 to 2008

Natural emissions

HCFC accelerated phaseout 

(Montreal – 2007 Adjustment)

WMO: World Meteorological Organization 

UNEP: United Nations Environment 

 Programme 

IPCC: Intergovernmental Panel

 on Climate Change 

TEAP: Technical and Economic Assessment

 Panel of the Montreal Protocol 

EESC: Equivalent Effective 

Stratospheric Chlorine 

ODSs: Ozone-depleting substances
Q.3

Introduction 20 Questions: 2010 Update

Figure Q0-1. Stratospheric ozone depletion milestones. This timeline highlights milestones related to the history of 

ozone depletion. Events represent the occurrence of important scientific findings, the completion of international scientific 

assessments, and highlights of the Montreal Protocol. The graph shows the history and near future of annual total emissions 

of ozone-depleting substances (ODSs) combined with natural emissions of halogen source gases. ODSs are halogen source 

gases controlled under the Montreal Protocol. The emissions, when weighted by their potential to destroy ozone, peaked 

near 1990 after several decades of steady increases (see Q19). Between 1990 and the present, emissions have decreased 

substantially as a result of the Montreal Protocol and its subsequent Amendments and Adjustments coming into force (see 

Q15). The Protocol began with the Vienna Convention for the Protection of the Ozone Layer in 1985. The provisions of the 

Protocol and its Amendments and Adjustments decisions have depended on information embodied in international scientific 

assessments of ozone depletion that have been produced periodically since 1989 under the auspices of UNEP and WMO. 

Atmospheric observations of ozone, CFCs, and other ODSs have increased substantially since the early 1970s. For example, 

the SAGE and TOMS satellite instruments have provided essential global views of stratospheric ozone for several decades. 

The Nobel Prize in Chemistry in 1995 was awarded for research that identified the threat to ozone posed by CFCs and that 

described key reactive processes in the stratosphere. By 2008, stratospheric chlorine abundances in the stratosphere were 

10% lower than their peak values reached in the late 1990s and were continuing to decrease. January 2010 marked the end of 

global production of CFCs and halons under the Protocol. (A megatonne = 1 billion (109) kilograms.)

 IPCC: Intergovernmental Panel on Climate Change

 TEAP: Technology and Economic Assessment Panel 

 of the Montreal Protocol

 UNEP: United Nations Environment Programme

1

 The update of this component of the Assessment was discussed by the 74 scientists who attended the Panel Review Meeting for the 2010 Ozone 

Assessment (Les Diablerets, Switzerland, 28 June–2 July 2010). In addition, subsequent contributions, reviews, or comments were provided by 

the following individuals: Ross J. Salawitch (Special Recognition), Stephen A. Montzka (Special Recognition), Stephen O. Andersen (Special 

Recognition), Pieter J. Aucamp, Alkiviadis F. Bais, Peter F. Bernath, Gregory E. Bodeker, Janet F. Bornman, Geir O. Braathen, Peter Braesicke, 

Irene Cionni, Martin Dameris, John S. Daniel, Susana B. Diaz, Ellsworth G. Dutton, James W. Elkins, Christine A. Ennis, Veronika Eyring, 

Vitali E. Fioletov, Marvin A. Geller, Sophie Godin-Beekmann, Malcolm K.W. Ko, Kirstin Krüger, Lambert Kuijpers, Michael J. Kurylo, Igor 

Larin, Gloria L. Manney, C. Thomas McElroy, Rolf Müller, Eric R. Nash, Paul A. Newman, Samuel J. Oltmans, Nigel D. Paul, Judith Perlwitz, 

Jean-Pierre Pommereau, Claire E. Reeves, Stefan Reimann, Alan Robock, Michelle L. Santee, Dian J. Seidel, Theodore G. Shepherd, Peter Simmonds, Anne K. Smith, Richard S. Stolarski, Matthew B. Tully, Guus J.M. Velders, Elizabeth C. Weatherhead, Ann R. Webb, Ray F. Weiss, and 

Durwood Zaelke.

EESC: Equivalent effective stratospheric chlorine

ODS: Ozone-depleting substance

WMO: World Meteorological Organization
Q.4

Section I: OZONE IN OUR ATMOSPHERE

Q1

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 

(about 90%) resides in the stratosphere 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.”

What is ozone and where is it in the atmosphere? 

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 mid1800s. 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), 

meaning “to smell.” Ozone has a pungent odor that allows it 

to be detected even at very low amounts. Ozone reacts rapidly 

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 of textiles and food products.

Ozone location. Most ozone (about 90%) is found in 

the stratosphere, which begins about 10–16 kilometers (6–10 

miles) above Earth’s surface and extends up to about 50 kilometers (31 miles) altitude. The stratospheric region with 

the highest ozone concentration is commonly known as the 

“ozone layer” (see Figure Q1-2). The ozone layer extends over 

the entire globe with some variation in altitude and thickness. 

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 low relative 

abundance in the atmosphere. In the stratosphere near the 

peak concentration of the ozone layer, there are typically a few 

thousand 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 result when ozone is 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 layer of gas extending over the globe. The resulting 

layer of pure ozone would have an average thickness of about 

three millimeters (about one-tenth inch) (see Q4). Nonetheless, this extremely small fraction of the atmosphere plays a 

vital role in protecting life on Earth (see Q3).

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 the 

gases in Earth’s atmosphere, contain two oxygen atoms 

bound together.

Oxygen

atom (O)

Oxygen

molecule (O2)

Ozone

molecule (O3)

Ozone and Oxygen
Q.5

Section I: OZONE IN OUR ATMOSPHERE 20 Questions: 2010 Update

Figure Q1-2. Ozone in the atmosphere. Ozone is present throughout the troposphere and stratosphere. This profile shows schematically how ozone changes with 

altitude in the tropics. Most ozone 

resides in the stratospheric “ozone 

layer.” The vertical extent or thickness of this layer varies from region 

to region and with season over the 

globe (see Q4). Increases in ozone 

occur near the surface as a result of 

pollution from human activities.

Altitude (kilometers)

Ozone abundance (mPa)

Altitude (miles)

0

5

10

15

20

25

30

35

Stratospheric

ozone

Tropospheric

ozone

20

15

10

5

Ozone Layer

Ozone

increases

from pollution

Ozone in the Atmosphere 
Q.6

 20 Questions: 2010 Update Section I: OZONE IN OUR ATMOSPHERE

Stratospheric ozone. Stratospheric ozone is formed 

naturally by chemical reactions involving solar ultraviolet 

radiation (sunlight) and oxygen molecules, which make up 

21% of the atmosphere. In the first step, solar ultraviolet 

radiation breaks apart one oxygen molecule (O2) to produce 

two oxygen atoms (2 O) (see Figure Q2-1). In the second step, 

each of these highly reactive atoms combines with an oxygen 

molecule to produce an ozone molecule (O3). These reactions 

occur continually whenever solar ultraviolet radiation is present in the stratosphere. As a result, the largest ozone production occurs in the tropical stratosphere.

The production of stratospheric ozone is balanced by its 

destruction in chemical reactions. Ozone reacts continually with sunlight and 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 

hydrogen and nitrogen oxides and those containing chlorine 

and bromine (see Q8).

Some stratospheric ozone is regularly transported down 

into the troposphere and can occasionally influence ozone 

amounts at Earth’s surface, particularly in remote, unpolluted 

regions of the globe.

Tropospheric ozone. Near Earth’s surface, ozone is 

produced by chemical reactions involving naturally occurring gases and gases from pollution sources. Ozone production reactions primarily involve hydrocarbon and nitrogen 

oxide gases, as well as ozone itself, and all require sunlight 

for completion. Fossil fuel combustion is a primary source 

of pollutant gases that lead to tropospheric ozone production. The production of ozone near the surface 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 by naturally occurring chemical reactions and 

by reactions involving human-produced chemicals. Tropospheric ozone can also be destroyed when ozone reacts with a 

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 by a different set of chemical reactions that involve naturally occurring gases and 

those from pollution sources.

Q2 How is ozone formed in the atmosphere? 

Figure Q2-1. Stratospheric ozone production.

Ozone is naturally produced in the stratosphere by a twostep reactive process. In the first step, solar ultraviolet 

radiation (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.

Stratospheric Ozone Production

Overall reaction: 3 O2 2 O3

sunlight
Q.7

 Section I: OZONE IN OUR ATMOSPHERE 20 Questions: 2010 Update

variety of surfaces, such as those of soils and plants.

Balance of chemical processes. Ozone abundances 

in the stratosphere and troposphere are determined by the 

balance between chemical processes that produce and destroy 

ozone. The balance is determined by the amounts of reactive 

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 increase. Similarly, if conditions 

change to favor other reactions that destroy ozone, abundances decrease. The balance of production and loss reactions combined with atmospheric air motions determines the 

global distribution of ozone on timescales of days to many 

months. Global ozone has decreased during the past several 

decades because the amounts of reactive gases containing 

chlorine and bromine have increased in the stratosphere (see 

Q13) due to human activities.
Q.8

 20 Questions: 2010 Update Section I: OZONE IN OUR ATMOSPHERE

Ozone in the stratosphere absorbs a large part of the Sun’s biologically harmful ultraviolet radiation. Stratospheric ozone 

is considered “good” ozone because of this beneficial role. In contrast, ozone formed at Earth’s surface in excess of natural 

amounts is considered “bad” ozone because it is harmful to humans, plants, and animals. Natural ozone near the surface 

and in the lower atmosphere plays an important beneficial role in chemically removing pollutants from the atmosphere.

Q3 Why do we care about atmospheric ozone?

Good ozone. Stratospheric ozone is considered good 

for humans and other life forms because it absorbs ultraviolet-B (UV-B) radiation from the Sun (see Figure Q3-1). If 

not absorbed, UV-B radiation would reach Earth’s surface 

in amounts that are harmful to a variety of life forms. In 

humans, increased exposure to UV-B radiation increases 

the risks of skin cancer, cataracts, and a suppressed immune 

system. UV-B radiation exposure before adulthood and 

cumulative exposure are both important health risk factors. 

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.

Protecting good ozone. In the mid-1970s, it was discovered that gases containing chlorine and bromine atoms 

released by human activities could cause stratospheric ozone 

depletion (see Q6). These gases, referred to as halogen source 

gases, and as ozone-depleting substances (ODSs), chemically 

release their chlorine and bromine atoms after they reach the 

stratosphere. Ozone depletion increases surface UV-B radiation above naturally occurring amounts. International efforts 

have been successful in protecting the ozone layer through controls on ODS production and consumption (see Q15 and Q16).

Bad ozone. Ozone near Earth’s surface in excess of natural amounts is considered bad ozone. It is formed by reactions 

involving human-made pollutant gases. Increasing surface 

ozone above natural levels is harmful to humans, plants, and 

other living systems because ozone reacts strongly to destroy 

or alter many biological molecules. High ozone exposure 

reduces crop yields and forest growth. In humans, expoFigure Q3-1. UV 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 strongly 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), visible 

light, and other solar radiation are not strongly 

absorbed by the ozone layer. Human exposure 

to UV-B radiation increases the risks of skin 

cancer, cataracts, and a suppressed immune 

system. UV-B radiation exposure can also 

damage terrestrial plant life, single-cell organisms, and aquatic ecosystems.

UV Protection by the Ozone Layer
Q.9

sure 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 because ozone is a greenhouse gas (see Q18). 

The negative effects of excess tropospheric ozone contrast 

sharply with the protection from harmful UV-B radiation 

afforded by an abundance of stratospheric ozone.

Reducing bad ozone. Limiting the emission of certain 

common pollutants reduces the production of excess ozone 

in the air surrounding humans, plants, and animals. Natural 

emissions from the biosphere, mainly from trees, also participate in reactions that produce ozone. 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 or limiting the 

emission of pollutants that cause production of excess ozone 

near Earth’s surface.

Natural ozone. In the absence of human activities, ozone 

would still be present near Earth’s surface and throughout the 

troposphere and stratosphere because ozone is a natural component of the clean atmosphere. Ozone plays important roles 

in the atmosphere beyond absorbing UV radiation. For example, ozone initiates the chemical removal of many pollutants, 

such as carbon monoxide (CO) and nitrogen oxides (NOx), 

as well as some greenhouse gases, such as methane (CH4). In 

addition, the absorption of UV-B radiation by ozone is a natural source of heat in the stratosphere, causing temperatures to 

increase with altitude. Stratospheric temperatures affect the 

balance of ozone production and destruction processes (see 

Q2) and air motions that redistribute ozone throughout the 

stratosphere.

Section I: OZONE IN OUR ATMOSPHERE 20 Questions: 2010 Update
Q.10

Total ozone. Total ozone at any location on the globe is 

defined as the sum of all the ozone in the atmosphere directly 

above that location. Most ozone resides in the stratospheric 

ozone layer and a small percentage (about 10%) is distributed 

throughout the troposphere (see Q1). Total ozone values 

are often reported in Dobson units denoted as “DU.” Typical values vary between 200 and 500 DU over the globe (see 

Figure Q4-1). The ozone molecules required for total ozone 

to be 500 DU around the globe, for example, could also form 

a layer of pure ozone gas at Earth’s surface having a thickness 

of only 5 millimeters (0.2 inches) (see Q1).

Global distribution. Total ozone varies strongly with 

latitude over the globe, with the largest values occurring at 

middle and high latitudes during all seasons (see Figure Q4-1). 

This is the result of ozone production rates from solar ultraviolet radiation that are highest on average in the tropics, and 

the large-scale air circulation in the stratosphere that slowly 

transports tropical ozone toward the poles. Ozone accumulates at middle and high latitudes, increasing the thickness (or 

vertical extent) of the ozone layer and, at the same time, total 

ozone. In contrast, the values of total ozone are the lowest in 

the tropics in all seasons (except in the ozone hole) because 

the thickness of the ozone layer is smallest there. 

Seasonal distribution. Total ozone also varies with 

season, as is shown in Figure Q4-1 using two-week averages 

of ozone taken from 2009 satellite observations. March and 

September plots represent the early spring and fall seasons in 

the Northern and Southern Hemispheres. June and December plots similarly represent the early summer and winter seasons. Total ozone shows a maximum at high latitudes during 

spring as a result of increased transport of ozone from its 

source region in the tropics toward the polar regions during 

late fall and winter. This ozone transport is much weaker 

during the summer and early fall periods and is weaker overall in the Southern Hemisphere. An important feature of 

seasonal ozone changes is the natural chemical destruction 

that occurs when daylight is continuous in the summer polar 

stratosphere, which causes total ozone to decrease gradually 

toward its lowest values in early fall.

This natural seasonal cycle can be observed clearly in the 

Northern Hemisphere as shown in Figure Q4-1, with increasing values in Arctic total ozone during winter, a clear maximum in spring, and decreasing values from summer to fall. 

In the Antarctic, however, a pronounced minimum in total 

ozone is observed during spring. The minimum is a consequence of the “ozone hole,” which describes the widespread 

chemical destruction of ozone by ozone-depleting substances 

(see Q6 and Q11) in spring. In the late 1970s, before the ozone 

hole appeared each year, much higher ozone values were 

found in Antarctic spring (see Q11). Now, the lowest values 

of total ozone across the globe and all seasons are found every 

spring in the Antarctic as shown in Figure Q4-1. After spring, 

these low values disappear from total ozone maps as polar air 

mixes with lower-latitude air containing much higher ozone 

values.

In the tropics, the total ozone changes through the springsummer-fall-winter progression of the seasons are much 

smaller than in the polar regions. This is because seasonal 

changes in both sunlight and ozone transport are smaller in 

the tropics than in the polar regions.

Natural variations. Total ozone varies strongly with 

latitude and longitude within the seasonal plots in Figure 

Q4-1. These patterns, which change on daily to weekly 

timescales, come about for two reasons. First, natural air 

motions mix and blend air between regions of the stratosphere that have high ozone values and those that have low 

ozone values. Tropospheric weather systems can temporarily change the thickness of the ozone layer in a region, and 

thereby change total ozone. The geographical variation in 

these air motions in turn causes variations in the distribution of total ozone.

Second, ozone variations occur as a result of changes in 

the balance of chemical production and loss processes as air 

moves to and from different locations over the globe. This 

balance, for example, is very sensitive to the amount of sunlight in a region.

The distribution of total ozone over the Earth varies with location on timescales that range from daily to seasonal. The variations are caused by large-scale movements of stratospheric air and the chemical production and destruction of ozone. Total 

ozone is generally lowest at the equator and highest in polar regions.

Q4 How is total ozone distributed over the globe? 

 20 Questions: 2010 Update Section I: OZONE IN OUR ATMOSPHERE
Q.11

There is 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 routinely monitored by a large group of investigators using satellite, airborne, and ground-based instruments. 

The continued analysis of these observations provides an 

important basis to quantify the contribution of human activities to ozone depletion.

Section I: OZONE IN OUR ATMOSPHERE 20 Questions: 2010 Update

Figure Q4-1. Total ozone. Total ozone at any location on the globe is defined as the sum of all the ozone molecules in 

the atmosphere directly above that location. Total ozone varies with latitude, longitude, and season, with the largest values 

at high latitudes and the lowest values in tropical regions. The variations are demonstrated here with two-week averages 

of total ozone in 2009 as measured with a satellite instrument. Total ozone shows little variation in the tropics (20°N–20°S 

latitudes) over all seasons. Total ozone outside the tropics varies more strongly with time on a daily to seasonal basis as 

ozone-rich air is moved from the tropics and accumulates at higher latitudes. The low total ozone values over Antarctica in 

September constitute the “ozone hole” in 2009. Since the 1980s, the ozone hole in late winter and early spring represents the 

lowest values of total ozone that occur over all seasons and latitudes (see Q11).

16 – 30 December 16 – 30 June

16 – 30 March

16 – 30 September

Global Satellite Maps of Total Ozone in 2009

Early spring 

Early fall 

Early spring

Early fall 

Early summer

Early winter Early summer

Early winter

100

200

300

400

500

600

Total ozon

e

(Dobson units)
Q.12

 20 Questions: 2010 Update Section I: OZONE IN OUR ATMOSPHERE

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’s detection chamber, ozone is measured by its absorption of ultraviolet (UV) light or by the electrical current or 

light produced in a chemical reaction involving ozone. The 

last 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 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. 

Ozone measurements are also being made routinely on some 

commercial aircraft flights.

Remote measurements. Remote measurements of 

total ozone amounts and the altitude distributions of ozone 

are obtained by detecting ozone at large distances 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, lasers, and starlight. For example, satellites use the absorption of solar UV radiation 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. Lasers are routinely deployed at ground 

sites and on research aircraft to detect ozone over a distance 

of many kilometers along the laser light path. A network of 

ground-based detectors measures ozone by detecting small 

changes in the amount of the Sun’s UV radiation that reaches 

Earth’s surface. Other instruments measure ozone using its 

absorption of infrared or visible radiation or its emission of 

The amount of ozone in the atmosphere is measured by instruments on the ground and carried aloft on balloons, 

aircraft, and satellites. Some instruments measure ozone locally by continuously drawing air samples into a small 

detection chamber. Other instruments measure ozone remotely over long distances by using ozone’s unique optical 

absorption or emission properties.

Q5 How is ozone measured in the atmosphere? 

Figure Q5-1. Ozone measurements. Ozone is measured throughout the atmosphere with instruments on 

the ground, 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; detect the microwave emissions from ozone; or use chemical reactions 

that are unique to ozone. At many locations over the 

globe, regular measurements are made to monitor total 

ozone amounts and their variations over time.

Measuring Ozone in the Atmosphere

High-altitude

aircraft

Low-altitude

aircraft

Laser

beams

Ground-based

systems

Balloon

sondes

Satellites
Q.13

Section I: OZONE IN OUR ATMOSPHERE 20 Questions: 2010 Update

microwave or infrared radiation. Emission measurements 

have the advantage of providing remote ozone measurements 

at night, which is particularly valuable for sampling polar 

regions in continuous darkness.

Global Ozone Network

The first instrument for routinely monitoring total ozone was developed by Gordon M.B. Dobson in the United Kingdom 

in the 1920s. The instrument, called a Féry spectrometer, made its measurements by examining the wavelength spectrum of 

solar ultraviolet radiation (sunlight) using a photographic plate. A small network of instruments distributed around Europe 

allowed Dobson to make important discoveries about how total ozone varies with location and time. In the 1930s a new 

instrument was developed by Dobson, now called a Dobson spectrophotometer, which precisely 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 provides a measure of total ozone above the instrument location.

A global network of total-ozone observing stations was established in 1957 as part of the International Geophysical 

Year. Today, there are about 100 sites located around the world ranging from South Pole, Antarctica (90°S), to Ellesmere 

Island, Canada (83°N), that routinely measure total ozone. The accuracy of these observations is maintained by regular 

instrument calibrations and intercomparisons. Data from the network have been essential for understanding the effects of 

chlorofluorocarbons (CFCs) and other ozone-depleting substances on the global ozone layer, starting before the launch of 

space-based ozone-measuring instruments and continuing to the present day. Ground-based instruments with excellent 

long-term stability and accuracy 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).
Q.14

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, often referred to as 

ozone-depleting substances (ODSs), include manufactured chemicals released to the atmosphere in a variety of applications, such as 

refrigeration, air conditioning, and foam blowing. Chlorofluorocarbons (CFCs) are an important example of chlorine-containing 

gases. Emitted source gases accumulate in the lower atmosphere 

(troposphere) and are transported to the stratosphere by natural air 

motions. The accumulation occurs because most source gases are 

highly unreactive in the lower atmosphere. Small amounts of these 

gases dissolve in ocean waters. The low reactivity of these manufactured halogenated gases is one property that makes them well suited 

for specialized applications such as refrigeration.

Some halogen gases are emitted in substantial quantities from 

natural sources (see Q7). These emissions also accumulate in the 

troposphere, are transported to the stratosphere, and participate in 

ozone destruction reactions. These naturally emitted gases are part 

of the natural balance of ozone production and destruction that 

predates the large release of manufactured halogenated gases.

The initial step in the depletion of stratospheric ozone by human activities is the emission, at Earth’s surface, of 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. Natural air motions transport these accumulated gases to the stratosphere, where 

they are converted to more reactive gases. Some of these 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.

What are the principal steps in stratospheric ozone depletion Q6 caused by human activities? 

Section II: THE OZONE DEPLETION PROCESS

Principal Steps in the Depletion

of Stratospheric Ozone

Halogen source gases are emitted at Earth's 

surface by human activities and natural processes.

Halogen source gases accumulate in the

atmosphere and are globally distributed throughout

the lower atmosphere by winds and other air motions.

Halogen source gases are transported to the 

stratosphere by air motions.

Most halogen source gases are converted in the

stratosphere to reactive halogen gases in chemical 

reactions involving ultraviolet radiation from the Sun.

Air containing reactive halogen gases returns

to the troposphere where the gases are removed 

by moisture in clouds and rain.

Reactive halogen gases cause chemical depletion

of stratospheric ozone over the globe.

Emissions

Accumulation

Transport

Chemical reaction

Conversion

Removal

Low-temperature surface reactions on polar

stratospheric clouds (PSCs) significantly 

increase reactive halogen gases and 

thereby cause severe ozone loss in polar 

regions in late winter and early spring. 

Figure Q6-1. Principal steps in stratospheric ozone depletion. The stratospheric ozone depletion process begins with 

the emission of halogen source gases by human activities and 

natural processes. Those emitted by human activities are also 

called ozone-depleting substances (ODSs). Subsequent steps 

are accumulation, transport, conversion, chemical reaction, and 

removal. Ozone depletion by halogen source gases occurs globally. Large seasonal ozone losses occur in polar regions as a result 

of reactions involving polar stratospheric clouds (PSCs). Ozone 

depletion ends when reactive halogen gases are removed by rain 

and snow in the troposphere and deposited on Earth’s surface.
Q.15

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). Gases with longer lifetimes have 

slower conversion rates and survive longer in the atmosphere 

after emission. Lifetimes of the principal ODSs vary from 1 to 

100 years (see Q7). Emitted gas molecules with atmospheric 

lifetimes greater than a few years circulate between the troposphere and stratosphere multiple times, on average, before 

conversion occurs.

The reactive gases formed from halogen source gases react

chemically to destroy ozone in the stratosphere (see Q9). The 

average depletion of total ozone attributed to reactive gases is 

smallest in the tropics and largest at high latitudes (see Q13). 

In polar regions, surface reactions that occur at low temperatures on polar stratospheric clouds (PSCs) greatly increase 

the abundance of the most reactive chlorine gas, chlorine 

monoxide (ClO) (see Q10). This results in substantial ozone 

destruction in polar regions in late winter and early 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 or deposited on Earth’s land or ocean surfaces. 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 (less than 1 year) undergo significant chemical conversion in the troposphere, producing reactive halogen gases and other compounds. Source gas molecules that 

are not converted are transported to the stratosphere. Only 

small portions of reactive halogen gases produced in the troposphere are transported to the stratosphere because most 

are removed by precipitation. Important examples of halogen 

gases that undergo some tropospheric removal are the hydrochlorofluorocarbons (HCFCs), methyl bromide (CH3Br), and 

gases containing iodine (see Q7).

Section II: THE OZONE DEPLETION PROCESS 20 Questions: 2010 Update

Understanding Stratospheric Ozone Depletion 

Our understanding of stratospheric ozone depletion has been obtained through a combination of laboratory studies, 

computer models, and atmospheric observations. The wide variety of chemical reactions that occur in the stratosphere 

have been discovered and studied in laboratory studies. Chemical reactions between two gases follow well-defined physical 

rules. Some of these reactions occur on the surfaces of polar stratospheric clouds (PSCs) formed in the winter stratosphere. 

Reactions have been studied that involve many different molecules containing chlorine, bromine, fluorine, and iodine and 

other atmospheric constituents such as carbon, oxygen, nitrogen, and hydrogen. These studies have shown that several 

reactions involving chlorine and bromine directly or indirectly destroy ozone in the stratosphere.

Computer models have been used to examine the combined effect of the large group of known reactions that occur in 

the stratosphere. These models simulate the stratosphere by including representative chemical abundances, winds, air 

temperatures, and the daily and seasonal changes in sunlight. These analyses show that under certain conditions chlorine 

and bromine react in catalytic cycles in which one chlorine or bromine atom destroys many thousands of ozone molecules. 

Models are also used to simulate ozone amounts observed in previous years as a strong test of our understanding of atmospheric processes and to evaluate the importance of new reactions found in laboratory studies. The responses of ozone 

to possible future changes in the abundances of trace gases, temperatures, and other atmospheric parameters have been 

extensively explored with specialized computer models (see Q20).

Atmospheric observations have shown what gases are present in different regions of the stratosphere and how their 

abundances vary. Gas and particle abundances have been monitored over time periods spanning a daily cycle to decades. 

Observations show that halogen source gases and reactive halogen gases are present in the stratosphere at the amounts 

required to cause observed ozone depletion. Ozone and chlorine monoxide (ClO), for example, have been observed extensively 

with a variety of instruments. ClO is a highly reactive gas that is involved in catalytic ozone destruction cycles throughout the 

stratosphere (see Q9). Instruments on the ground and on satellites, balloons, and aircraft now routinely detect ozone and ClO 

remotely using optical and microwave signals. High-altitude aircraft and balloon instruments are also used to detect both gases 

locally in the stratosphere (see Q5). The observations of ozone and reactive gases made in past decades are used extensively 

in comparisons with computer models in order to increase confidence in our understanding of stratospheric ozone depletion.
Q.16

Halogen source gases versus ODSs. Those halogen 

source gases emitted by human activities and controlled by 

the Montreal Protocol are referred to as ODSs within the Montreal Protocol, by the media, and in the scientific literature. 

The Montreal Protocol now controls the global production 

and consumption of ODSs (see Q15). Halogen source gases 

that have only natural sources are not classified as ODSs. The 

contributions of ODSs and natural halogen source gases to 

chlorine and bromine entering the stratosphere in 2008 are 

shown in Figure Q7-1.

Ozone-depleting substances (ODSs). ODSs are manufactured for specific industrial uses or consumer products, 

most of which result in the eventual emission of these gases to 

the atmosphere. ODS emissions have increased substantially 

from the middle to the late 20th century, reached a peak in the 

late 1980s, and are now in decline (see Figure Q0-1). A large 

fraction of ODS emissions reach the stratosphere and lead to 

ozone depletion because chlorine and bromine atoms react to 

destroy ozone. ODSs 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 halogen source gases emitted 

by human activities. These and other chlorine-containing 

ODSs have been used in many applications, including refrigeration, air conditioning, foam blowing, aerosol propellants, 

and cleaning of metals and electronic components.

Another category of ODSs contains bromine. The most 

important of these gases are the halons and methyl bromide 

(CH3Br). Halons are halocarbon gases originally developed 

to extinguish fires. Halons were widely used to protect large 

computer installations, military hardware, and commercial 

aircraft engines. As a consequence, halons are often released 

directly into the atmosphere upon use. Halon-1211 and 

halon-1301 are the most abundant halons emitted by human 

activities (see Figure Q7-1). Methyl bromide is used primarily 

as an agricultural and pre-shipping fumigant.

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 contributed about 17% of the chlorine in the stratosphere in 2008 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 biological activity in the oceans. Only a 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 28%, but this has a large uncertainty. 

The contribution to stratospheric chlorine of short-lived chlorinated gases from natural and human sources is much smaller 

(less than 3%) and is included in Other gases in Figure Q7-1. 

The amounts of chlorine and bromine in the stratosphere from 

natural sources are believed to have been fairly constant since 

the middle of the 20th century and, therefore, cannot be the 

cause of ozone depletion as observed since the 1980s.

Other human sources of chlorine and bromine. 

Other chlorine- and bromine-containing gases are released 

regularly from human activities. Common examples are 

the use of chlorine gases to disinfect swimming pools and 

wastewater, fossil fuel burning, biomass burning, and various 

industrial processes. These emissions do not contribute significantly to stratospheric amounts of chlorine and bromine 

because either the global source is small, or the emitted gases 

and their degradation products are short-lived (very reactive 

or highly soluble). As a consequence, the chlorine and bromine content of these gases is prevented from reaching the 

stratosphere in significant amounts.

Certain industrial processes and consumer products result in the emission of ozone-depleting substances (ODSs) to the 

atmosphere. ODSs are manufactured halogen source gases that are controlled worldwide by the Montreal Protocol. These 

gases bring chlorine and bromine atoms to the stratosphere, where they destroy ozone in chemical reactions. Important 

examples are the chlorofluorocarbons (CFCs), once used in almost all refrigeration and air conditioning systems, and the 

halons, which were used in fire extinguishers. Current ODS abundances in the atmosphere are known directly from air 

sample measurements.

 20 Questions: 2010 Update Section II: THE OZONE DEPLETION PROCESS

Q7 What emissions from human activities lead to ozone depletion?
Q.17

Figure Q7-1. Stratospheric source gases. A variety of halogen source gases emitted from natural sources and by human 

activities transport chlorine and bromine into the stratosphere. Ozone-depleting substances (ODSs) are the subset of these 

gases emitted by human activities that are controlled by the Montreal Protocol. These partitioned columns show the sources 

and abundances of chlorine- and bromine-containing gases entering the stratosphere in 2008. The approximate amounts 

are derived from tropospheric observations of each gas in 2008. Note the large difference in the vertical scales: total chlorine 

entering the stratosphere is 150 times more abundant than total bromine. Human activities are the largest source of chlorine 

reaching the stratosphere and the CFCs are the most abundant chlorine-containing gases. Methyl chloride is the primary 

natural source of chlorine. HCFCs, which are substitute gases for CFCs and also controlled under the Montreal Protocol, are a 

small but growing fraction of chlorine-containing gases. For bromine entering the stratosphere, halons and methyl bromide 

are the largest contributors. Methyl bromide has an additional, much larger, natural source. Natural sources provide a much 

larger fraction of total bromine entering the stratosphere 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 equals the presence of one molecule of a gas per trillion 

(=1012) total air molecules.)

Lifetimes and emissions. After emission, halogen source 

gases are either naturally removed from the atmosphere or 

undergo chemical conversion in the troposphere or stratosphere. The time to remove or convert about 60% of a gas 

is often called its atmospheric lifetime. Lifetimes vary from 

less than 1 year to 100 years for the principal chlorine- and 

bromine-containing gases (see Table Q7-1). The long-lived 

gases are primarily destroyed in the stratosphere and essentially all of the emitted halogen is available to participate in 

the destruction of stratospheric ozone. Gases with the short 

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 the emitted 

halogen contributes to ozone depletion in the stratosphere.

The amount of an emitted gas that is present in the atmosphere represents a balance between the emission rate and the 

lifetime of the gas. Emission rates and atmospheric lifetimes 

vary greatly for the source gases, as indicated in Table Q7-1. 

For example, the atmospheric abundances of most of the 

principal CFCs and halons have decreased since 1990 while 

those of the leading substitute gases, the hydrochlorofluorocarbons (HCFCs), continue to increase under the provisions 

of the Montreal Protocol (see Q16). In the coming decades, 

the emissions and atmospheric abundances of all controlled 

gases are expected to decrease under these provisions.

Ozone Depletion Potential (ODP). Halogen source 

gases are compared in their effectiveness to destroy stratospheric ozone using the ODP, as listed in Table Q7-1 (see Q18). 

A gas with a larger ODP destroys more ozone over its atmospheric lifetime. The ODP is calculated relative to CFC-11, 

Section II: THE OZONE DEPLETION PROCESS 20 Questions: 2010 Update

Chlorine source gases Bromine source gases

1000

2000

3000

3320

0

16.6%

32.3% CFC-12 (CCl2F2)

CFC-11 (CCl3F)

10.8% Carbon tetrachloride (CCl4)

22%

Methyl chloroform (CH3CCl3) 1%

Methyl chloride (CH3Cl)

Other gases 2.9%

10

20

0

21.6 Other gases 4.3%

15

5

7.5% HCFCs (e.g., HCFC-22 = CHClF2)

6.9% CFC-113 (CCl2FCClF2)

Very short-lived gases

(e.g., bromoform = CHBr3)

27.8%

 29.8%

19.3% Halon-1211 (CBrClF2)

14.7% Halon-1301 (CBrF3)

4.1%

Human-activity sources

Ozone-depleting substances (ODSs)

Human-activity

sources

ODSs

Natural

sources

Natural sources

Halogen Source Gases Entering the Stratosphere in 2008

Total chlorine amount (parts per trillion)

Total bromine amount (parts per trillion)

Methyl bromide

(CH3Br)
Q.18

 20 Questions: 2010 Update Section II: THE OZONE DEPLETION PROCESS

which has an ODP defined to be 1. The calculations, which 

require the use of computer models of the atmosphere, use 

as the basis of comparison the ozone depletion from an equal 

mass of each gas emitted to the atmosphere. 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. The gases 

with small ODP values generally have short atmospheric lifetimes or contain fewer chlorine and bromine atoms.

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. As a consequence, halogen 

source gases that contain fluorine and no other halogens are 

not classified as ODSs. An important category is the hydrofluorocarbons (HFCs), which are included in Table Q7-1 because 

they are ODS substitute gases with ODPs of zero (see Q18).

Iodine is a component of several gases that are naturally 

emitted from the oceans. Although iodine can participate in 

ozone destruction reactions, these iodine-containing source 

gases generally have very short lifetimes and, as a result, only 

Table Q7-1. Atmospheric lifetimes, global emissions, Ozone Depletion Potentials, and Global Warming Potentials of some halogen source gases and HFC substitute gases.

Gas Atmospheric

Lifetime (years)

Global Emissions

in 2008 (Kt/yr) a

Ozone Depletion

Potential (ODP) c

Global Warming 

Potential (GWP) c

Halogen source gases

Chlorine gases

CFC-11 45 52–91 1 4750

CFC-12 100 41–99 0.82 10900

CFC-113 85 3–8 0.85 6130

Carbon tetrachloride (CCl4) 26 40–80 0.82 1400

HCFCs 1–17 385–481 0.01–0.12 77–2220

Methyl chloroform (CH3CCl3) 5 Less than 10 0.16 146

Methyl chloride (CH3Cl) 1 3600–4600 0.02 13

Bromine gases

Halon-1301 65 1–3 15.9 7140

Halon-1211 16 4–7 7.9 1890

Methyl bromide (CH3Br) 0.8 110–150 0.66 5

Very short-lived gases (e.g., CHBr3) Less than 0.5 b b very low b very low

Hydrofluorocarbons (HFCs)

HFC-134a 13.4 149 ± 27 0 1370

HFC-23 222 12 0 14200

HFC-143a 47.1 17 0 4180

HFC-125 28.2 22 0 3420

HFC-152a 1.5 50 0 133

HFC-32 5.2 8.9 0 716

a

 Includes both human activities (production and banks) and natural sources. Emissions are in units of kilotonnes per year 

(1 kilotonne = 1000 metric tons = 1 gigagram = 109 grams).

b

 Estimates are very uncertain for most species.

c

 100-year GWPs. ODPs and GWPs are discussed in Q18. Values are calculated for emissions of an equal mass of each gas.
Q.19

Section II: THE OZONE DEPLETION PROCESS 20 Questions: 2010 Update

a very small fraction reaches the stratosphere. There are large 

uncertainties in how these emissions vary with season and 

geographical region.

Other non-halogen gases. Other non-halogen gases 

that influence stratospheric ozone abundances have also 

increased in the stratosphere as a result of emissions from 

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 participate in 

the destruction of stratospheric ozone (see Q2). Increasing 

abundances of N2O and CH4, as well as CO2, are expected to 

significantly affect future stratospheric ozone through combined effects on temperature, winds, and chemistry (see Q20). 

Although all of these gases are part of the Kyoto Protocol (see 

Q15) because they are climate gases, they are not classified as 

ODSs under the Montreal Protocol. Although past emissions 

of ODSs still dominate global ozone depletion, the current

emissions of N2O from human activities will destroy more 

stratospheric ozone than the current emissions of any ODS.

Heavier-Than-Air CFCs

CFCs and other ozone-depleting substances reach the stratosphere despite the fact that they are “heavier than air.” For 

example, molecules of CFC-11 (CCl3F) and CFC-12 (CCl2F2) are approximately 4–5 times heavier than the average molecule 

of air, since air is composed primarily of oxygen and nitrogen. The emissions of long-lived gases accumulate in the lower 

atmosphere (troposphere). The distribution of these gases in the troposphere and stratosphere is not controlled by the 

molecular weight of each gas because air is in continual motion in these regions as a result of winds and convection. Continual air motions ensure that new emissions of long-lived gases are horizontally and vertically well mixed throughout the 

troposphere within a few months. It is this well-mixed air that enters the lower stratosphere from upward air motions in 

tropical regions, bringing with it ozone-depleting substances emitted from any location on Earth’s surface.

Atmospheric measurements confirm that ozone-depleting substances 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 generally 

consistent with the emission 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, and 

not found only in the upper atmosphere. Noble gases from very light helium to very heavy xenon, which all have very long 

atmospheric lifetimes, are also uniformly distributed throughout the troposphere and stratosphere. 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.
Q.20

Figure Q8-1. Conversion of halogen source gases. Halogen source gases containing chlorine and bromine are chemically converted to reactive halogen gases primarily in the stratosphere. The conversion requires solar ultraviolet radiation 

and a few chemical reactions. The shorter-lived gases undergo some conversion in the troposphere. The reactive halogen 

gases contain all the chlorine and bromine originally present in the source gases before conversion. The reactive gases can 

be grouped into the reservoir gases, which do not directly destroy ozone, and the most reactive gases, which participate in 

ozone destruction cycles (see Q9). A principal reactive gas, ClO, is formed by reactions of the reservoir gases HCl and ClONO2

that occur on the surfaces of liquid and solid polar stratospheric clouds (PSCs) (see Q10).

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 (Figure Q8-1). The source gases, which 

include ozone-depleting substances (ODSs), 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 for most 

gases instead of the troposphere because solar ultraviolet 

radiation (sunlight) is more intense in the stratosphere.

Reactive halogen gases. The chemical conversion of 

halogen source gases, which involves solar ultraviolet radiation (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 

bromine-containing gases that form in the stratosphere are 

shown in Figure Q8-1. Throughout the stratosphere, the most 

abundant are typically hydrogen chloride (HCl) and chlorine 

nitrate (ClONO2). These two gases are considered important 

reservoir gases because, while they don’t react directly with 

ozone, they can be converted to the most reactive forms that 

 20 Questions: 2010 Update Section II: THE OZONE DEPLETION PROCESS

Emissions from human activities and natural processes represent a large source of chlorine- and bromine-containing gases 

that enter the stratosphere. When exposed to ultraviolet radiation from the Sun, these halogen source gases are converted 

to more reactive gases containing chlorine and bromine. Some reactive gases act as chemical reservoirs that convert to form 

the most reactive gases, namely chlorine monoxide (ClO) and bromine monoxide (BrO). The most reactive gases participate 

in catalytic reactions that efficiently destroy ozone. Most volcanoes emit some reactive halogen gases that readily dissolve 

in water and are usually washed out of the atmosphere before they can reach the stratosphere.

Q8 What are the reactive halogen gases that destroy

stratospheric ozone? 

Stratospheric Halogen Gases

Chemical

conversion

Solar

ultraviolet (UV)

radiation and

chemical

reactions
Q.21

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 reactive bromine is generally in the form of BrO, 

whereas usually only a small fraction of reactive chlorine is 

in the form of ClO. The special conditions that occur in the 

polar regions in winter cause the reservoir gases ClONO2 and 

HCl to undergo nearly complete conversion to ClO in reactions on polar stratospheric clouds (PSCs) (see Q10).

Reactive chlorine at midlatitudes. Reactive chlorine 

gases have been observed extensively in the stratosphere with 

both local and remote measurement techniques. The measurements from space displayed in Figure Q8-2 are representative of how the amounts of chlorine-containing gases change 

between the surface and the upper stratosphere at midlatitudes. 

Available chlorine (see red line in Figure Q8-2) is the sum of 

chlorine contained in halogen source gases and the reactive 

gases (e.g., HCl, ClONO2, ClO). Available chlorine is constant 

to within about 15 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 

the reactive chlorine gases. At the highest altitudes, available 

chlorine is all in the form of the reactive chlorine gases.

In the altitude range of the ozone layer at midlatitudes, as 

shown in Figure Q8-2, the reactive reservoir gases HCl and 

ClONO2 account for most of the available chlorine. ClO, the 

most reactive gas in ozone depletion, is a small fraction of available chlorine. The low abundance of ClO limits the amount of 

ozone destruction that occurs outside of polar regions.

Reactive chlorine in polar regions. Reactive chlorine 

gases in polar regions undergo large changes between the fall 

and late winter. Meteorological and chemical conditions in both 

polar regions are now routinely observed from space in all seasons. Fall and winter conditions over the Antarctic are contrasted 

in Figure Q8-3 using seasonal observations made near the center 

of the ozone layer (about 18 km (11.3 miles)) (see Figure Q12-3). 

In fall (May), ozone values are high over the entire Antarctic continent and beyond. Temperatures are mid-range, HCl 

and nitric acid (HNO3) are high, and ClO is very low. High 

HCl indicates that substantial conversion of halogen source 

gases has occurred in the stratosphere. In the past decades, 

HCl and ClONO2 reactive reservoir gases have increased substantially in the stratosphere following increased emissions 

of halogen source gases. HNO3 is an abundant, naturally 

Figure Q8-2. Reactive chlorine gas observations. The abundances of chlorine source 

gases and reactive chlorine gases as measured 

from space in 1994 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, the abundances of reactive chlorine 

gases increase with altitude as the amounts of 

chlorine source gases decline. This is a consequence of chemical reactions initiated by solar 

ultraviolet radiation that convert source gases 

to reactive gases (see Figure Q8-1). The principal reactive chlorine gases formed are HCl, 

ClONO2, and ClO. Summing the source gases 

with the reactive gases gives “Available chlorine,” which is nearly constant with altitude 

up to 47 km. In the ozone layer (18–35 km), 

chlorine source gases are still present and HCl 

and ClONO2 are the most abundant reactive 

chlorine gases. (The unit “parts per trillion” is 

defined in the caption of Figure Q7-1.)

Section II: THE OZONE DEPLETION PROCESS 20 Questions: 2010 Update

40

30

20

10

0

20

25

15

10

0

5

0 1000 2000 3000 4000

Chlorine abundance (parts per trillion)

Altitude (kilometers)

Altitude (miles)

Measurements of Reactive Chlorine from Space

November 1994 (35° – 49°N)

Hydrogen

chloride (HCl)

Chlorine source gases

(CFCs, HCFCs, carbon

tetrachloride, etc.)

Chlorine nitrate (ClONO2)

Chlorine monoxide (ClO)

Other gases

Available chlorine
Q.22

 20 Questions: 2010 Update Section II: THE OZONE DEPLETION PROCESS

Temperature Nitric acid (HNO3)

Hydrogen chloride (HCl) Chlorine monoxide (ClO)

Normal ozone amounts in fall (1 May 2008) at 18-km altitude

Chemical Conditions Observed in the Ozone Layer Over Antarctica 

Ozone

Ozone

Temperature

Hydrogen chloride (HCl) Chlorine monoxide (ClO)

Large ozone depletion in late winter (15 September 2008) at 18-km altitude

Nitric acid (HNO3)

Low High

Temperatures and chemical abundances

Ozone

Temperature

HCl

ClO

HNO3

2800 ppb

-60°C

1.7 ppb

1.1 ppb

9.0 ppb

1800

-85

0.2

0.1

1.0
Q.23

Section II: THE OZONE DEPLETION PROCESS 20 Questions: 2010 Update

occurring stratospheric compound that moderates ozone 

destruction chemistry and also condenses to help form polar 

stratospheric clouds (PSCs) (see Q10). Low ClO indicates that 

little conversion of the reactive reservoirs occurs in the fall, 

thereby limiting catalytic ozone destruction.

By late winter (September), ozone amounts reflect substantial depletion at 18-km altitude over an area larger than 

the Antarctic continent. Depletion throughout much of 

the altitude range of the ozone layer created the 2008 ozone 

hole shown in Figure Q11-3 (see altitude profile in Figure 

Q12-3). The associated meteorological and chemical conditions in late winter are very different from those found in 

fall: very low temperatures, very low HCl and HNO3, and 

very high ClO. Low stratospheric temperatures are characteristic of winter when solar heating is reduced. Low HCl 

and high ClO reflect the conversion of the reactive halogen 

reservoir compounds, HCl and ClONO2, to the most reactive chlorine form, ClO. This conversion occurs selectively 

in winter on PSCs and other stratospheric particles, which 

form at the very low temperatures (see Q10). Low HNO3 is 

indicative of its condensation to form PSCs, some of which 

subsequently move to lower altitudes through gravitational 

settling. High ClO abundances generally cause ozone depletion to continue in the Antarctic region until mid-October 

(spring), when the lowest ozone values usually are observed 

(see Q11). As temperatures rise at the end of the winter, halting further PSC formation, ClO is converted back into the 

reactive reservoir species HCl and ClONO2 (see Q10), and 

ozone destruction is curtailed.

Similar though less dramatic changes in meteorological 

and chemical conditions are also observed between fall and 

winter in the Arctic, where winter ozone depletion is less 

severe than in the Antarctic.

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), which can be observed from 

space. Estimates of reactive bromine abundances in the 

stratosphere are larger than expected from the conversion of 

the halons and methyl bromide source gases, suggesting that 

the contribution of the very short-lived bromine-containing 

gases to reactive bromine must also be significant (see Q7).

Other sources. Other emission sources of reactive halogen gases exist that are associated with natural processes and 

human activities. Most emissions become trapped in the 

lower atmosphere when dissolved in water, and ultimately are 

returned to Earth’s surface before they can reach the stratosphere. Volcanoes are an important episodic source of reactive halogen gases (e.g., HCl) that do not reach the stratosphere 

in appreciable quantities (see Q14). Other examples include 

reactive chlorine produced by evaporation of ocean spray. Sea 

salt products dissolve in water and are removed in the lower 

atmosphere. Solid rocket motors, such as those used on the 

Space Shuttle, release reactive chlorine gases directly into the 

troposphere and stratosphere. At current launch rates, the 

emitted quantities are very small in comparison with halogen 

emissions from other human activities.

Figure Q8-3. Chemical conditions in the ozone layer over Antarctica. Observations of the chemical conditions in the 

Antarctic region highlight the changes associated with severe ozone depletion that forms the ozone hole. Satellite instruments now routinely monitor ozone, reactive chlorine gases, and temperatures in the global stratosphere. Results are shown 

here for fall (May) and late winter (September) seasons in Antarctic winter for a narrow altitude region near 18 kilometers (11.2 

miles) within the ozone layer (see Figure Q12-3). Ozone has normal high values in fall before the onset of ozone destruction 

reactions causes wide spread depletion by late winter. High ozone is accompanied by moderate temperatures, normal high 

values of HCl and HNO3, and normal very low ClO amounts. When ClO is not enhanced, significant ozone destruction from 

ozone-depleting substances does not occur. Chemical conditions are quite different in late winter when ozone undergoes 

severe depletion. Temperatures are much lower, HCl has been converted to ClO, the most reactive chlorine gas, and HNO3

has been removed by the gravitational settling of PSC particles. ClO values closely surrounding the South Pole are low in 

September because ClO formation requires sunlight, which is still gradually returning to the highest latitudes. The high ClO 

values in late winter 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 typically reaches its minimum values in early to mid-October 

(see Q12). Note that the first and last colors in the color bar represent values outside the indicated range of values. (The unit 

“parts per billion,” abbreviated “ppb,” is used here as a measure of the relative abundance of a gas in air: 1 part per billion 

equals the presence of one molecule of a gas per billion (=109) total air molecules (compare to ppt in Figure Q7-1)).
Q.24

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: Cl + O3

and ClO + O. 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 re-formed. 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, and up to tens of thousands of ozone molecules 

during the total time of its stay in the stratosphere.

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 clouds (PSCs) (see 

Q8 and 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 

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 thousands of ozone molecules before it 

leaves the stratosphere. In this way, a small amount of reactive chlorine or bromine has a large impact on the ozone layer. 

A special situation develops in polar regions in the late winter/early spring season where large enhancements in the abundance of the most reactive gas, chlorine monoxide, leads to severe ozone depletion.

What are the chlorine and bromine reactions that destroy Q9 stratospheric ozone? 

 20 Questions: 2010 Update Section II: THE OZONE DEPLETION PROCESS

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. Then, 

Cl reacts with ozone and re-forms ClO, consuming ozone in the process. The cycle then begins 

again with another reaction of ClO with O. Chlorine is considered a catalyst for ozone destruction 

because Cl and ClO are re-formed each time the 

reaction cycle is completed, and ozone is simply 

removed. Atomic oxygen (O) is formed when solar 

ultraviolet radiation (sunlight) reacts with ozone 

and oxygen molecules. Cycle 1 is most important 

in the stratosphere at tropical and middle latitudes, 

where solar ultraviolet radiation is most intense.

Cl 

Cl 

Ozone Destruction Cycle 1

Chlor ine atom (Cl) Oz one (O 3 ) 

Oz one 

destr uction 

Chlor ine mono xide (ClO) 

Chlorine

catalytic

cycle

Oxygen molecule (O 2 ) 

Oxygen atom (O) 

ClO + O 

reaction 

Cl + O 3 reaction 

Net: O + O3 2 O2

Oxygen molecule (O 2 ) 

ClO + O Cl + O2

Cl + O 3 ClO + O2
Q.25

Section II: THE OZONE DEPLETION PROCESS 20 Questions: 2010 Update

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 

re-formed each time the reaction cycle is completed. Visible sunlight is required to complete each cycle and to help form and 

maintain ClO abundances. During polar night and other periods of darkness, ozone cannot be destroyed by these reactions.

in the late winter/early spring season (see Q11 and Q12). At 

high ClO abundances, the rate of polar 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 solar ultraviolet 

radiation. Cycle 1 is most important in the stratosphere at 

tropical and middle latitudes, where solar ultraviolet radiation (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 stratosphere, 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 stratosphere. The visible sunlight needed in 

Cycles 2 and 3 is not sufficient to form ozone because this 

process requires solar ultraviolet radiation (see Q1). In the 

stratosphere in the late winter/early spring period, solar ultraviolet radiation is weak at low Sun angles. As a result, ozone 

destruction by Cycles 2 and 3 in the sunlit winter stratosphere 

greatly exceeds ozone production.

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 ozone-destruction 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. 

The sources of reactive hydrogen and nitrogen gases are the 

naturally occurring gases methane (CH4) and nitrous oxide 

(N2O), respectively. The importance of reactive hydrogen and 

nitrogen gases relative to reactive halogen gases is expected to 

increase in the future, because the atmospheric abundances 

of the reactive halogen gases are decreasing as a result of the 

Montreal Protocol, while CH4 and N2O abundances are projected to increase substantially due to human activities.

ClO + ClO

(ClO)2 + sunlight

ClOO

2(Cl + O3

Net: 2O3

Cycle 2

(ClO)2

ClOO + Cl

Cl + O2

ClO + O2)

3O2

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 3

or( )

()g

ClOO

2(Cl + O3

Net: 2O3

Cl + O2

ClO + O2)

3O2

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 3

or( )

Ozone Destruction Cycles in Polar Regions
Q.26

 20 Questions: 2010 Update Section II: THE OZONE DEPLETION PROCESS

Replacing Lost Ozone in the Stratosphere

The idea is sometimes put forth that humans could compensate for lost global stratospheric ozone by replacing it. Ozone 

could be manufactured, stored, transported to the stratosphere, and released in depleted regions. Unfortunately, the idea 

has substantial practical limitations.

Ozone amounts in the stratosphere reflect a continual balance between chemical production and destruction (see Q2). 

The addition of chlorine and bromine to the stratosphere from human activities has changed the natural balance by increasing ozone destruction and, thereby, lowering stratospheric ozone amounts. Chlorine and bromine destroy ozone in catalytic reactions that allow each atom to destroy many thousands of ozone molecules (see Q9). A one-time injection of manufactured ozone to the stratosphere would not restore the natural balance because the added ozone would be destroyed in 

the same chemical reactions with chlorine and bromine within about a year. Thus, ozone additions would need to be large 

and continuous as long as stratospheric chlorine and bromine amounts remained enhanced above natural amounts, a condition expected to persist for several decades (see Q16). Continuous replacement for decades would pose unprecedented 

technical and resource challenges that will likely be unacceptable to decision makers in the international community.

Specific technical 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. Compensating for the average global ozone loss, currently about 3%, would require 90 

megatons of ozone to be distributed regularly throughout the stratosphere 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 annually 

in the United States, for example, which is now approximately 4 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 developed. Concerns for a 

global delivery system would include further significant energy use and unforeseen environmental consequences.
Q.27

The severe depletion of stratospheric ozone in late winter and early spring in the Antarctic is known as the “ozone 

hole” (see Q11). The ozone hole first appeared over Antarctica because atmospheric and chemical conditions unique to 

this region increase the effectiveness of ozone destruction by 

reactive halogen gases (see Q8). In addition to an abundance 

of these reactive gases, the formation of the Antarctic ozone 

hole requires temperatures low enough to form polar stratospheric clouds (PSCs), isolation from air in other stratospheric 

regions, and sunlight.

Distributing halogen gases. Halogen source gases emitted at Earth’s surface are present in comparable abundances 

Ozone-depleting substances 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 atmospheric and chemical conditions that exist there and nowhere else on the globe. The very low 

winter temperatures in the Antarctic stratosphere cause polar stratospheric clouds (PSCs) to form. Special reactions that 

occur on PSCs, combined with the relative isolation of polar stratospheric air, allow chlorine and bromine reactions to produce the ozone hole in Antarctic springtime.

Why has an “ozone hole” appeared over Antarctica when ozone- Q10 depleting substances are present throughout the stratosphere? 

Section II: THE OZONE DEPLETION PROCESS 20 Questions: 2010 Update

Figure Q10-1. Arctic and Antarctic temperatures. Air temperatures in both polar 

regions reach minimum values in the lower 

stratosphere in the winter season. Average 

daily 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 late December and 

January. Polar stratospheric clouds (PSCs) 

are formed in the polar ozone layer when 

winter minimum temperatures fall below 

the formation temperature of 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 liquid and solid PSC particles 

cause the highly reactive chlorine gas ClO 

to be formed, which catalytically destroys 

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 contrast, PSC formation temperatures are always present for many months somewhere in the 

Antarctic, and severe ozone depletion now occurs in each winter season (see Q11). (Note that the dashed black lines denote 

the upper limits of the Antarctic temperature range where they overlap with the Arctic temperature range.)

-60 -50

-55

-60

-65

-70

-75

-80

-85

-90

-95

-70

-80

-90

-100

-110

-120

-130

-140

Arctic winter

Antarctic winter

Nov Dec Jan Feb March April

May June July August Sept Oct

Temperature (degrees Fahrenheit)

Temperature (degrees Celsius) Minimum Air Temperatures in the Polar Stratosphere

Antarctic 1979 to 2009

Arctic 1978-79 to 2009-10

Average daily values 40° to 90° Latitude

Arctic

Antarctic

PSC formation temperature

Ranges of daily values
Q.28

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 significant natural removal processes in the lower atmosphere and because winds and convection redistribute and 

mix air efficiently throughout the troposphere on the timescale of weeks to months. Halogen gases (in the form of source 

gases and some reactive products) enter the stratosphere primarily from the tropical upper troposphere. Stratospheric air 

motions then transport these gases upward and toward the 

pole in both hemispheres.

Low polar temperatures. The severe ozone destruction 

represented by the ozone hole requires that low temperatures 

be present over a range of stratospheric altitudes, over large 

geographical regions, and for extended time periods. Low 

temperatures are important because they allow liquid and 

solid PSCs to form. Reactions on the surfaces of these PSCs 

initiate a remarkable increase in the most reactive chlorine 

gas, chlorine monoxide (ClO) (see below and Q8). Stratospheric temperatures are lowest in both polar regions in 

winter. In the Antarctic winter, minimum daily temperatures are generally much 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. These and other meteorological differences occur because of the unequal distribution among land, 

ocean, and mountains between the hemispheres at middle 

and high latitudes. The winter temperatures are low enough 

for PSCs to form somewhere in the Antarctic for nearly the 

entire winter (about 5 months) and in the Arctic for only limited periods (10–60 days) in most winters.

Isolated conditions. Stratospheric air in the polar 

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, forming 

a polar vortex, which prevents substantial motion of air into 

or out of the polar stratosphere. This circulation strengthens 

in winter as stratospheric temperatures decrease, with the 

result that the isolation of air in the vortex is much more effective in the Antarctic than the Arctic. Once chemical changes 

occur in polar regions from reactions on PSCs, the isolation 

preserves those changes for many weeks to months.

Polar stratospheric clouds (PSCs). Reactions on the 

surfaces of liquid and solid PSCs can substantially increase 

the relative abundances of the most reactive chlorine gases. 

These reactions convert the reservoir forms of reactive chlorine gases, chlorine nitrate (ClONO2) and hydrogen chloride 

(HCl), to the most reactive form, ClO (see Figure Q8-1). ClO 

increases from a small fraction of available reactive chlorine to comprise nearly all chlorine that is available. With 

increased ClO, additional catalytic cycles involving ClO and 

BrO become active in the chemical destruction of ozone 

whenever sunlight is available (see Q9).

Different types of liquid and solid PSC particles 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 

significant altitude ranges. With a temperature threshold of 

 20 Questions: 2010 Update Section II: THE OZONE DEPLETION PROCESS 

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 in the ozone layer during winters in the Arctic and Antarctic stratospheres wherever 

low temperatures occur (see Figure Q10-1). 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 (PSCs)
Q.29

−78°C, PSCs exist in larger regions and for longer time periods 

in the Antarctic than the Arctic. The most common type of 

PSC forms from nitric acid (HNO3) and water condensing on 

pre-existing liquid sulfuric acid-containing particles. Some of 

these particles freeze to form reactive solid particles. At even 

lower temperatures (−85°C or −121°F), water condenses to form 

ice particles. 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, which increases condensation of water and HNO3.

When average temperatures begin increasing by late 

winter, PSCs form less frequently and their surface conversion reactions produce less ClO. Without continued ClO production, ClO amounts decrease and other chemical reactions 

re-form the reactive reservoirs, ClONO2 and HCl. When PSC 

temperatures no longer occur, on average, either by late January to early February in the Arctic or by mid-October in the 

Antarctic, the most intense period of ozone depletion ends.

Nitric acid and water removal. Once formed, PSC 

particles fall to lower altitudes because of gravity. The largest particles can descend several kilometers or more in the 

stratosphere during the low-temperature winter/spring 

period that lasts several months in Antarctica. Because PSCs 

often contain a significant fraction of available HNO3, their 

descent removes HNO3 from regions of the ozone layer. This 

process is called denitrification of the stratosphere. With less 

HNO3, the highly reactive chlorine gas ClO remains chemically active for a longer period, thereby increasing chemical 

ozone destruction. Significant denitrification occurs each 

winter in the Antarctic and in some, but not all, Arctic winters, because PSC formation temperatures must be sustained 

over an extensive altitude region and time period to effect 

denitrification (see Figure Q10-1).

Ice particles form when temperatures are a few degrees 

lower than PSC formation temperatures. If ice temperatures 

persist for weeks to months over extensive altitude regions, 

ice particles will also fall several kilometers due to gravity. 

As a result, a significant fraction of water vapor can also be 

removed from regions of the ozone layer. This process is called 

dehydration of the stratosphere. With the low temperatures 

required to form ice, dehydration is common in the Antarctic 

and rare in the Arctic winters. The removal of water vapor 

does not directly affect the catalytic reactions that destroy 

ozone. Dehydration indirectly affects ozone destruction by 

suppressing PSC formation, which reduces ClO production in 

PSC reactions.

Discovering the role of PSCs. Ground-based observations of PSCs, and knowledge of their formation processes, 

were available many years before the role of PSCs in polar 

ozone destruction was recognized. 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 PSC particles 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 chemical role of PSC particles developed from laboratory 

studies of their surface reactivity, computer modeling studies of polar stratospheric chemistry, and measurements that 

directly sampled particles and reactive chlorine gases, such as 

ClO, in the polar stratosphere.

Section II: THE OZONE DEPLETION PROCESS 20 Questions: 2010 Update
Q.30

 20 Questions: 2010 Update Section II: THE OZONE DEPLETION PROCESS 

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 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 Japan Meteorological Agency and the British Antarctic Survey. The results became widely 

known to the world after three scientists from the British Antarctic Survey published their observations in the scientific 

journal Nature in 1985 and speculated that CFCs were the cause. 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 as a description of the very 

low total ozone values that encircled the Antarctic continent in satellite images for many weeks (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.

Very early Antarctic ozone measurements. The first total ozone measurements made in Antarctica with Dobson 

spectrophotometers occurred in the 1950s following extensive measurements in the Northern Hemisphere and Arctic 

region. Total ozone values found in spring were around 300 DU, surprisingly lower than in the Arctic spring, because the 

assumption then was that the two polar regions would have similar values. We now know that these Antarctic values were 

not anomalous; indeed they are similar to those observed there in the 1970s before the ozone hole appeared (see Figure 

Q11-3). We also now know that Antarctic total ozone values, in comparison with Arctic values, are systematically lower in 

the early spring because the polar vortex is much stronger and, therefore, much more effective in reducing the transport of 

ozone-rich air from midlatitudes to the pole (compare Figures Q11-3 and Q12-2).

In 1958, measurements of total ozone were made at the Dumont d’Urville station (66.7°S, 140°E) in Antarctica using a photographic plate method to analyze solar ultraviolet radiation after it passed through the ozone layer. The reported measurements were anomalously low, reaching 110–120 DU in September and October. These values are similar to minimum ozone 

hole values now routinely observed over Antarctica in the same months (see Figure Q11-2). Some have speculated that 

these limited observations provide evidence that an ozone hole existed before ODS emissions were large enough to cause 

the depletion. However, analyses of the more extensive Dobson spectrophotometer measurements made at several other 

Antarctic locations in 1958 did not confirm the low total ozone values. These measurements indicate that the photographic 

plate determinations were not a reliable source of total ozone values at the Dumont d’Urville station in 1958.
Q.31

Figure Q11-1. Antarctic ozone hole. Total ozone values 

are shown for high southern latitudes on 17 September 

2009 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 Antarctic springtime values of about 350 DU 

(see Figure Q11-3). The ozone hole area is usually defined as 

the geographical area within the 220-DU contour (see white 

line) on total ozone maps. In late spring or early summer 

(November–December) the ozone hole disappears in satellite images as ozone-depleted air is displaced and mixed 

with ozone-rich air transported toward the pole from outside the ozone hole. Note that maximum total ozone values 

in the Southern Hemisphere in this late winter period are 

generally located in a crescent-shaped region surrounding 

the ozone hole.

Section III: STRATOSPHERIC OZONE DEPLETION

The severe depletion of Antarctic ozone, known as the “ozone hole,” was first reported in the mid-1980s. The 

depletion is attributable to chemical destruction by reactive 

halogen gases that increased in the stratosphere in the latter 

half of the 20th century (see Q16). Conditions in the Antarctic 

winter and early spring stratosphere are highly suitable for 

ozone depletion because of (1) the long periods of extremely 

low temperatures, which cause polar stratospheric clouds 

(PSCs) to form; (2) the large abundance of reactive halogen 

gases produced in reactions on PSCs; and (3) the isolation of 

stratospheric air, which allows time for chemical destruction 

processes to occur. The severity of Antarctic ozone depletion 

can be seen using satellite observations of total ozone, ozone 

altitude profiles, and long-term changes in total ozone.

Antarctic ozone hole. The most widely used images of 

Antarctic ozone depletion are derived from measurements of 

total ozone made with satellite instruments. A map of Antarctic early spring measurements shows 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 maps. The area of the ozone hole is defined here as the 

geographical region contained within the 220-Dobson unit 

(DU) contour in total ozone maps (see white line in Figure 

Severe depletion of the Antarctic ozone layer was first reported in the mid-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, thereby reducing total ozone by as much as two-thirds at some 

locations. This severe depletion creates the “ozone hole” apparent in images of Antarctic total ozone made using satellite 

observations. In most years the maximum area of the ozone hole far exceeds the size of the Antarctic continent.

Q11 How severe is the depletion of the Antarctic ozone layer?

100 300 400 500 600

17 September 2009

200

Antarctic Ozone Hole

Total ozone (Dobson units)
Q.32

 20 Questions: 2010 Update Section III: STRATOSPHERIC OZONE DEPLETION

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 to mid-October are near 100 DU, 

which is two-thirds below normal springtime values of about 

350 DU (see Figures Q11-3 and Q12-1). Low total ozone inside 

the ozone hole contrasts strongly with the distribution of much 

larger values outside the ozone hole. This common feature can 

also be seen in Figure Q11-1, where a crescent-shaped region 

with maximum values of about 500 DU surrounds a large portion of the ozone hole in September 2009.

Altitude profiles of Antarctic ozone. The low total 

ozone values within the ozone hole are caused by nearly complete removal of ozone in the lower stratosphere. Balloonborne 

instruments (see Q5) demonstrate that this depletion occurs 

within the ozone layer, the altitude region that normally contains the highest ozone abundances. At geographic locations 

with the lowest total ozone values, balloon measurements 

show that the chemical destruction of ozone is complete over 

an altitude region of several kilometers. For example, in the 

ozone profile over South Pole, Antarctica, on 9 October 2006 

(see Figure Q12-3), ozone abundances are essentially zero over 

the altitude region of 14 to 21 kilometers. The lowest winter 

temperatures and highest reactive chlorine (ClO) abundances 

occur in this altitude region (see Figure Q8-3). The differences 

in the average South Pole ozone profiles between the decades 

1962–1971 and 1990–2009 in Figure Q12-3 show how reactive 

halogen gases have dramatically altered the ozone layer. In 

the 1960s, a normal ozone layer is clearly evident in the October average profile with a peak near 16 kilometers altitude. 

In the 1990–2009 average profile, a broad minimum centered 

near 16 kilometers now occurs with ozone values reduced by 

up to 90% of normal values.

Long-term total ozone changes. The significant 

springtime ozone depletion represented by the ozone hole 

has only been observed since the early 1980s even though low 

winter temperatures and isolated conditions occur each year 

in the Antarctic stratosphere. Prior to 1980, the amounts of 

reactive halogen gases in the stratosphere were insufficient to 

cause significant depletion. Satellite observations can be used 

in multiple ways to examine how ozone depletion has changed 

in the Antarctic region over the last 30–40 years:

▶ First, ozone hole areas displayed in Figure Q11-2 show that 

depletion has increased since 1980 to become fairly stable 

in the 1990s and early 2000s near a value of 25 million 

Figure Q11-2. Antarctic ozone hole features. Long-term changes are shown for key 

aspects of the Antarctic ozone hole: the area 

enclosed by the 220-DU contour on maps of 

total ozone (upper panel) and the minimum 

total ozone amount within the 220-DU contour (lower panel). The values are derived 

from satellite observations and averaged for 

each year at a time near the peak of ozone 

depletion, as defined by the dates shown 

in each panel. The areas of continents are 

included for reference in the upper panel. 

The magnitude of ozone depletion gradually 

increased beginning in 1980. In the 1990s, 

the depletion reached steady year-to-year 

values, except for the unusually low depletion 

in 2002 (see Figure Q11-4 and following box). 

The magnitude of Antarctic ozone depletion 

will decrease as ODSs are removed from the 

atmosphere (see Figure Q16-1). The return 

of Antarctic total ozone to 1980 values is not 

expected to occur before 2050 (see Q20).

1980

200

30

10

0

20

150

100

1990 2000 2010

Year

Minimum total ozone

(Dobson units)

Area (million squar

e

kilometers)

Antarctic Ozone Depletion

Average daily minimum

(21 September–16 October)

Average daily area

(21–30 September, 220 DU)

Africa

North

America

South

America

Antarctica

Europe

Australia
Q.33

 Section III: STRATOSPHERIC OZONE DEPLETION 20 Questions: 2010 Update

square kilometers. The exception is the unexpected low 

depletion in 2002, which is explained in the box at the end 

of this Question. The ozone-hole area is defined here as 

the geographical area inside the 220-DU contour on total 

ozone maps (see Figure Q11-1).

▶ Second, minimum Antarctic ozone amounts displayed in 

Figure Q11-2 show that the severity of the depletion has 

increased beginning around 1980 along with the ozone 

hole area. Fairly constant minimum values near 100 DU 

were observed in the 1990s and 2000s with 2002 being an 

exception.

▶ Third, total ozone maps over the Antarctic and surrounding regions show how the ozone hole has developed over 

time (see Figure Q11-3). October averages of total ozone 

confirm the absence of an ozone hole in the 1970s and its 

persistent occurrence in the late 2000s. 

▶ Fourth, values of total ozone poleward of 63°S averaged for 

each October show how total ozone has changed when 

Figure Q11-3. Antarctic total ozone. Long-term changes in Antarctic total ozone are demonstrated with this series of 

total ozone maps derived from satellite observations. Each map is an average during October, the month of maximum ozone 

depletion over Antarctica. In the 1970s, no ozone hole was observed as defined by a significant region with total ozone values 

less than 220 DU (dark blue and purple colors). Starting in the 1980s, the ozone hole began to appear and increased in size 

(not shown here). A large ozone hole has occurred each year since the early 1990s as shown in Figure Q11-2. The maps from 

the 2000s show the large extent (about 25 million square kilometers) of the most recent ozone holes. The largest total-ozone 

values at high southern latitudes are still found in a crescent-shaped region outside the ozone hole in October. The maps 

show that these maximum values and their extent have significantly diminished since the 1970s. (Note the difference in color 

scale with Figure Q12-2.)

1970 1971 1972 1979

2006 2007 2008 2009

Antarctic Total Ozone

(October monthly averages)

100 200 300

Total ozone (Dobson units)

400 500
Q.34

 20 Questions: 2010 Update Section III: STRATOSPHERIC OZONE DEPLETION

averaged over the ozone hole and its surrounding areas 

(see Figure Q12-1). The values decrease strongly starting 

from those found in the 1970s and now are approximately 

37% less than in pre-ozone-hole years (1970–1982). The 

average values show a larger year-to-year variability than 

found for the other ozone measures noted above because 

the average includes areas outside the ozone hole where the 

distribution of total ozone varies in response to meteorological conditions. The maps in Q11-3 show also how the 

maximum in total ozone surrounding the ozone hole each 

year has diminished over the last three decades, adding to 

the decreases noted in Figure Q12-1.

Disappearance of the ozone hole in spring. The 

severe depletion of Antarctic ozone occurs in the late winter/

early spring season. In spring, temperatures in the polar 

lower stratosphere increase (see Figure Q10-1), ending PSC 

formation and reactions on aerosols and, consequently, the 

most effective chemical cycles that destroy ozone. Wintertime isolation of high-latitude air ends during this time with 

increasing exchange of air between the Antarctic stratosphere 

and lower latitudes. This allows substantial amounts of 

ozone-rich air to be transported poleward, where it displaces 

or mixes with air depleted in ozone. As a result of these largescale transport and mixing processes, the ozone hole disappears by December.

Figure Q11-4. Unusual 2002 ozone hole. Views from space of the Antarctic ozone hole are shown as observed on 24 

September in the years 2001, 2002, and 2003. The ozone holes in 2001 and 2003 are considered typical of those observed 

since the early 1990s. An initially circular hole in 2002 was transformed into two smaller, depleted regions in the days 

preceding 24 September. This unusual event is attributable to an early warming of the polar stratosphere caused by air 

disturbances that originated in the lower atmosphere (troposphere) at midlatitudes. Warmer temperatures reduced the rate 

of ozone depletion. As a consequence, total ozone depletion was unusually low that year in comparison with 2001 and 2003 

and all other years since the early 1990s (see Figure Q11-2).

Unusual 2002 Antarctic Ozone Hole

24 September 2001 24 September 2002 24 September 2003

100 200 300 400 500 600

Total ozone (Dobson units)
Q.35

 Section III: STRATOSPHERIC OZONE DEPLETION 20 Questions: 2010 Update

The 2002 Antarctic Ozone Hole

The 2002 Antarctic ozone hole showed features that looked surprising at the time (see Figure Q11-4). It had much less 

ozone depletion as measured by the area of the ozone hole or minimum total-ozone amounts in comparison with the 2001 

ozone hole. The 2002 values now stand out clearly in the year-to-year changes in these quantities displayed in Figure Q11-2. 

There were no forecasts of an ozone hole with unusual features in 2002 because the chemical and meteorological conditions required to deplete ozone, namely low temperatures and available reactive halogen gases, were present that year and 

did not differ substantially from previous years. The ozone hole initially formed as expected in August and early September 

2002. Later, during the last week of September, an unexpected and surprisingly strong meteorological event occurred that 

dramatically reshaped the ozone hole into two separate depleted regions. As a result of this disturbance, the combined 

area of these two regions in late September and early October was significantly less than that observed for the previous or 

subsequent ozone holes. 

The unexpected meteorological influence in 2002 resulted from specific atmospheric air motions that sometimes occur 

in polar regions. Meteorological analyses of the Antarctic stratosphere show that it was warmed by very strong, large-scale 

weather systems that originated in the lower atmosphere (troposphere) at midlatitudes in late September. At that time, Antarctic temperatures are generally very low (see Q10) and ozone destruction rates are near their peak values. The influence 

of these tropospheric systems extended poleward and upward into the stratosphere, disturbing the normal circumpolar 

wind flow (polar vortex) and warming the lower stratosphere where ozone depletion was ongoing. Higher temperatures 

reduced the rate of ozone depletion and led to the higher minimum values observed for total ozone in Figure Q11-2. The 

higher-than-normal impact of these weather disturbances during the critical time period for ozone loss reduced the total 

loss of ozone in 2002.

The strong influence of the 2002 warming event is unique in the many decades of Antarctic meteorological observations. 

Another warming event occurred in 1988 causing smaller changes in the ozone hole features in Figure Q11-2. Large warming 

events are difficult to predict because of the complex conditions leading to their formation. 

In 2003 through 2009, ozone hole features returned to values observed from the mid-1990s to 2001 (see Figure 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, leading to full recovery of total ozone, requires comparable, 

sustained reductions of ODSs in the stratosphere. Even with the source gas reductions already underway (see Q16), the 

return of Antarctic total ozone to 1980 values is not expected to occur before 2050.
Q.36

Significant depletion of total ozone has been observed in the Arctic stratosphere in recent decades. 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). Arctic depletion also occurs in the late 

winter/early spring period (January–March) over a somewhat 

shorter period than in the Antarctic (July–October). Similar 

to the Antarctic (see Q11), Arctic depletion occurs because of 

(1) the periods of extremely low temperatures, which cause 

polar stratospheric clouds (PSCs) to form; (2) the large abundance of reactive halogen gases produced in reactions on 

PSCs; and (3) the isolation of stratospheric air, which allows 

time for chemical destruction processes to occur. 

Arctic ozone depletion is much less than that observed each 

Antarctic winter/spring season. Large and recurrent ozone 

holes as found in the Antarctic stratosphere do not occur in 

the Arctic. Depletion is limited because, in comparison to 

Antarctic conditions, Arctic average temperatures are always 

significantly higher (see Figure Q10-1) and the isolation of 

stratospheric air is less effective. Temperature and other 

meteorological differences occur because northern polar latitudes have more land and mountainous regions than southern polar latitudes (compare Figures Q11-3 and Q12-2). In a 

few Arctic winters, for example, PSCs did not form because 

temperatures never reached low enough values. These differences cause the extent and timing of Arctic ozone depletion 

to vary considerably from year to year. Depletion in some 

winter/spring seasons occurs over many weeks; in others only 

for brief early or late periods; and in some not at all.

Long-term total ozone changes. Satellite observations 

can be used in two important ways to examine the average 

total ozone abundances in the Arctic region for the last 30–40 

years and to contrast these values with Antarctic abundances:

▶ First, total ozone poleward of 63°N averaged for each March 

shows quantitatively how total ozone has changed in the 

Arctic (see Figure Q12-1). The seasonal poleward and 

downward transport of ozone-rich air is naturally stronger in the Northern Hemisphere. As a result, total ozone 

values at the beginning of each winter season in the Arctic 

are considerably higher than in the Antarctic. Before 

depletion sets in, normal Arctic values are close to 450 

DU and Antarctic values near 330 DU. Decreases from 

pre-ozone-hole average values (1970–1982) were observed 

in the Arctic by the mid-1980s, when larger changes were 

already occurring in the Antarctic. The decreases in total 

ozone have reached a maximum of about 30% and generally remain smaller than those found in the Antarctic. The 

low value of Arctic total ozone in March 1997 relative to 

1970–1982 observations is the most comparable to Antarctic depletion. In the 1996/1997 Arctic winter, low temperatures facilitated large amounts of chemical depletion, while 

meteorological conditions kept ozone transport to high 

latitudes below average values.

 Overall, the Arctic values tend to show larger year-toyear variability than in the Antarctic. Ozone differences 

from the 1970-to-1982 average value are due to a combination of chemical destruction by ODSs and meteorological (natural) variations. In the last two decades, these two 

aspects have contributed about equally to observed ozone 

changes. The amount of chemical destruction depends in 

large part on stratospheric temperatures. Meteorological 

conditions determine how well Arctic stratospheric air is 

isolated from ozone-rich air at lower latitudes and influence the extent of low temperatures.

▶ Second, total ozone maps over the Arctic and surrounding regions (see Figure Q12-2) show year-to-year changes 

in March total ozone. In the 1970s, total ozone values 

were near 450 DU when averaged over the Arctic region in 

March. Beginning in the 1990s and continuing into the late 

2000s, values above 450 DU were increasingly absent from 

the March average maps. A comparison of the 1971 and 

2009 maps, for example, shows a striking reduction of total 

ozone throughout the Arctic region. The large geographical 

extent of low total ozone in the map of March 1997 is rare in 

Yes, significant depletion of the Arctic ozone layer now occurs in most years in the late winter/early spring period (January–

March). 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.

Q12 Is there depletion of the Arctic ozone layer?

 20 Questions: 2010 Update Section III: STRATOSPHERIC OZONE DEPLETION
Q.37

the Arctic observational record of the last three decades as 

noted above in the discussion of Figure Q12-1. 

Altitude profiles of Arctic ozone. Arctic ozone is 

measured using a variety of instruments (see Q5), as in the 

Antarctic, to document daily to seasonal changes within 

the ozone layer. Spring Arctic and Antarctic balloonborne 

measurements are contrasted in Figure Q12-3 using Arctic 

profiles from the Ny-Ålesund research station at 79°N. For 

1991–2009, the March average reveals a substantial ozone 

layer and total ozone of 382 DU, contrasting sharply with the 

severely depleted Antarctic ozone layer in the October average for these years. This further demonstrates how higher 

stratospheric temperatures and meteorological variability 

have protected the Arctic ozone layer from the greater ozone 

losses that occur in the Antarctic, despite similar reactive 

halogen abundances in the two regions.

The separate Arctic profile shown for 29 March 1996 is one 

of the most severely depleted in the two-decade record from 

Ny-Ålesund. Although significant, this depletion is modest in 

comparison to that routinely observed in the Antarctic, such 

as in the profile from 9 October 2006. The near-complete 

depletion of ozone over many kilometers in altitude, as is 

now common in the Antarctic stratosphere, has never been 

observed in the Arctic.

Restoring ozone in spring. As in the Antarctic, ozone 

depletion in the Arctic is confined to the late winter/early 

spring season. In spring, temperatures in the polar lower 

stratosphere increase (see Figure Q10-1), ending PSC formation and reactions on aerosols, as well as the most effective 

chemical cycles that destroy ozone. Wintertime isolation 

of high-latitude air ends during this time with increasing 

exchange of air between the Arctic stratosphere and lower 

latitudes. This allows more ozone-rich air to be transported 

poleward, where it displaces or mixes with air in which ozone 

may have been depleted. As a result of this large-scale transport and mixing processes, any ozone depletion disappears by 

April or earlier.

 Section III: STRATOSPHERIC OZONE DEPLETION 20 Questions: 2010 Update

Figure Q12-1. Average total 

ozone in polar regions. Longterm changes in average total ozone 

are shown for the Antarctic and Arctic 

for the respective regions defined by 

latitudes between 63° and 90°. Total 

ozone is measured with satellite 

instruments. The reference values 

(red lines) are averages of springtime 

total ozone values available from 

observations between 1970 and 1982. 

Each point represents a monthly average for October in the Antarctic or 

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 ozone, particularly in the Arctic. Essentially all of the decrease in the 

Antarctic and usually about 50% of the decrease in the Arctic each year are attributable to chemical destruction by reactive 

halogen gases. In the Arctic, the other 50% is attributable to natural variations in the amounts of ozone transported to polar 

regions before and during winter. Average total ozone values over the Arctic are naturally larger at the beginning of each 

winter season because, in the preceding months, more ozone is transported poleward in the Northern Hemisphere than in 

the Southern Hemisphere.

Average Total Ozone in Polar Regions

500

450

400

350

300

250

200

Total ozone (Dobson units)

1970 1980 1990 2000 2010

Year

Chemical destruction

and natural variation

of ozone

1970 – 1982 Average

(63°–90° Average)

Arctic region in March

Antarctic region in October
Q.38

Figure Q12-2. Arctic total ozone. Long-term changes in Arctic total ozone are evident in this series of total ozone maps 

derived from satellite observations. Each map is an average during March, the month when some ozone depletion is usually 

observed in the Arctic. In the 1970s, the Arctic region had normal ozone values in March, with values of 450 DU and above 

(orange and red colors). Ozone depletion on the scale of the Antarctic ozone hole does not occur in the Arctic. Instead, late 

winter/early spring ozone depletion has eroded the normal high values of total ozone. On the maps from the late 2000s, 

the extent of values of 450 DU and above is greatly reduced in comparison with the 1970s maps. The large region of low 

total ozone in 1997 is unusual in the Arctic record and is attributable to meteorological conditions that led to below-average 

stratospheric temperatures and a strong polar vortex that winter. (Note the difference in color scale with Figure Q11-3.)

1971 1972 1979 1980

1997 2006 2008 2009

Arctic Total Ozone

(March monthly averages)

200 300

Total ozone (Dobson units)

400 500 600

 20 Questions: 2010 Update Section III: STRATOSPHERIC OZONE DEPLETION
Q.39

 Section III: STRATOSPHERIC OZONE DEPLETION 20 Questions: 2010 Update

Figure Q12-3. Vertical distribution of Arctic and Antarctic ozone. Most stratospheric ozone resides between about 

10 and 30 kilometers (6 to 19 miles) above Earth’s surface. Long-term observations of the ozone layer with balloonborne 

instruments allow winter ozone altitude profiles to be compared between the Antarctic and Arctic regions. In the Antarctic 

at the South Pole (left panel), a normal ozone layer was observed to be present between 1962 and 1971. In more recent 

years, as shown here for 9 October 2006, ozone is almost completely destroyed between 14 and 21 kilometers (9 to 13 miles) 

in the Antarctic in spring. Average October values in the last decades (1990–2009) are 90% lower than pre-1980 values at 

the peak altitude of the ozone layer (16 kilometers). In contrast, the Arctic ozone layer is still present in spring as shown by 

the average March profile for 1991–2009 obtained over the Ny-Ålesund site (right panel). No Ny-Ålesund data are available 

for the 1962–1971 period before significant ODS destruction began. Some March profiles do reveal significant depletion, 

as shown here for 29 March 1996. In such years, winter minimum temperatures are generally lower than normal, allowing 

PSC formation for longer periods. Arctic profiles with depletion similar to that shown for 9 October 2006 at the South Pole 

have never been observed. The number in parentheses for each profile is the total ozone value in Dobson units (DU). Ozone 

abundances are shown here as the pressure of ozone at each altitude using the unit “milli-Pascals” (mPa) (100 million mPa = 

atmospheric sea-level pressure).

Polar Ozone Depletion

0 5 10 15 0 5 10 15

0

5

10

15

20

25

30

Altitude (km)

Ozone abundance (mPa)

Antarctic ozone

South Pole (90°S)

Arctic ozone

Ny-Ålesund (79°N)

October averages

1962–1971 (298 DU)

1990–2009 (132 DU)

9 October 2006 (93 DU)

March averages

1962–1971 (no data)

1991– 2009 (382 DU)

29 March 1996 (258 DU)

Altitude (miles) 0

5

10

15
Q.40

Global total ozone has decreased beginning in the 1980s (see Figure Q13-1). The decreases have occurred in the 

stratospheric ozone layer where most ozone resides (see Figure 

Q1-2). In the early 1990s, the depletion of global total ozone 

reached a maximum of about 5% below the 1964–1980 average. The depletion has lessened since then and now averages 

about 3.5% for 2006–2009. The observations shown in Figure 

Q13-1 have been smoothed to remove regular ozone changes 

that are due to natural seasonal and solar effects (see Q14). 

The depleted amounts are larger than the remaining natural 

variations in global total ozone amounts. 

The observed global ozone depletion in the last three 

decades is attributable to increases in reactive halogen gases 

in the stratosphere. The lowest global total ozone values since 

1980 have occurred in the years following the eruption of Mt. 

Pinatubo in 1991, which temporarily increased the number 

of sulfuric acid-containing particles throughout the stratosphere. These particles significantly increased the effectiveness of reactive halogen gases in destroying ozone (see Q14) 

and, thereby, increased global ozone depletion by 1–2% for 

several years following the eruption. 

Polar regions. Observed total ozone depletion varies 

significantly with latitude on the globe (see Figure Q13-1). 

The largest reductions have occurred at high 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 high latitudes of the Northern Hemisphere, 

caused in part by winter losses over the Arctic. Although the 

depletion in polar regions is larger than at lower latitudes, the 

influence of polar regions on global ozone is limited by their 

small geographical area. Latitudes above 60° account for only 

about 13% of Earth’s surface.

Midlatitude regions. Ozone depletion is also observed 

at the midlatitudes spanning the region between equatorial 

and polar latitudes. In comparison with the 1964–1980 averages, total ozone averaged for 2005–2009 is about 3.5% lower 

in northern midlatitudes (35°N–60°N) and about 6% lower 

at southern midlatitudes (35°S–60°S). Midlatitude depletion 

has two contributing factors. First, ozone-depleted air over 

both polar regions is dispersed away from the poles during 

and after each winter/spring period, thereby reducing average 

ozone at nonpolar latitudes. Second, chemical destruction 

occurring at midlatitudes contributes to observed depletion 

in these regions. This contribution is much smaller than in 

polar regions because the amounts of reactive halogen gases 

are lower and a seasonal increase of the most reactive halogen 

gases, such as the increase in ClO in Antarctic late winter (see 

Figure Q8-3), does not occur in midlatitude regions.

Tropical region. There has been little or no depletion of 

total ozone in the tropics (20°N–20°S latitude). In this region 

of the lower stratosphere, air has only recently (less than 18 

months) been transported from the lower atmosphere (troposphere). As a result, the conversion of ozone-depleting substances (ODSs) to reactive halogen gases is very small. With 

so little reactive halogen amounts, total ozone depletion in 

this region is also very small. In addition, ozone production is 

high in the tropical stratosphere because average solar ultraviolet radiation is highest in the tropics. In contrast, stratospheric air in polar regions has been in the stratosphere for an 

average of 4 to 7 years, allowing time for significant conversion of ODSs to reactive halogen gases. The systematic differences in the age of stratospheric air are a well-understood 

consequence of the large-scale atmospheric transport: air 

enters the stratosphere in the tropics, moves poleward into 

both hemispheres, and then descends and ultimately returns 

to the lower atmosphere.

Recovery of global ozone. Global ozone is no longer 

declining as it was in the 1980s and early 1990s because 

ODSs are no longer increasing in the atmosphere (see Q16). 

During recovery from ODSs, global total ozone is expected 

to reach 1980 and earlier values in the coming decades. The 

recovery process depends on the slow removal of ODSs from 

Depletion of the global ozone layer began gradually in the 1980s and reached a maximum of about 5% in the early 1990s. 

The depletion has lessened since then and now is about 3.5% averaged over the globe. The average depletion exceeds the 

natural year-to-year variations of global total ozone. The ozone loss is very small near the equator and increases with latitude toward the poles. The larger polar depletion is attributed to the late winter/early spring ozone destruction that occurs 

there each year.

Q13 How large is the depletion of the global ozone layer?

 20 Questions: 2010 Update Section III: STRATOSPHERIC OZONE DEPLETION
Q.41

 Section III: STRATOSPHERIC OZONE DEPLETION 20 Questions: 2010 Update

Figure Q13-1. Global total ozone changes. Satellite 

observations show depletion of global total ozone beginning in the 1980s. The top panel compares annual averages of global ozone with the average from the period 

1964 to 1980 before the ozone hole appeared. Seasonal 

and solar effects have been removed from the observational data set. On average, global ozone decreased each 

year between 1980 and 1990. The depletion worsened for 

a few years after 1991 due to the effect of volcanic aerosol 

from the Mt. Pinatubo eruption. Average global ozone 

for 2005–2009 is about 3.5% below the 1964-to-1980 

average. The bottom panel shows how the 2005–2009 

depletion varies with latitude over the globe. The largest 

decreases have occurred at high 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 than 

at mid or high latitudes and ozone production rates are 

greater.

the stratosphere following emission reductions. Future 

changes in climate parameters will also influence ozone. The 

global ozone increases observed in the last 20 years cannot 

be attributed solely to reductions in ODSs that began in the 

1990s, because global ozone also responded strongly to the 

Mt. Pinatubo eruption. The projections of long-term changes 

in total ozone for different regions of the globe are described 

in Q20.

Global Total Ozone Changes

Changes from 1964–1980 average

1970 1980 1990 2000 2010

0

-2

2

-4

-6

-8

Year

Annual averages

Range of observations

Ozone change (%)

Latitude (degrees)

-90 -60 -30 0 30 60 90

0

-10

Ozone change (%) -20

South North

Average

Uncertainty range

Changes between 1964–1980

and 2005–2009 averages
Q.42

Changes in solar radiation and increases in stratospheric particles from volcanic eruptions both affect the abundance of stratospheric ozone. Over the last three decades, 

global total ozone has decreased over the globe and is now 

about 3.5% below pre-1980 values (see Q13). The depletion is 

attributed to changes in reactive halogen gases, which are represented by changes in equivalent effective stratospheric chlorine (EESC). EESC values account for stratospheric chlorine 

and bromine abundances and their different effectiveness in 

destroying ozone (see definition in Q16). A comparison of the 

smooth year-to-year changes in ozone and EESC shows that 

the quantities are inversely related to each other, with ozone 

first decreasing while EESC increases (see Figure Q14-1). 

After the mid-1990s, the annual changes in both quantities 

are sharply reduced. Changes in solar output and volcanic 

activity do not show such smooth long-term changes, as discussed below, and therefore are not considered to be the cause 

of long-term global ozone depletion.

Total ozone and 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-documented 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 incoming solar radiation at a wavelength of 10.7 cm are 

often used as a surrogate for changes in solar output at UV 

wavelengths. The long-term changes in the 10.7-cm output in 

Figure Q14-1 clearly show alternating periods of maximum 

and minimum values in total solar output separated by about 

5–6 years. If changes in solar output were the cause of global 

ozone depletion, a gradually decreasing output would have 

been observed around 1980 or earlier, slowing sharply in the 

mid-1990s. Since such a decrease was not observed, nor is 

expected based on longer-term solar observations, the longterm decreases in global ozone cannot result from changes 

in solar output alone. Most analyses presented in this and 

previous international scientific assessments quantitatively 

account for the influence of the 11-year solar cycle on longterm variations in ozone.

Total ozone and past volcanoes. Explosive volcanic 

eruptions inject sulfur gases directly into the stratosphere, 

causing new sulfate particles to be formed. The particles initially form in the stratosphere downwind of the volcano and 

then spread throughout the hemisphere or globally as air is 

transported by stratospheric winds. One method of detecting the presence of volcanic particles in the stratosphere uses 

observations of the transmission of solar radiation through 

the atmosphere (see Figure Q14-1). When large amounts of 

new particles are formed in the stratosphere over an extensive 

region, solar transmission is measurably reduced. The eruptions of Mt. Agung (1963), El Chichón (1982), and Mt. Pinatubo (1991) are the most recent examples of sulfur injections 

that temporarily reduced solar transmission.

Laboratory measurements and stratospheric observations 

have shown that chemical reactions on the surfaces of volcanically produced particles can increase ozone destruction 

by increasing the amounts of the highly reactive chlorine gas 

chlorine monoxide (ClO). The ozone response depends on 

the total abundance of EESC after the eruption (see Q16). At 

times of relatively low EESC, such as the early 1980s, ozone is 

not very sensitive to stratospheric injection of volcanic sulfate 

particles. At times of higher EESC amounts, such as from 1980 

to the present, global ozone is expected to decrease significantly 

following large explosive eruptions. 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). EESC 

was too low for ozone depletion to occur after the Mt. Agung 

eruption in 1963. The effect on ozone diminishes during the 

years following an eruption as volcanic particles are gradually 

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 three decades. If large volcanic eruptions occur in the coming decades, ozone depletion will increase for several 

years afterwards.

Do changes in the Sun and volcanic eruptions affect Q14 the ozone layer?

 20 Questions: 2010 Update Section III: STRATOSPHERIC OZONE DEPLETION
Q.43

Figure Q14-1. Solar changes and volcanoes.

Global total ozone values have decreased beginning in the early 1980s. Ozone values shown 

are the 3-month averages without corrections 

for solar or volcanic effects and annual averages 

corrected for seasonal and solar effects (from 

Figure Q13-1). These long-term ozone decreases

are primarily attributable to equivalent effective 

stratospheric chlorine (EESC), which has increased

over the same time period. Since the mid-1990s, 

changes in both quantities have slowed. Incoming solar radiation varies on a well-recognized 

11-year cycle related to sunspot activity. The 

amount of incoming solar radiation at a wavelength of 10.7 cm is often used as a surrogate 

for incoming solar radiation at UV wavelengths 

that produce stratospheric ozone. The 10.7-cm 

radiation values clearly show the recent periods 

of solar maximum and minimum. A comparison 

of the solar radiation and ozone changes strongly 

indicates that the cyclic changes in solar output 

alone cannot account for the long-term decrease 

in total ozone. Following large volcanic eruptions, transmission of solar radiation to Earth’s 

surface is reduced by the large number of new 

sulfur-containing particles formed in the stratosphere. The three large volcanic eruptions that 

occurred between 1960 and 2010 temporarily 

decreased solar transmission as measured in 

Hawaii. Volcanic particles increase ozone depletion only for a few years before they are removed 

from the stratosphere by natural processes. As a 

consequence, the two most recent volcanic eruptions cannot be the cause of the continuous longterm decrease found in global total ozone.

removed from the stratosphere by natural air circulation. As 

particles are removed, solar transmission is restored. Based 

on the short residence time of volcanic particles in the stratosphere, the two large eruptions in the past three decades cannot 

account directly for the continuous long-term decreases in 

global total ozone observed over the same period.

Reactive chlorine from volcanoes. Explosive volcanic 

plumes generally contain large quantities of reactive chlorine 

in the form of hydrogen chloride (HCl). HCl is a reactive halogen gas that can be converted to ClO, which rapidly destroys 

ozone (see Figure Q8-3). The plumes also contain a considerable amount of water vapor, which forms rainwater and ice in 

the rising fresh plume. Rainwater and ice efficiently scavenge 

and remove HCl while it is still in the lower atmosphere (troposphere). As a result, most of the HCl in explosive volcanic 

plumes does not enter the stratosphere. After recent explosive 

 Section III: STRATOSPHERIC OZONE DEPLETION 20 Questions: 2010 Update

100

50

150

200

250

1960 1970 1980 1990 2000 2010

Year

95

90

85

80

75

The Solar Cycle, Volcanic Eruptions, Global Ozone,

and Equivalent Effective Stratospheric Chlorine

Incoming solar

radiation

Solar transmission

at Hawaii (20°N)

Major volcanic

eruptions

Mt. Agung

(8°S)

El Chichón

(17°N)

Mt. Pinatubo

(15°N )

Change from 1964 –1980

average (%)

10.7– cm solar radiation

(relative amounts)

Transmission

(%)

 Annual averages

(seasonal and solar corrections)

 3-month averages

(no solar or volcanic corrections)

Global total ozone changes

2

0

-2

-4

-6

EESC

(relative amounts)

Equivalent effective

stratospheric

chlorine (EESC) 

Natural sources

1

0
Q.44

 20 Questions: 2010 Update Section III: STRATOSPHERIC OZONE DEPLETION

eruptions, observations of HCl in the stratosphere have 

confirmed that increases are small compared with the total 

amount of chlorine in the stratosphere from other sources.

Antarctic volcanoes. Volcanoes on the Antarctic continent are of special interest due to their proximity to the Antarctic ozone hole. An explosive eruption could in principle 

inject volcanic aerosol and small amounts of HCl directly into 

the stratosphere over Antarctica, which could lead to ozone 

depletion. However, to be a possible cause of the annually 

recurring ozone hole beginning in the early 1980s, explosive 

Antarctic eruptions would need to have occurred at least every 

few years to maintain volcanic emissions in the stratosphere. 

This is not the case. Only the Mt. Erebus volcano is currently 

active in Antarctica. No explosive eruptions of Mt. Erebus or 

any other Antarctic volcano have occurred since 1980. Therefore, explosive volcanic eruptions in the last three decades have 

not caused the Antarctic ozone hole and, as noted above, have 

not been sufficient to cause the long-term depletion of global 

total ozone.

Total ozone and 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 EESC present in the stratosphere in the early 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 EESC 

values were much lower. EESC values will remain substantial 

in the early decades of the 21st century even as ODSs decline 

globally, with 1980 values reached by about 2050 (see Figures 

Q16-1 and Q20-2). Large volcanic eruptions in the intervening years will cause more ozone depletion. If an explosive 

eruption larger than Mt. Pinatubo were to occur, peak ozone 

losses could be larger than previously observed and substantial ozone losses could persist for longer time periods. As 

halogen gas abundances gradually decline to 1980 values, the 

effect of volcanic eruptions on ozone will lessen.
Q.45

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 was a framework agreement that 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 sufficient country ratification, entered into force in 1989. The Protocol has been successful in establishing legally binding controls for developed 

and developing nations on the production and consumption 

of halogen source gases known to cause ozone depletion. 

Halogen source gases containing chlorine and bromine controlled under the Montreal Protocol are referred to as ozonedepleting substances (ODSs). National consumption of an 

ODS is defined as the amount that production and imports of 

the substance exceed its export to other nations. The Protocol 

provisions are structured for developed countries to act first 

and for developing countries to follow with some financial 

assistance. In 2010, the Montreal Protocol became the first 

international treaty to achieve universal ratification by all 196 

United Nations members.

Amendments and Adjustments. As the scientific basis 

of ozone depletion became more certain after 1987 and substitutes and alternatives became available to replace ODSs, 

the Montreal Protocol was strengthened with Amendments 

and Adjustments. Each Amendment is named after the city 

in which the meeting of the Parties to the Montreal Protocol 

took place and by the year of the meeting. The timeline in 

Figure Q0-1 shows some of the major decisions that have been 

adopted in the last two decades. These decisions have put 

additional ODSs under control, accelerated existing control 

measures, and prescribed phase-out dates for the production 

and consumption of certain gases. The initial Protocol called 

for a 50% reduction in chlorofluorocarbon (CFC) production 

and a freeze on halon production. The 1990 London Amendment called for a phase-out of the production and consumption of the most damaging ODSs in developed nations by 2000 

and in developing nations by 2010. The 1992 Copenhagen 

Amendment accelerated the phase-out date to 1996 in developed nations. Further controls on ODSs were agreed upon in 

later meetings in Vienna (1995), Montreal (1997, 2007), and 

Beijing (1999).

Influence of the Montreal Protocol. Montreal Protocol controls are based on several factors that are considered 

separately for each ODS. The factors include (1) the effectiveness in depleting ozone in comparison with other substances 

(see Ozone Depletion Potential (ODP) in Q18), (2) the availability of suitable substitutes for domestic and industrial 

use, and (3) the potential impact of controls on developing 

nations. The influence of Montreal Protocol provisions on 

stratospheric ODS abundances can be demonstrated with 

long-term changes in equivalent effective stratospheric chlorine (EESC). Calculations of EESC combine the amounts of 

chlorine and bromine present in surface air to form a measure 

of the potential for ozone destruction in a particular stratospheric region on an annual basis (see definition in Q16). 

The long-term changes in EESC at midlatitudes are shown in 

Figure Q15-1 for several cases:

▶ No Protocol. Without the Montreal Protocol the production, use, and emissions of CFCs and other ozonedepleting substances is expected to have increased after 

1987 with an annual growth rate of about 3% (businessas-usual scenario). As a result, EESC is projected to have 

increased nearly 10-fold by the mid-2050s compared with 

the 1980 value. Computer models of the atmosphere indicate that such high EESC values would have at least doubled global total ozone depletion between 1990 and 2010 

and increased it far beyond that by midcentury. As a result, 

Yes, the production and consumption of ozone-depleting substances are controlled under a 1987 international agreement 

known as the “Montreal Protocol on Substances that Deplete the Ozone Layer” and by its subsequent Amendments and 

Adjustments. The Protocol, now ratified by all 196 United Nations members, establishes legally binding controls on national 

production and consumption of ozone-depleting substances (ODSs). Production and consumption of all principal ODSs by 

developed and developing nations will be almost completely phased out before the middle of the 21st century.

Are there controls on the production of Q15 ozone-depleting substances?

Section IV: CONTROLLING OZONE-DEPLETING SUBSTANCES
Q.46

 20 Questions: 2010 Update Section IV: CONTROLLING OZONE-DEPLETING SUBSTANCES

harmful UV-B radiation would have increased substantially at Earth’s surface, causing a global rise in skin cancer 

and cataract cases (see Q17).

▶ Montreal Protocol provisions. International compliance with only the 1987 provisions of the Montreal Protocol and the later 1990 London Amendment would have 

substantially slowed the projected growth of EESC. Not 

until the 1992 Copenhagen Amendments and Adjustments 

did the Protocol projections show a decrease in future 

EESC values. The provisions became more stringent with 

the Amendments and Adjustments adopted in Beijing in 

1999 and Montreal in 1997 and 2007. Now, with full compliance to the Protocol, most ODSs will be almost completely phased out, with some exemptions for critical uses 

(see Q16). Global EESC is slowly decaying from its peak 

value in the late 1990s and is expected to reach 1980 values 

in the mid-21st century. The success of the Montreal Protocol to date is demonstrated by the decline in ODP-weighted 

emissions of ODSs shown in Figure Q0-1. Total emissions 

peaked in 1988 at values about 10-fold higher than natural 

emissions. Between 1988 and 2010, ODS emissions from 

human activities have decreased by over 80%.

▶ Zero emissions. EESC values in the coming decades 

will be influenced by (1) the slow natural removal of ODSs 

still present in the atmosphere, (2) emissions from continued production and use of ODSs, and (3) emissions from 

currently existing banks containing a variety of ODSs. 

ODS banks are associated with applications that involve 

long-term containment of halogen gases. Examples are 

CFCs in refrigeration equipment and insulating foams, 

and halons in fire-fighting equipment. New emissions are 

projected based on continued production and consumption of ODSs, particularly in developing nations, under 

existing Protocol provisions.

 The zero-emissions case demonstrates the EESC values 

that would occur if it were possible to set all ODS emissions to zero beginning in 2011. This would eliminate the 

contributions from new production and bank emissions. 

Significant differences from the Montreal 2007 projections 

are evident in the first decades following 2011 because the 

Figure Q15-1. Effect of the Montreal Protocol. The objective of the Montreal Protocol is 

the protection of the ozone layer through control of the global production and consumption 

of ODSs. Projections of the future abundances 

of ODSs expressed as equivalent effective 

stratospheric chlorine (EESC) values (see Q16) 

are shown separately for the midlatitude stratosphere for (1) no Protocol provisions, (2) the 

provisions of the original 1987 Montreal Protocol 

and some of its subsequent Amendments and 

Adjustments, and (3) zero emissions of ODSs 

starting in 2011. The city names and years indicate where and when changes to the original 

1987 Protocol provisions were agreed upon (see 

Figure Q0-1). EESC is a relative measure of the 

potential for stratospheric ozone depletion that 

combines the contributions of chlorine and bromine from ODS surface observations (see Q16). 

Without the Protocol, EESC values are projected 

to have increased significantly in the 21st century. Only with the Copenhagen (1992) and subsequent Amendments and Adjustments did projected EESC values show a long-term decrease.

1960 1980

0

2000 2020 2040 2060 2080 2100

Year

Montreal

1987

London

1990

Beijing

1999

Montreal

2007

Zero Emissions

in 2011

Copenhagen

1992

No Protocol

EESC (relative amounts)

Effect of the Montreal Protocol

Long-term changes in equivalent effective 

stratospheric chlorine (EESC)

Natural sources

1

3

2

4
Q.47

Section IV: CONTROLLING OZONE-DEPLETING SUBSTANCES 20 Questions: 2010 Update

phase-out of all ODS production under the Protocol is not 

yet complete and bank emissions are substantial. Zero 

emissions would bring forward the return of EESC to 1980 

levels by about 13 years.

HCFC substitute gases. The Montreal Protocol provides for the use of hydrochlorofluorocarbons (HCFCs) as 

transitional, short-term substitute compounds for ODSs with 

higher ODPs, such as CFC-12. HCFCs are used for refrigeration, in making insulating foams, and as solvents, all of 

which were primary uses of CFCs. HCFCs are generally more 

reactive in the troposphere than other ODSs because they 

contain hydrogen (H) in addition to chlorine, fluorine, and 

carbon. HCFCs are 88 to 98% less effective than CFC-12 in 

depleting stratospheric ozone because their chemical removal 

occurs primarily in the troposphere (see ODPs in Table Q7-1). 

This removal protects stratospheric ozone from most of the 

halogen content of HCFC emissions. In contrast, CFCs and 

many other ODSs release all of their halogen content in the 

stratosphere because they are chemically inert in the troposphere (see Q6).

Under the provisions of the Montreal Protocol, developed 

and developing countries may continue to use HCFCs as ODS 

substitutes in the coming decades before they are ultimately 

phased out. In the most recent Adjustment to the Protocol 

(Montreal 2007), the phase-out of HCFCs for all Parties was 

accelerated so that it will be complete in 2030, a decade earlier than in previous provisions. In adopting this decision, the 

Parties reduced the contribution of HCFC emissions to both 

long-term ozone depletion and future climate forcing (see Q18 

and Q19).

HFC substitute gases. Hydrofluorocarbons (HFCs) are 

used as long-term substitute compounds for CFCs, HCFCs, 

and other ODSs. HFCs contain hydrogen, fluorine, and 

carbon. HFCs do not contribute to ozone depletion because 

they contain no chlorine, bromine, or iodine. As a consequence, HFCs are not ODSs and are not subject to the provisions of the Montreal Protocol. HFCs and all ODSs are 

radiatively active gases that contribute to human-induced 

climate change based on their accumulation in the atmosphere (see Q18). HFCs are included in the basket of gases 

being controlled by the Kyoto Protocol of the United Nations 

Framework Convention on Climate Change (UNFCCC). The 

Kyoto Protocol is an international treaty designed to protect 

climate by controlling emissions of HFCs, carbon dioxide 

(CO2), methane (CH4), nitrous oxide (N2O), perfluorocarbons (PFCs), and sulfur hexafluoride (SF6). HFC emissions 

are expected to grow substantially in coming decades as the 

demand for their use as substitute gases and in new applications increases in developed and developing countries. The 

climate change contribution from future HFC emissions will 

be minimized if HFCs with very low Global Warming Potentials (GWPs) (less than 100) are chosen to meet the growing 

demand (see Q18).
Q.48

The Montreal Protocol and its Amendments and Adjustments have been very successful in reducing the atmospheric abundance of ozone-depleting substances (ODSs). 

ODSs are halogen source gases released by human activities. 

Under the Protocol, the production and consumption of individual ODSs are now controlled in all 196 nations that are 

Parties to the Protocol (see Q15). The success of the Montreal 

Protocol controls is documented by (1) observed changes and 

future projections of the atmospheric abundance of the principal ODSs and (2) long-term values of equivalent effective 

stratospheric chlorine (EESC).

Individual ODS reductions. The reduction in the atmospheric abundance of an ODS in response to controls on production and consumption depends principally on two factors: 

(1) how rapidly an ODS is used and released to the atmosphere 

after being produced and (2) the lifetime for the removal of the 

ODS from the atmosphere (see Table Q7-1). For example, the 

abundances of ODSs with short lifetimes respond quickly to 

emission reductions. Long-term changes in ODS abundances 

are constructed from (1) estimates of historical emissions 

and bank quantities using industry reports, (2) abundances 

measured in air trapped for years in accumulated snow (firn) 

in polar regions, (3) observed atmospheric abundances using 

ground-based measurements, and (4) projections of future 

abundances based on compliance with Montreal Protocol 

provisions and patterns of use in developed and developing 

countries. The results for individual ODSs and the natural 

chlorine source gas, methyl chloride (CH3Cl), are shown in 

Figure Q16-1 and described as follows:

▶ CFCs. Chlorofluorocarbons (CFCs) include some of the 

most destructive chlorine-containing ODSs. CFC-11 and 

CFC-12, each with an Ozone Depletion Potential (ODP) 

near 1, are the most abundant ODSs in the atmosphere 

owing to large historical emissions and long atmospheric 

lifetimes (45–100 years). Production and consumption of 

CFCs in developed countries ended in 1996 and that in 

developing countries ended by January 2010. As a consequence, CFC-11 and CFC-113 abundances have peaked in 

the atmosphere and have been declining for more than a 

decade. In contrast, CFC-12 abundances have only recently 

shown a decrease, owing to its longer lifetime (100 years) 

and continuing emissions from CFC-12 banks, namely, 

refrigeration and air conditioning equipment and thermal 

insulating foams. With no further global production of the 

principal CFCs allowed except for limited exempted uses 

and with some continuing emissions from banks, CFC 

abundances are projected to decline steadily throughout 

this century.

▶ Halons. Halons are the most destructive brominecontaining ODSs. The most abundant in the atmosphere, 

halon-1211 and halon-1301, have abundances about 100 

times less than CFC-11 and CFC-12 and account for a 

significant fraction of bromine from all ODSs (see Figure 

Q7-1). Production and consumption of halons in developed countries ended in 1994 and that in developing countries ended by January 2010. In the 2005–2008 period, 

atmospheric abundances of halon-1211 showed a significant decrease for the first time. Halon-2402 abundances 

have been decreasing for a few years while those of halon1301 continue to increase. The increase is likely due to substantial banks in fire-extinguishing and other equipment 

that gradually release halon-1301 to the atmosphere. The 

abundance of halon-1301 is expected to remain high well 

into the 21st century because of its long lifetime (65 years) 

and continued release.

▶ Methyl chloroform. The largest reduction to date in the 

abundance of an ODS (about 90% from its peak value) has 

been observed for methyl chloroform. Production and 

consumption of methyl chloroform in developed countries 

ended in January 1996 and that in developing countries 

Yes, as a result of the Montreal Protocol, the overall abundance of ozone-depleting substances (ODSs) in the atmosphere 

has been decreasing for about a decade. If the nations of the world continue to comply with the provisions of the Montreal 

Protocol, the decrease will continue throughout the 21st century. Those gases that are still increasing in the atmosphere, such 

as halon-1301 and HCFC-22, will begin to decrease in the coming decades if compliance with the Protocol continues. Only 

after midcentury will the effective abundance of ODSs fall to values that were present before the Antarctic ozone hole was 

observed in the early 1980s.

Has the Montreal Protocol been successful in reducing Q16 ozone-depleting substances in the atmosphere?

 20 Questions: 2010 Update Section IV: CONTROLLING OZONE-DEPLETING SUBSTANCES
Q.49

Section IV: CONTROLLING OZONE-DEPLETING SUBSTANCES 20 Questions: 2010 Update

Figure Q16-1. Halogen source gas changes. The atmospheric abundances of individual ozone-depleting substances 

(ODSs) at Earth’s surface were obtained using a combination of direct atmospheric measurements, estimates of historical abundance, and future projections of abundance. The gases shown are all ODSs except for methyl chloride. The past increases 

of CFCs, along with those of carbon tetrachloride, methyl chloroform, and halon-1211, have slowed and reversed in the last 

two decades. HCFCs, which are used as transitional substances to replace CFCs, will likely continue to increase in the coming 

decades before a complete phase-out. Halon-1301 abundances will also likely continue to grow in the next decade while current halon banks are depleted. Future decreases in methyl bromide are expected to be small. Abundances of methyl chloride, 

which is not controlled under the Montreal Protocol and has large natural sources, are projected to be constant in the future.

The rise in equivalent effective stratospheric chlorine (EESC) values in the 20th century has slowed and reversed in the last 

decade (top left panel). EESC is a measure of the potential for ozone depletion in the stratosphere derived from surface measurements of chlorine and bromine gases (see Q16). EESC values are reduced as ODS emissions decrease. EESC, as shown here 

for the stratosphere at midlatitudes, returns to 1980 values around 2050. In polar regions, the return to 1980 values occurs 

almost 10 years later. International compliance with the Montreal Protocol provisions will ensure that EESC values will continue to decrease as projected (see Q15). (The unit “parts per trillion” is defined in the caption of Figure Q7-1.)

Methyl bromide (CH3Br)

Methyl chloride (CH3Cl)

Halon-1211

Halon-1301

CFC-12

HCFC-22 CFC-11

CFC-113

Halon-2402

Estimates

Observations

Future projections

Carbon tetrachloride

CCl4

Methyl chloroform

CH3CCl3

HCFC-141b

HCFC-142b

Past and Projected Atmospheric Abundances of Halogen Source Gases

2

1

0 0

0

4

3

6

8

10

12

2

1

500

400

1950 2000 2050 2100 1950 2000 2050 2100

Year Year

EESC (relative amount) Atmospheric abundance (parts per trillion)

Atmospheric abundance (parts per trillion) 100

200

300

400

600

500

Midlatitudes

1980

level

CH3Br (natural + human sources)

CH3Cl (natural sources)

CFCs

Halons

HCFCs

CCl4

CH3CCl3

0

20

40

60

80

100

120

300

Equivalent effective stratospheric

chlorine (EESC)
Q.50

 20 Questions: 2010 Update Section IV: CONTROLLING OZONE-DEPLETING SUBSTANCES

is scheduled to end by January 2015. Atmospheric abundances responded rapidly to the reduced emissions starting 

in the mid-1990s 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. It will approach complete 

removal from the atmosphere after the phase-out in developing countries is complete.

▶ HCFC substitute gases. The Montreal Protocol allows 

for the use of hydrochlorofluorocarbons (HCFCs) as shortterm, transitional substitutes for CFCs and in other specific applications. As a result, the abundances of HCFC22, HCFC-141b, and HCFC-142b continue to grow in the 

atmosphere with rates that have remained constant or 

increased in recent years in response to greater application demand. HCFCs pose a lesser threat to the ozone 

layer than CFCs because they have low ODP values (less 

than 0.12). The most recent Adjustment to the Montreal 

Protocol (Montreal 2007) accelerated the phase-out of 

HCFCs by a decade for both developed (2020) and developing countries (2030) (see Q15). Even with the accelerated 

phase-out, future projections show HCFC abundances that 

continue to increase, reach peak values before the mid-21st

century, and steadily decrease thereafter. The response of 

atmospheric abundances to decreasing emissions will be 

relatively rapid because of short atmospheric lifetimes of 

HCFCs (less than 20 years).

▶ Carbon tetrachloride. Carbon tetrachloride has been 

phased out in both developed countries (January 1996) 

and developing countries (January 2010). As a result, 

atmospheric abundances of carbon tetrachloride have been 

decreasing for two decades. The decline is somewhat less 

rapid than expected, suggesting that actual emissions are 

larger than reported or the atmospheric lifetime is greater 

than current estimates. Production of carbon tetrachloride 

for use as raw material (feedstock) to make other chemicals 

is exempted under the Protocol because emission to the 

atmosphere does not occur in this case.

▶ Methyl chloride and methyl bromide. Both methyl 

chloride and methyl bromide are distinct among halogen 

source gases because substantial fractions of their emissions 

are associated with natural processes (see Q7). Methyl chloride is not an ODS under the Montreal Protocol because it 

is not manufactured or used in appreciable quantities. Its 

abundance in the atmosphere has remained fairly constant 

throughout the last 60 years and will remain constant if 

the balance of its natural production and loss processes 

remains unchanged. Methyl bromide is controlled under 

the Protocol because it is manufactured for use as a fumigant. Developed country production and consumption of 

methyl bromide ended in January 2005 and that in developing countries is scheduled to end by January 2015. The 

Protocol currently provides exemptions for some methyl 

bromide production and use as an agricultural and preshipment fumigant. Atmospheric abundances of methyl 

bromide responded rapidly to the reduced emissions starting in 1999 because its atmospheric removal lifetime is less 

than 1 year. Future projections show only small changes 

in methyl bromide abundances based on the assumptions 

of unchanged contributions from natural sources and small 

use in developing countries. An important uncertainty in 

these projections is the future amounts that will be produced 

and emitted under Montreal Protocol use exemptions.

Equivalent effective stratospheric chlorine (EESC). 

Important measures of the success of the Montreal Protocol 

are the past and projected changes in the values of equivalent 

effective stratospheric chlorine (EESC), which was introduced 

in Figures Q14-1 and Q15-1. EESC is designed as one measure of the potential for ozone depletion in the stratosphere 

that can be calculated from atmospheric surface abundances 

of ODSs and natural chlorine and bromine gases. All gases 

used in the calculation are shown in Figure Q7-1. For both 

past and future EESC values, the required atmospheric abundances are derived from measurements, historical estimates, 

or future projections based on compliance with Montreal 

Protocol provisions.

EESC is expressed as a hypothetical amount of chlorine 

available in the stratosphere to deplete ozone. The term equivalent indicates that bromine gases, scaled by their greater peratom effectiveness in depleting ozone, are included in EESC. 

The term effective indicates that only the estimated fraction of 

ODSs that are currently in the form of reactive halogen gases 

in the stratosphere is included in an EESC value (see Q8). 

Although chlorine is much more abundant in the stratosphere 

than bromine (about 150-fold) (see Figure Q7-1), bromine 

atoms are about 60 times more effective than chlorine atoms 

in chemically destroying ozone in the lower stratosphere. 

EESC generally depends on the year and latitude region in the 

stratosphere being considered.

Another quantitative measure of the potential for ozone 

depletion in the stratosphere that is highly related to EESC 

is effective stratospheric chlorine (ESC). ESC, as defined in 

Q20, is calculated with chemistry-climate models rather than 

directly from ODS observations.
Q.51

Section IV: CONTROLLING OZONE-DEPLETING SUBSTANCES 20 Questions: 2010 Update

Long-term changes in EESC. In the latter half of the 

20th century up until the 1990s, EESC values steadily increased 

(see Figure Q16-1), causing global ozone depletion. As a result 

of the Montreal Protocol regulations, the long-term increase 

in EESC slowed, values reached a peak, and EESC began to 

decrease in the 1990s. The initial decrease came primarily 

from the substantial, rapid reductions in emissions of methyl 

chloroform, which has a lifetime of only 5 years. The decrease 

is continuing with declining emissions of CFCs and other 

long-lived ODSs. Decreasing EESC means that the potential 

for stratospheric ozone depletion is now lessening each year 

as a result of the Montreal Protocol. Decreases in EESC are 

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 are reduced, 

natural processes continue to gradually remove halogencontaining gases from the global atmosphere. Reduction of 

EESC values to 1980 values or lower will require several more 

decades because the most abundant ODS molecules now in the 

atmosphere have lifetimes ranging from 10 to 100 years.
Q.52

The depletion of stratospheric ozone leads to an increase in solar ultraviolet radiation at Earth’s surface. 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, which 

is invisible to the human eye. Long-term changes in UV-B 

radiation reaching the surface have been measured directly 

and can be estimated from total ozone changes.

UV-B radiation can harm humans, other life forms, and 

materials (see Q3). Most of the effects of sunlight on the 

human body are caused by UV-B exposure. A principal effect 

is skin erythema, which leads to sunburning. Excess exposure 

may lead to skin cancers. Erythemal radiation is regularly 

reported to the public in many countries in the form of the 

“UV Index.” The long-term changes in surface UV-B radiation are important to study because of its potential harmful 

effects and relationship to ozone depletion.

Surface UV-B radiation. The amount of UV-B radiation 

reaching Earth’s surface at a particular location depends in 

large part on total ozone at that location. Ozone molecules in 

the stratosphere and in the troposphere absorb UV-B radiation, thereby significantly reducing the amount that reaches 

Earth’s surface (see Q3). If conditions occur that reduce the 

abundance of ozone molecules somewhere in the troposphere 

or stratosphere, total ozone is reduced and the amount of 

UV-B radiation reaching Earth’s surface below is increased 

proportionately. This relationship between total ozone and 

surface UV-B radiation has been confirmed at a variety of 

locations with direct measurements of both quantities.

Additional causes of UV changes. The actual amount 

of UV-B radiation reaching the Earth’s surface at a specific 

location and time depends on a number of factors in addition 

to total ozone. The primary additional factor is the position 

of the Sun in the sky, which changes with daily and seasonal 

cycles. Other factors include local cloudiness, the altitude of 

the location, the amount of ice or snow cover, and the amounts 

of atmospheric particles (aerosols) in the atmosphere above the 

location. Changes in clouds and aerosols are partially related 

to air pollution and greenhouse gas emissions from human 

activities. Measurements indicate that both increases and 

decreases in UV radiation at certain locations have resulted 

from changes in one or more of these factors. Estimating the 

impact of changes in these factors is complex. For example, 

an increase in cloud cover usually results in a reduction of 

UV radiation below the clouds and could at the same time 

increase surface radiation in any nearby mountainous regions 

above the clouds.

Long-term surface UV changes. Long-term changes 

in UV-B radiation have been estimated from measurements 

made with special UV monitoring instruments at several 

surface locations since about 1990. For example, as a consequence of Antarctic ozone depletion, the average UV-B measured at the South Pole during spring between 1991 and 2006 

was 55–85% larger than estimated for the years 1963–1980. 

In addition, satellite observations of ozone changes have been 

used to estimate changes in surface UV-B radiation that have 

occurred over the past three decades. With satellite observations, the UV-B radiation changes can be separately attributed to changes in ozone and clouds. The results show that 

erythemal radiation has increased by up to 6% between 1979 

and 2008 over a wide range of latitudes outside the tropics 

(see lower panel of Figure Q17-1). The largest percentage 

increases have occurred at high polar latitudes in both hemispheres, where the largest annual decreases in total ozone 

are observed (see Figure Q13-1). Over this time period the 

UV increases due to ozone depletion are partially offset by 

changes in cloudiness, primarily in the high latitudes of the 

Southern Hemisphere (see top panel in Figure Q17-1). Without changes in cloudiness, the increases in erythemal radiation at these latitudes would have reached a maximum close 

to 9%. The smallest changes in erythemal UV have been in 

the tropics, where long-term changes in total ozone are smallest (see Q13). In the tropics and in the Northern Hemisphere 

the average changes in clouds during this period were very 

small. As a result, the net increases in erythemal radiation in 

these regions are determined primarily by ozone depletion.

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 

provide evidence that surface ultraviolet radiation has increased in large geographic regions in response to ozone depletion.

Q17 Does depletion of the ozone layer increase 

ground-level ultraviolet radiation?

Section V: IMPLICATIONS OF OZONE DEPLETION AND THE MONTREAL PROTOCOL
Q.53

UV Index changes. The UV Index is a measure of the 

erythemal radiation that occurs at a particular surface location and time. The index is used internationally to increase 

public awareness about the detrimental effects of UV on 

human health and to guide the need for personal protective 

measures. The maximum daily UV Index varies with location and season, as shown for three sites in Figure Q17-2. The 

UV Index increases when moving from high to low latitudes 

and is highest in summer when the midday Sun is closest 

to overhead. UV Index values in San Diego, California, at 

32°N, for example, are generally higher than those in Barrow, 

Alaska, at 71°N. At all latitudes, UV Index values increase in 

mountainous regions and over snow- or ice-covered regions. 

The UV Index is zero during periods of continuous darkness 

in winter at high-latitude locations.

The UV Index over Antarctica has increased dramatically due to ozone depletion, as illustrated in Figure Q17-2. 

Normal index values for Palmer, Antarctica, at 64°S in spring 

were estimated from satellite measurements made during the 

period 1978–1980, before the appearance of the ozone hole 

over Antarctica. In the period from 1991 to 2006, the severe 

and persistent ozone depletion that occurred in late winter and 

early spring over Antarctica increased the average UV Index 

well above normal values for several months. Now, the spring 

UV Index measured in Palmer, Antarctica, routinely equals 

or exceeds spring and summer values measured in San Diego, 

California, located at a much lower latitude (32°N).

UV changes and human health. Over the past several 

decades, depletion of the stratospheric ozone layer together 

with societal changes in lifestyle have increased UV-B radiation exposure for many people. Increased exposure has 

adverse health effects, primarily associated with eye and skin 

disorders. UV radiation is a recognized risk factor for some 

types of cataracts. For the skin, the most common threat is 

Section V: IMPLICATIONS OF OZONE DEPLETION AND THE MONTREAL PROTOCOL 20 Questions: 2010 Update

Figure Q17-1. Long-term changes in surface erythemal radiation. Solar ultraviolet 

(UV) radiation at Earth’s surface has increased 

over much of the globe since 1979. Erythemal 

radiation, which can lead to sunburning, is a 

component of surface UV radiation that is 

harmful to humans and other life forms. Surface erythemal radiation responds to changes 

in total ozone as well as clouds and aerosols. 

Satellite observations have been used to 

estimate the long-term changes of erythemal 

radiation due to changes in ozone, clouds, and 

aerosols. These estimates show that erythemal radiation at Earth’s surface has increased 

over much of the globe over the period 1979 

to 2008, particularly at midlatitudes in both 

hemispheres (bottom panel). The increases 

in the Southern Hemisphere would have been 

larger without the offsetting changes due to 

clouds and aerosols (upper panel). The smallest estimated changes in erythemal radiation 

are in the tropics because observed total 

ozone changes over this period are smallest 

there (see Figure Q13-1). The shaded areas 

surrounding the figure lines represent the 

uncertainties in the estimated changes.

Changes in Surface Erythemal UV Radiation

Changes between 1979 and 2008

Latitude (degrees)

Average change (%) Average change (%)

8

6

4

2

0

10

12

8

6

4

2

0

-2

-4

0

-2

-60 -40 -20 20 40 60

South North

. . . from combined

 ozone, cloud, and

 aerosol changes

. . . from total

 ozone changes

Averages

Uncertainty ranges

. . . from cloud and

aerosol changes

Average

Uncertainty range
Q.54

 20 Questions: 2010 Update Section V: IMPLICATIONS OF OZONE DEPLETION AND THE MONTREAL PROTOCOL

Figure Q17-2. Long-term changes in the 

UV Index. The UV Index is a measure of the 

erythemal radiation that reaches the surface 

at a particular location and time. Sunburning in humans is caused by erythemal radiation. The UV Index varies with total ozone 

above a location, as well as the Sun’s midday 

position in the local sky, which depends on 

latitude and season. The highest values of 

the UV Index occur in the tropics, where the 

midday Sun is highest throughout the year 

and where total ozone values are low. The 

lowest average UV Index values occur at high 

latitudes. The figure shows daily maximum 

UV Index values at several locations. Values 

are higher throughout the year in San Diego, 

a low-latitude location, than in Barrow, a 

high-latitude location. UV Index values are 

zero at high latitudes in winter when darkness 

is continuous. Measurements at Palmer, Antarctica, demonstrate the effect of Antarctic 

ozone depletion. Normal values estimated 

for Palmer are shown for 1978–1980, a period 

before the appearance of the “ozone hole” 

(see red dashed line). In the period 1990–

2006, Antarctic ozone depletion increased 

the UV Index at Palmer throughout spring and into summer (see yellow shaded region). Springtime values at Palmer now 

routinely equal or exceed those measured in spring and summer in San Diego, which is located at much lower latitude.

Seasonal Changes in the UV Index

Southern Hemisphere months

Northern Hemisphere months

Winter Spring Summer Fall

J F M A M J J A S O N D

J J F M A M J

16

14

12

10

8

6

4

2

0

A S O N D

Daily UV Index

Palmer, Antarctica (64°S) San Diego, California (32°N)

Barrow, Alaska (71°N)

1991– 2007

1990 – 2006 1992 – 2006

1978 – 1980

(pre-ozone-hole)

Long-term averages of the UV Index

skin cancer. Over the past decades, the incidence of several 

types of skin tumors has risen significantly among people of 

all skin types. Skin cancer in humans occurs long after exposure to sunburning UV. With current Montreal Protocol provisions, projections of additional skin cancer cases associated 

with ozone depletion are largest in the early to middle decades 

of the 21st century and represent a significant global health 

issue. An important human health benefit of UV-B radiation exposure is the production of vitamin D, which plays a 

significant role in bone metabolism and the immune system. 

Human exposure to solar UV-B radiation requires a careful 

balance to maintain adequate vitamin D levels while minimizing long-term risks of skin and eye disorders.
Q.55

Figure Q18-1. Radiative forcing of greenhouse gases. Human activities since the 

start of the Industrial Era (around 1750) have 

caused increases in the abundances of several 

short-lived and long-lived gases, known as 

greenhouse gases, that all contribute to the 

radiative forcing of climate, also known as 

climate forcing. Radiative forcing is expressed 

in units of watts per square meter (W/m2). As 

shown in the figure, the largest forcings are 

those of carbon dioxide (CO2), followed by 

methane (CH4), tropospheric ozone, halocarbon gases, and nitrous oxide (N2O). The black 

whiskers on each bar show uncertainties in 

the values. Tropospheric ozone increases 

result from the emission of pollutant gases 

and create a positive ozone forcing. Positive 

forcings lead 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. Halocarbons 

include all ODSs, their substitutes, and a few 

other gases (see Figure Q18-2). In the coming decades, ODS abundances and stratospheric ozone depletion are expected to 

be reduced, along with their associated radiative forcings.

While stratospheric ozone depletion is not the principal cause of climate change, aspects of ozone depletion 

and climate change are closely linked. Both processes involve 

gases released to the atmosphere by human activities. The 

links are best understood by examining the contribution to 

climate change of the gases involved: ozone; ozone-depleting 

substances (ODSs) (or halogen source gases) and their substitutes; and other leading greenhouse gases.

Greenhouse gases and the radiative forcing of 

climate. The warming of the Earth by the Sun is enhanced 

by the presence of natural greenhouse gases, of which water 

vapor is an important example. Without this natural greenhouse effect, the Earth’s surface would be much colder. 

Human activities since the preindustrial era have led to longterm increases in the atmospheric abundances of a number 

of long-lived and short-lived greenhouse gases. This group 

No, ozone depletion itself is not the principal cause of climate change. Changes in ozone and climate are directly linked 

because ozone absorbs solar radiation and is also a greenhouse gas. Stratospheric ozone depletion and increases in global 

tropospheric ozone that have occurred in recent decades have opposing contributions to climate change. The ozonedepletion contribution, while leading to surface cooling, is small compared with the contribution from all other greenhouse 

gas increases, which leads to surface warming. The total forcing from these other greenhouse gases is the principal cause 

of observed and projected climate change. Ozone depletion and climate change are indirectly linked because both ozonedepleting substances and their substitutes are greenhouse gases.

Is depletion of the ozone layer the principal cause

 of climate change?

Section V: IMPLICATIONS OF OZONE DEPLETION AND THE MONTREAL PROTOCOL 20 Questions: 2010 Update

Q18

Radiative Forcing of Climate Change

Radiative forcing (W/m2)

Surface cooling Surface warming

-1 0 1 2

Carbon dioxide (CO2)

Nitrous oxide (N2O)

Methane (CH4)

Halocarbons

Tropospheric

ozone increase

Stratospheric

ozone decrease

From changes in greenhouse gases caused by 

human activities between 1750 and 2005
Q.56

Figure Q18-2. Halocarbons and radiative forcing of climate change. Halocarbon gases in the atmosphere represent 

an important contribution to the radiative forcing of climate between 2005 and the preindustrial era (see Figure Q18-1). Halocarbons are all gases containing chlorine, bromine, or fluorine atoms that are now controlled as ozone-depleting substances 

(ODSs) by the Montreal Protocol or as climate change gases by the Kyoto Protocol (see color shading). Shown in the figure 

are the separate contributions of each gas or group of gases, as estimated using atmospheric abundance histories and Global 

Warming Potentials (GWPs) (see Figure Q18-3). The gases listed in the right hand labels begin with the largest contribution in 

each group, except for CFC-13, CFC-114, CFC-115, and halons, which are shown as one total value. The individual forcing terms 

add together to form the bottom bar representing the total halocarbon forcing. The forcings of CFC-11 and CFC-12, the largest 

halocarbon contributions, are already decreasing and will continue to decrease as CFCs are gradually removed from the atmosphere (see Figure Q16-1). In contrast, the contributions of the intermediate-term ODS substitute gases, HCFCs, are projected 

to grow for another two decades before decreasing. The future contributions of the long-term ODS substitute gases, HFCs, are 

also expected to increase. In this case, the total contribution will depend strongly on which HFCs are used because the GWPs 

of individual HFCs vary widely (see Figure Q18-3).

20 Questions: 2010 Update Section V: IMPLICATIONS OF OZONE DEPLETION AND THE MONTREAL PROTOCOL

Radiative Forcing of Climate by Halocarbons in 2005

From increases in all controlled gases containing chlorine, bromine,

and fluorine from human activities between 1750 and 2005

Major CFCs

CFC-12,

 CFC-11,

 CFC-113

Other ODSs:

Carbon tetrachloride

Methyl chloroform

CCl4,

 CH3CCl3

Minor CFCs

 and halons

(CFC-13, -114, -115

and halons)

Hydrochlorofluorocarbons

(HCFCs)

HCFC-22,

 HCFC-142b,

 HCFC-141b

HFCs,

 PFCs,

 SF6

Hydrofluorocarbons (HFCs)

Perfluorocarbons (PFCs)

Sulfur hexafluoride (SF6)

Total halocarbon

climate forcing

Radiative forcing (W/m2)

Increasing surface warming

0 0.1 0.2 0.3 0.4

Montreal Protocol Gases

Kyoto Protocol Gases
Q.57

includes stratospheric and tropospheric ozone, halocarbons, 

carbon dioxide (CO2), methane (CH4), and nitrous oxide 

(N2O). ODSs and their substitutes make up a large fraction 

of the halocarbons in today’s atmosphere. The natural abundances of these gases in Earth’s atmosphere change the balance between incoming solar radiation and outgoing infrared 

radiation, warming the atmosphere and surface. Increases 

in the abundance of these gases from human activities cause 

more outgoing radiation to be absorbed, which further warms 

the atmosphere and surface. This change in Earth’s radiative 

balance caused by human activities is called a radiative forcing 

of climate or, more simply, a climate forcing. The magnitude 

of this energy imbalance is usually evaluated at the top of the 

troposphere (tropopause) and is expressed using units of watts 

per square meter (W/m2). The potential for climate change 

increases as this radiative forcing increases.

A summary of radiative forcings in 2005 resulting from the 

increases in the principal long-lived and short-lived greenhouse gases during 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. Climate forcings also 

lead to other changes, such as in precipitation patterns and 

extreme weather events. International climate assessments 

conclude that much of the observed surface warming and 

changes in other climate parameters over the last decades is 

due to increases in the abundance of carbon dioxide and other 

greenhouse gases caused by human activities.

Stratospheric and tropospheric ozone. Stratospheric 

and tropospheric ozone both absorb infrared radiation emitted by Earth’s surface, trapping heat in the atmosphere. Stratospheric ozone also significantly absorbs solar radiation. As a 

result, increases or decreases in stratospheric or tropospheric 

ozone induce a climate forcing and, therefore, represent direct 

links between ozone and climate. In recent decades, global 

stratospheric ozone has decreased due to rising reactive chlorine and bromine amounts in the atmosphere, while global 

tropospheric ozone in the Industrial Era has increased due 

to pollution from human activities (see Q3). Stratospheric 

ozone depletion has caused a small negative radiative forcing since preindustrial times, while increases in tropospheric 

ozone have caused a positive radiative forcing (see Figure Q18-

1). Summing the positive forcing due to tropospheric ozone 

increases with the smaller negative forcing from stratospheric 

ozone depletion yields a net positive radiative forcing. The 

large uncertainty in tropospheric ozone forcing reflects the 

difficulty in quantifying tropospheric ozone trends and in 

modeling the complex production and loss processes that control its abundance. The negative radiative forcing from stratospheric ozone depletion will diminish in the coming decades 

as ODSs are gradually removed from the atmosphere.

Stratospheric ozone depletion cannot be a principal cause 

of present-day global climate change for two reasons: first, the 

climate forcing from ozone depletion is negative, which leads 

to surface cooling. Second, the total forcing from other longlived and short-lived gases in Figure Q18-1 is positive and far 

larger. The total forcing from these other gases is the principal cause of observed and projected climate change.

Carbon dioxide, methane, and nitrous oxide. The 

accumulation of carbon dioxide during the Industrial Era 

represents the largest climate forcing related to human activities. Carbon dioxide concentrations continue to increase in 

the atmosphere primarily as the result of burning fossil fuels 

(coal, oil, and natural gas) for energy and transportation, as 

well as from cement manufacturing. The atmospheric abundance of carbon dioxide in 2005 was about 36% above what 

it was 260 years ago in preindustrial times. Carbon dioxide 

is considered a long-lived gas, since a significant fraction 

remains in the atmosphere 100–1000 years after emission.

Methane is a short-lived climate gas (atmospheric lifetime 

of about 10 years) that has both human and natural sources. 

Human sources include livestock, rice agriculture, and landfills. Natural sources include wetlands, oceans, and forests.

Nitrous oxide is a long-lived climate gas (atmospheric lifetime of about 110 years) that also has both human and natural 

sources. The largest human source is agricultural activities, 

especially related to fertilization. Microbial processes in soils 

that are part of natural biogeochemical cycles represent the 

largest natural source. In the stratosphere, nitrous oxide is 

the principal source of reactive nitrogen species, which participate in ozone destruction cycles (see Q2 and Q7).

Halocarbons. Halocarbons in the atmosphere contribute 

to both ozone depletion and climate change. As used here, halocarbons represent those gases containing chlorine, bromine, or 

fluorine atoms that are now controlled substances under the 

Montreal Protocol or the Kyoto Protocol. ODSs are the halocarbons controlled under the Montreal Protocol. HFC substitute 

gases, perfluorocarbons (PFCs), and sulfur hexafluoride (SF6) 

are controlled under the Kyoto Protocol. In 2005, the halocarbon contribution to climate forcing was 0.34 W/m2

, which is the 

third or fourth largest following carbon dioxide and methane 

(see Figure Q18-1). The contributions of individual halocarbon 

gases are highlighted in Figure Q18-2. Within the halocarbons, 

CFCs contribute the largest percentage (80%) to the 2005 climate 

Section V: IMPLICATIONS OF OZONE DEPLETION AND THE MONTREAL PROTOCOL 20 Questions: 2010 Update
Q.58

20 Questions: 2010 Update Section V: IMPLICATIONS OF OZONE DEPLETION AND THE MONTREAL PROTOCOL

forcing. HCFCs, the intermediate-term ODS substitutes, make 

the next largest contribution (12%). The atmospheric abundance of HFCs, the longer-term ODS substitutes, contributes 

only 3% to the 2005 halocarbon climate forcing.

The large contribution of the CFCs is expected to gradually 

decrease following the projected decline in their atmospheric 

abundance (see Figure Q16-1). Based on their long lifetimes, 

CFCs will still make a significant contribution, and most 

likely the largest ODS contribution, to halocarbon climate 

forcing at the end of the 21st century. Halocarbons controlled 

under the Kyoto Protocol (HFCs, PFCs, and SF6) represent 

about 5% of halocarbon climate forcing in 2005. With the 

projected growth of HFC production and consumption in 

developing nations, this percentage contribution is expected 

to increase substantially in the coming decades.

Ozone Depletion Potentials and Global Warming 

Potentials. An important way of comparing the influence 

of individual halocarbons on ozone depletion and climate 

change is to use Ozone Depletion Potentials (ODPs) and 

Global Warming Potentials (GWPs). The ODP and GWP of 

a gas quantify its effectiveness in causing ozone depletion and 

climate forcing, respectively (see Table Q7-1). The principal 

halocarbon gases are contrasted with each other in Figure 

Q18-3. The ODP of CFC-11 and the GWP of carbon dioxide are assigned reference values of 1. The CFCs and carbon 

tetrachloride all have ODPs near 1, indicating comparable 

Figure Q18-3. ODPs and GWPs. ODSs and their substitutes can be compared via their Ozone Depletion Potentials (ODPs) 

and Global Warming Potentials (GWPs) (see Table Q7-1). Larger ODPs or GWPs indicate greater potential for ozone depletion 

or climate warming, respectively, when a gas is emitted to the atmosphere. The ODP and GWP values are derived assuming 

an equal mass of each gas is emitted. The GWPs shown here are evaluated for a 100-year time interval after emission. The 

ODP of CFC-11 and the GWP of CO2 are assigned reference values of 1.0. The CFCs, halons, and HCFCs are ozone-depleting 

substances (see Q7) while HFCs, used as ODS substitutes, do not destroy ozone (ODPs equal 0). The ODPs of the halons far 

exceed those of the CFCs. All ODSs and their substitutes shown here have a non-zero GWP, with values spanning the wide 

range of 4 to 14,000.

Evaluation of Selected Ozone-Depleting Substances and Substitute Gases

Relative importance of equal mass emissions for ozone depletion and climate change

0 0 5 10 15 20 4000 8000 12000 16000

Ozone Depletion Potential (ODP) Global Warming Potential (GWP, 100-yr)

Ozone-depleting substances Substitute gases

HFC-143a

HFC-23

HFC-125

HFC-152a

Halon-1211

CFC-11

HCFC-22

Methyl bromide

Carbon tetrachloride

(0)

(0)

(0)

(0)

(0.05)

(0)

Hydrofluorocarbons

(HFCs)

Bromine-containing

gases

(133)

(5)

(1)

Increasing ozone depletion Increasing surface warming 

HFC-134a

Halon-1301

CFC-113

CFC-12

Chlorine-containing

gases

(0) HFC-1234yf (4)
Q.59

Section V: IMPLICATIONS OF OZONE DEPLETION AND THE MONTREAL PROTOCOL 20 Questions: 2010 Update

effectiveness in causing ozone depletion. The principal halons 

have ODPs greater than 7, making them the most effective 

ozone-depleting substances. HFCs have ODPs of zero since 

they cause no ozone depletion (see Q7).

All halocarbons have non-zero GWPs and, therefore, contribute to climate forcing. The GWP does not correspond 

strongly with the ODP of a gas because these quantities 

depend on different chemical and physical properties. For 

example, while HFC-134a does not destroy ozone (ODP equal 

0), each gram emitted is 1,370 times more effective than a 

gram of carbon dioxide in causing climate forcing. The future 

selection of specific HFCs as ODS substitutes or for use in new 

global applications will have important consequences for climate forcing. When these HFCs are eventually released to the 

atmosphere, the contribution to climate forcing will depend on 

their GWPs, which could vary over a wide range (4 to 14,000).

Montreal Protocol regulations have led to reductions in 

CFC emissions and increases in HCFC emissions (see Q16). 

As a result of these actions, the total radiative forcing from 

ODSs is slowly decreasing (see Q19). Overall halocarbon 

radiative forcing, however, is slowly increasing because of 

growing contributions from HFCs, PFCs, and SF6. It is 

important to note that, despite having a GWP that is small in 

comparison to many other halocarbons and other greenhouse 

gases, carbon dioxide is the most important greenhouse gas 

related to human activities because its emissions are large and 

its atmospheric abundance is far greater than the abundances 

of other emitted gases.

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 increase the extent 

and severity of winter polar ozone depletion. While the Earth’s 

surface is expected to continue warming in response to the 

net positive radiative forcing from greenhouse gas increases, 

the stratosphere is expected to continue cooling. A cooler 

stratosphere would extend the time period over which polar 

stratospheric clouds (PSCs) are present in winter and early 

spring and, as a result, might increase polar ozone depletion. 

In the upper stratosphere at altitudes above PSC formation 

regions, a cooler stratosphere is expected to increase ozone 

amounts because lower temperatures decrease the effectiveness of ozone loss reactions. Furthermore, climate change 

may alter the strength of the stratospheric circulation and with 

it the distribution of ozone in the stratosphere (see Q20).
Q.60

The success of the Montreal Protocol in controlling the production and consumption of ozone-depleting substances 

(ODSs) has protected the ozone layer (see Q15). The resulting 

reductions in atmospheric abundances of ODSs also reduced 

the human influence on climate because all ODSs are greenhouse gases (see Q18). By protecting both ozone and climate, 

the Montreal Protocol has provided a dual benefit to society 

and Earth’s ecosystems. In the following, the dual benefit of 

the Montreal Protocol is highlighted by considering long-term 

baseline and world-avoided scenarios of ODS emissions that 

use Ozone Depletion Potentials (ODPs), Global Warming 

Potentials (GWPs), equivalent effective stratospheric chlorine 

(EESC), and the radiative forcing of climate change.

Baseline ODS scenarios. The baseline scenarios of 

past and future ODS emissions presented here include the 

emissions of principal halogen source gases. They are constructed from (1) historical annual production and consumption of individual ODSs reported to the Montreal Protocol, 

(2) projected annual production and consumption of ODSs 

for future years based on provisions of the Protocol, (3) estimates of ODS banks, (4) atmospheric observations of ODSs 

and some naturally occurring halogen source gases, such as 

methyl chloride (CH3Cl), and (5) weighting factors related to 

ozone depletion and climate change.

In forming two of the baseline scenarios shown in Figure 

Q19-1 (upper panels), the emissions of each gas are added 

together after being weighted (multiplied) by the Ozone Depletion Potential (ODP) or the Global Warming Potential (GWP) 

of the respective gas (see Q18 and Table Q7-1). In the ODPweighted scenario, the emission sum is expressed as CFC-11-

equivalent emissions because CFC-11 is the reference gas with 

an assigned ODP value of 1. For example, in the sum, 1 kg of 

halon-1211 emissions is added as 7.9 kg of CFC-11-equivalent 

emissions because the ODP of halon-1211 is 7.9. Similarly, 

the GWP-weighted sum is expressed as CO2-equivalent emissions because CO2 is the reference gas with an assigned GWP 

of 1. For example, in the sum, 1 kg of carbon tetrachloride 

emissions is added as 1400 kg of CO2-equivalent emissions 

Yes. All ozone-depleting substances are also greenhouse gases that contribute to climate forcing when they accumulate 

in the atmosphere. Montreal Protocol controls have led to a substantial reduction in the emissions of ozone-depleting 

substances (ODSs) over the last two decades. These reductions have provided the added benefit of reducing the human 

contribution to climate change while protecting the ozone layer. Without Montreal Protocol controls, the climate forcing 

contribution from annual ODS emissions could now be 10-fold larger than its present value, which would be a significant 

fraction of the climate forcing from current carbon dioxide (CO2) emissions.

Have reductions of ozone-depleting substances under Q19 the Montreal Protocol also protected Earth’s climate?

Figure Q19-1. Montreal Protocol protection of ozone and climate. The provisions of the Montreal Protocol have 

substantially reduced ozone-depleting substances (ODSs) in the atmosphere. This has protected the ozone layer and also 

reduced the potential for climate change because ODSs are greenhouse gases. The scenarios and comparisons shown here 

demonstrate this dual benefit of the Montreal Protocol. Baseline scenarios for ODS emissions include all principal gases 

weighted by their Ozone Depletion Potentials (ODPs) or Global Warming Potentials (GWPs) (top panels). With these weightings, emissions are expressed as CFC-11-equivalent or CO2-equivalent mass per year. The lower panels show EESC and radiative forcing of climate as derived from the respective ODP- and GWP-weighted scenarios. The world-avoided emission scenarios assume ODS emission growth of 2 or 3% per year beyond 1987 abundances. Shown for reference are the emissions 

and radiative forcing of CO2, and the emissions reduction target of the first commitment period of the Kyoto Protocol. The 

contributions of natural halogen source gases are shown in the ODP-weighted and EESC scenarios (red dashed lines) and are 

negligible in the GWP-weighted and radiative forcing scenarios. The magnitude of the dual benefit has increased since about 

1987 as shown by differences between the world-avoided and baseline scenarios (blue shaded regions in each panel). For 

completeness, these differences can be adjusted by offsets due to additional ozone depletion and HFC emissions (see text). 

(A megatonne = 1 billion (109) kilograms. A gigatonne = 1 trillion (1012) kilograms.)

20 Questions: 2010 Update Section V: IMPLICATIONS OF OZONE DEPLETION AND THE MONTREAL PROTOCOL
Q.61

Section V: IMPLICATIONS OF OZONE DEPLETION AND THE MONTREAL PROTOCOL 20 Questions: 2010 Update

CO2 scenarios

Lower limits in 

CO2 scenarios

World-avoided scenarios

3% Annual growth from 1987

2% } without the Montreal Protocol

The Montreal Protocol Protection of Ozone and Climate 

From global emissions of all ozone-depleting substances (ODSs) and CO2

0

2

4

6

Emissions

(megatonnes CFC-11-equivalent per year)

0

0.5

1.0

1.5

2.0

Radiative forcing of climate

 Radiative forcing

(watts per square meter)

1960 1980 2000 2020

Year

1960 1980 2000 2020

Year

Emissions weighted by 

Ozone Depletion Potentials (ODPs)

Montreal

Protocol Ozone layer

protection

Emissions weighted by 

Global Warming Potentials (GWPs)

0

10

20

30

40

Emissions

(gigatonnes CO2-equivalent per year)

Montreal

Protocol

1960 1980 2000 2020

Year

Climate

protection

Montreal

Protocol

Climate

protection

Equivalent effective stratospheric

chlorine (EESC)

1960 1980 2000 2020

Year

 EESC

(relative amounts)

0

2

4

6

8

Montreal

Protocol

Natural halogen sources 

Baseline ODS scenarios

Ozone layer

protection

reduction target

for 2008–2012

 2.0 

Kyoto Protocol
Q.62

20 Questions: 2010 Update Section V: IMPLICATIONS OF OZONE DEPLETION AND THE MONTREAL PROTOCOL

because the GWP of carbon tetrachloride is 1400.

World-avoided ODS scenarios. The baseline scenario 

of ODS emissions can be contrasted with a scenario of ODS 

emissions that the world has avoided by agreeing to the Montreal Protocol (see Figure Q19-1). These world-avoided emissions are estimated by assuming that emissions of ODSs in the 

baseline scenario increase beyond 1987 values with a 2 or 3% 

annual growth rate. These growth rates are consistent with 

the strong market for ODSs in the late 1980s that included a 

wide variety of current and potential applications and that had 

potential for substantial new growth in developing countries.

CO2 emission scenarios. Long-term CO2 emission scenarios are also shown for comparison, as derived from past 

and projected CO2 emissions, because CO2 is the principal 

greenhouse gas related to human activities. The projected 

CO2 emissions have high and low scenarios that are derived 

using different basic assumptions about future economies, 

technical progress, and societal decisions.

ODP-weighted emissions scenarios. The ODPweighted emissions in the ODS baseline scenario are a measure of the overall threat to stratospheric ozone from ODSs 

(see Figure Q19-1, upper left panel). When ODP-weighted 

emissions increase (decrease) in a given year, more (less) 

ozone will be destroyed in future years. ODP-weighted emissions increased substantially in the baseline scenario between 

1960 and 1987, the year the Montreal Protocol was signed (see 

Figure Q19-1 and Q0-1). After 1987, ODP-weighted emissions began a long and steady decline to present-day values. 

The decline in emissions is expected to continue, causing the 

atmospheric abundances of individual ODSs to decrease (see 

Figure Q16-1). The reduction in ODP-weighted emissions 

from the 1987 value is a conservative measure of the annual 

emissions avoided by the Montreal Protocol since 1987 and, 

hence, of the success of the Montreal Protocol in protecting 

the ozone layer.

Annual ODP-weighted emissions in the world-avoided scenario are about double the 1987 values by 2020. The annual 

differences between the world-avoided emissions and the 

baseline scenario (blue shaded region in Figure Q19-1) provide reasonable upper-limits to the ODP-weighted emissions 

avoided by the Montreal Protocol each year since 1987.

GWP-weighted emissions scenarios. The GWPweighted emissions in the ODS baseline scenario are a measure of the overall threat to climate from ODSs (see Figure 

Q19-1, upper right panel). As ODS emissions accumulate in 

the atmosphere, their climate forcing contribution increases. 

The long-term changes in the GWP-weighted scenario are 

very similar to those in the ODP-weighted scenario. Both 

show an increase before 1987 and decrease afterwards. The 

similarity follows from the predominant role that CFC-11 and 

CFC-12 emissions play in ozone depletion and climate forcing from ODSs. The reduction in GWP-weighted emissions 

since 1987 is a conservative measure of the substantial success 

of the Montreal Protocol in reducing the potential for climate 

change from human activities. The annual differences since 

1987 between the world-avoided emissions and the baseline 

scenario (blue shaded region in Figure Q19-1) provide reasonable upper-limits to the GWP-weighted emissions avoided by 

the Montreal Protocol each year since 1987.

The climate protection calculated using differences 

between world-avoided emissions and the baseline scenario 

has two offsetting effects. The first is the additional ozone 

depletion that would be caused by world-avoided ODS emissions. Ozone depletion offsets ODS climate forcing because 

a greenhouse gas (ozone) is being removed from the atmosphere in response to ODS emissions (see Q18). The second 

effect is the increase in emissions of HFC substitute gases that 

occurred in response to ODS reductions from Montreal Protocol controls. More HFCs in the atmosphere offset the gain 

in climate protection from ODS reductions because HFCs are 

also greenhouse gases (see Q18).

The combined magnitude of these offsets in 2010, for example, 

is about 30% of the difference between the baseline and worldavoided scenarios. The resulting net GWP-weighted emission 

reduction in 2010 is about 9.7–12.5 gigatonnes CO2-equivalent 

per year. In contrast, the annual emissions reduction target 

adopted by the Kyoto Protocol during its first commitment 

period (2008–2012) is estimated as 2 gigatonnes CO2-equivalent per year (see Figure Q19-1). The reductions are expected 

to result from controlling the Kyoto Protocol basket of gases 

that includes HFCs and does not include ODSs (see Q18). As 

a result, the upper limit for the net reduction in annual GWPweighted emissions achieved by the Montreal Protocol in 2010 

is 5- to 6-fold larger than the Kyoto Protocol target.

Annual GWP-weighted emissions of ODSs were a large 

percentage (about 20–40%) of CO2-baseline emissions between 

1960 and 1989 (see Figure Q19-1). Thereafter, this percentage 

has steadily decreased and is projected to reach 2–3% by 2020. 

This projection stands in sharp contrast to the world-avoided 

scenario, in which the percentage increases to 40–75% of 

CO2-baseline emissions by 2020.

EESC scenarios. The EESC scenario in Figure Q19-1 

(lower left panel) provides a measure of the year-to-year 

potential of the atmospheric abundances of ODSs to destroy 
Q.63

Section V: IMPLICATIONS OF OZONE DEPLETION AND THE MONTREAL PROTOCOL 20 Questions: 2010 Update

stratospheric ozone. Changes in historical and projected 

atmospheric emissions of ODSs cause changes in their atmospheric abundances. The derivation of EESC from ODS 

atmospheric abundances is discussed in Q16 and similar 

EESC baseline scenarios are shown in Figures Q14-1, Q15-1, 

and Q16-1 for different time intervals. An increase in ODPweighted emissions always leads to some increase in EESC 

in the years following the emissions. When ODS-weighted 

emissions decreased after 1987, EESC did not proportionally decrease because of the long atmospheric lifetimes of the 

principal ODSs. In Figure Q19-1, for example, EESC reached 

its peak nearly a decade after the peak in ODP-weighted emissions, and by 2010 the decrease in EESC from its peak value 

was only about 10%, compared to the 70% decrease in ODPweighted emissions achieved by 2010.

Radiative forcing of climate change scenarios. The 

radiative forcing derived for the ODS baseline scenario in 

Figure Q19-1 (lower right panel) provides a measure of the 

year-to-year contribution to climate forcing from atmospheric 

ODS abundances. The radiative forcing of an ODS is proportional to its radiative efficiency and the net increase in its 

atmospheric abundance during the Industrial Era. Increases 

in abundance up to the present are derived from atmospheric 

observations. Future abundances rely on projected emissions and atmospheric lifetimes of each gas. In Figure Q19-1, 

radiative forcing due to ODSs increases smoothly from 1960 

onward, peaks in 2003 and decreases very gradually in subsequent years. Radiative forcing responds to ODS emission 

reductions in a manner similar to EESC, with the current slow 

decline attributable to the two principal contributing gases, 

CFC-11 and CFC-12, and their long atmospheric lifetimes 

(45–100 years).

The differences in ODS climate forcing between the worldavoided and baseline scenarios are offset by additional ozone 

depletion and HFC emissions in a manner similar to that 

noted above for differences in GWP-weighted emissions. 

After accounting for these two offsets, the climate forcing due 

to ODSs in the world-avoided scenario is approximately 70% 

higher than that in the baseline scenario in 2010 and approximately 30% of that due to CO2. 

The considerable contributions that ODSs could have made 

to climate forcing, if not controlled by the Montreal Protocol, attests to their potency as greenhouse gases. ODSs had 

negligible atmospheric abundances 50–60 years ago and, as 

a group, represent chlorine amounts that currently are about 

100,000 times less abundant in the atmosphere than CO2.

The future. Fewer control options are available to 

increase the dual benefit of the Montreal Protocol beyond 

2020 because the most effective and abundant ODSs have 

already been phased out under Montreal Protocol provisions 

(see Q16). The most recent Montreal Protocol action increased 

ozone and climate protection by accelerating the phase-out 

of HCFCs (Montreal in 2007) (see Q15). This provision is 

expected to reduce total GWP-weighted emissions of HCFCs 

by about 50% between 2010 and 2050, corresponding to about 

18 gigatonnes of CO2-equivalent emissions. The accelerated 

HCFC phase-out protects ozone by advancing the date that 

EESC returns to 1980 values by 4–5 years. 

Ozone and climate protection could be further enhanced 

with Montreal Protocol provisions that increase the effectiveness of capturing and destroying ODSs contained in banks, 

namely, those ODSs currently being used in refrigeration, air 

conditioning, and fire protection equipment, or stockpiled for 

servicing long-term applications.

Future projections suggest that growth in HFC production 

and consumption could result in GWP-weighted emissions 

of up to 8.8 gigatonnes CO2-equivalent per year by 2050, primarily in developing nations. The 2050 value is comparable 

to the peak in GWP-weighted ODS emissions in 1988 (see 

Figure Q19-1). The estimate assumes that HFC application 

demand would be met using the same suite of HFCs currently used in developed countries. If future HFC demand is 

met instead with lower-GWP substances, the 2050 estimate 

would be substantially reduced. International proposals 

have been put forth for the Montreal Protocol to expand its 

production and consumption controls to include HFCs. The 

expansion would occur in collaboration with the Kyoto Protocol, which currently includes HFCs in its basket of gases. 

If the proposals are successful, the Montreal Protocol would 

have the opportunity to guide the transition from ODSs to 

HFCs in a manner that would optimize the protection of the 

ozone layer and climate while minimizing the burden on 

participating nations. 
Q.64

Substantial recovery from the depletion of global and polar ozone caused by ODSs is expected in the later decades 

of this century. The recovery follows on the success of the 

Montreal Protocol in reducing the global production and 

consumption of ODSs. Now, the atmospheric abundances 

of most major ODSs and the associated values of equivalent 

effective stratospheric chlorine (EESC) are in decline (see 

Q16). In contrast to the diminishing role of ODSs, changes 

in climate are expected to have an increasing influence on 

future ozone. Important aspects are the projected growth 

of greenhouse gas abundances and the resulting changes in 

stratospheric temperatures and the stratospheric circulation. 

Chemistry-climate models can be used to project how ozone is 

expected to respond to changes in ODSs and climate in different geographical regions during the recovery period. Global 

events, such as major volcanic eruptions and geoengineering 

actions, may also influence future total ozone amounts.

Using chemistry-climate models. Projections of total 

ozone presented here are based on the results from a group 

of chemistry-climate models that take into account the influences of changes in ODSs and climate. These models show 

how changes in ozone are expected to vary across global 

regions by evaluating the complex interaction of the processes that control ozone and climate involving radiation, 

chemistry, and transport. Required model inputs include, 

for example, historical and projected emissions of ODSs, 

carbon dioxide (CO2), methane (CH4), and nitrous oxide 

(N2O), and changes in sea surface temperatures. The results 

from chemistry-climate model simulations are used to identify aspects of the models that are particularly important for 

future ozone abundances. For example, model projections for 

the coming decades show a strengthening in the atmospheric 

circulation that brings air to the stratosphere in the tropics, 

moves air poleward into both hemispheres, and then returns 

it to the troposphere at middle to high latitudes. These circulation changes will significantly alter the global distribution 

of ozone and the atmospheric lifetimes of ODSs and other 

long-lived gases.

Simulating recent ozone changes. Comparisons of 

model results with observations help confirm the causes of 

ozone depletion and increase confidence in model projections 

of future ozone amounts. Two important measures of ozone, 

global total ozone and minimum total ozone values in the 

Antarctic, are compared to the group of chemistry-climate 

model simulations in Figure Q20-1. Both ozone measures 

show substantial depletion since 1980. The average model 

value follows the observed decline in both ozone measures, 

indicating that the main processes involved in ozone depletion are reasonably well represented by the models. Some 

differences between the simulations and observations can 

be explained by unusual meteorological conditions, volcanic eruptions, changes in solar activity, or other natural 

influences, which are not all fully accounted for in the various models used here. In Figure Q20-1, global total ozone 

has increased in the last 20 years, while Antarctic minimum 

ozone values have been relatively constant. The global total 

ozone increases cannot be attributed solely to reductions in 

ODSs that began in the 1990s, because total ozone has also 

recovered from additional depletion caused by the 1991 Mt. 

Pinatubo eruption (see Q14).

Equivalent stratospheric chlorine (ESC) projections. Equivalent stratospheric chlorine (ESC) values are 

also projected in chemistry-climate models to represent how 

the potential for reactive halogen gases (see Q8) to destroy 

ozone varies over time. ESC is similar to equivalent effective 

stratospheric chlorine (EESC) (see Q16) in that both represent a weighted sum of reactive chlorine and bromine gases 

in the stratosphere. The weighting in each accounts for the 

greater effectiveness of bromine atoms in destroying ozone. 

ESC can be calculated more precisely and comprehensively 

Substantial recovery of the ozone layer from the effects of ozone-depleting substances (ODSs) is expected near the middle 

of the 21st century, assuming global compliance with the Montreal Protocol. Recovery will occur as ODSs and reactive halogen gases in the stratosphere decrease in the coming decades. In addition to responding to ODSs, future ozone amounts 

will increasingly be influenced by expected changes in climate. The resulting changes in stratospheric ozone will depend 

strongly on the geographic region. During the long recovery period, large volcanic eruptions could temporarily reduce 

global ozone amounts for several years.

Q20 How is ozone expected to change in the coming decades?

Section VI: STRATOSPHERIC OZONE IN THE FUTURE
Q.65

than EESC because the chemistry-climate models can derive 

ESC as a function of altitude, latitude, longitude, and time. 

The ESC calculations are based on the history and projections 

of ODS surface abundances and the chemical and transport 

processes that control (1) the conversion of ODSs to reactive 

halogen gases, (2) the distribution of reactive halogen gases in 

the global stratosphere and (3) their ultimate removal from 

the stratosphere. 

The long-term changes in ESC in the global and regional 

analyses shown in Figure Q20-2 are highly similar. In all 

regions, changes in ESC from 1960 values increase smoothly 

with time, reach a peak near the end of the 20th century, and 

decrease gradually until the end of the 21st century. Values at 

the end of the 21st century are approaching those in 1960, indicating that ODSs are largely removed from the stratosphere by 

that time. ESC returns to 1980 values several decades sooner 

than to 1960 values because of the slow rate of decline.

Peak ESC values around the year 2000 are highest in polar 

regions and lowest in the tropics. In the tropics, stratospheric 

air has only recently been transported from the troposphere, 

with the result that only a small fraction of ODSs has undergone conversion to reactive halogen gases (see Q8). In polar 

regions, the fraction is much larger because stratospheric air 

requires several years on average to journey from its entry 

point in the tropics to the polar lower stratosphere. During 

this time, a much larger fraction of ODSs undergoes conversion to reactive halogen gases.

Long-term total ozone projections. Total ozone 

changes derived from chemistry-climate models, as referenced 

to 1960 values, are shown in Figure Q20-2. The range of values 

from the group of models is included in the figure as one measure of the uncertainty in the model projections. Total ozone 

changes in the different regions are described as follows:

▶ Antarctic. Total ozone changes are largest in the Antarctic region in springtime (October). Chemistry-climate 

models show that ODSs are the predominant factor in 

Antarctic ozone depletion in the past and in the coming 

decades. Changes in climate parameters have a smaller 

role. As a result, total ozone changes closely mirror

changes in ESC: as ESC increases, ozone proportionately decreases; as ESC decreases, ozone proportionately 

increases. Antarctic total ozone is projected to return to 

1980 levels after midcentury and later than in any other 

region. Meteorological variability in polar regions in late 

Section VI: STRATOSPHERIC OZONE IN THE FUTURE 20 Questions: 2010 Update

Figure Q20-1. Simulations of ozone depletion.

Chemistry-climate models of the atmosphere that 

account for changes in ozone-depleting substances 

(ODSs) and climate parameters are widely used to 

assess past ozone changes and project future ozone. 

Agreement in comparisons of model results with 

observations increases confidence in the model projections and our understanding of the processes leading 

to ozone depletion. Observed values of global total 

ozone (top panel) and minimum total ozone values 

over Antarctica (bottom panel) decreased beginning 

in the early 1980s (red points). Ozone is no longer 

decreasing in either region. The observations lie mostly 

within the range of projections derived from a group of 

chemistry-climate models (blue shading). Differences 

between models and observations can occur due to 

unusual meteorological conditions, volcanic eruptions, 

changes in solar activity, or other influences, which are 

not fully accounted for in the different models. As ODS 

abundances decrease in the 21st century, chemistryclimate models project global total ozone to increase 

steadily and exceed 1980 values, and Antarctic minimum ozone to return to 1980 values.

Antarctic minimum total ozone

Average and range

of atmospheric

model predictions

Annually averaged

observations

Average and range

of atmospheric

model predictions

Averages from

observations

Global total ozone change from 1980

-20

-25

-10

-15

-5

0

5

10

200

100

0

250

150

50

Total ozone change (DU) Minimum total ozone (DU)

Year

1980 2000 2020 2040

Simulations of Stratospheric Ozone Depletion

Results from chemistry-climate models
Q.66

winter/early spring when ozone depletion occurs causes a 

large range in the model projections.

▶ Arctic. Total ozone changes in the Arctic region in 

springtime (March) are considerably smaller than in the 

Antarctic. In contrast to the Antarctic, ozone changes 

do not closely mirror changes in ESC. After midcentury, 

Arctic total ozone increases to values above those that 

would be expected from ESC reductions alone because of 

the strengthening of the atmospheric circulation and the 

enhanced stratospheric cooling associated with increases 

in CO2. By 2100 Arctic total ozone is projected to be well 

above both 1960 and 1980 values. The large range in projections compared to other nonpolar regions is due to greater 

meteorological variability as noted for the Antarctic. Arctic 

total ozone is projected to return to 1980 levels between 

2020 and 2035, which is two to three decades before ESC 

returns to 1980 levels, and to continue to increase until the 

end of the century.

▶ Northern and southern midlatitudes. The annual 

averages of total ozone changes in midlatitudes are much 

20 Questions: 2010 Update Section VI: STRATOSPHERIC OZONE IN THE FUTURE

Northern midlatitudes

40

0

-40

4

0

1980 2020 2060 2100

Southern midlatitudes

Tropics

40

0

-40

4

0

1980 2020 2060 2100

40

0

-40

4

0

1980 2020 2060 2100

Year

Global

40

0

-40

4

0

1980 2020 2060 2100

Year

Total Ozone and Equivalent Stratospheric Chlorine

Results from atmospheric chemistry-climate models for 1960 to 2100

(2090 – 2099 average)

200 300 400 450 

Total ozone (DU)

Antarctic (October)

4

0

1980 2020 2060 2100

Arctic (March)

40

0

-40

-80

80

4

0

8

1980 2020 2060 2100

Year

40

0

-40

-80

-160

-120

Ozone 

ESC

Ozone 

ESC

Ozone 

ESC

Ozone 

ESC

Ozone 

ESC

Ozone 

ESC

Total ozone (DU)

Equivalent stratospheric

chlorine (ESC) (ppb) near 

21-km altitude

Annual (or monthly) mean

changes from 1960 values

Ranges of model values

1980 values
Q.67

Section VI: STRATOSPHERIC OZONE IN THE FUTURE 20 Questions: 2010 Update

Figure Q20-2. Long-term changes in ozone and equivalent stratospheric chlorine (ESC). Chemistry-climate models 

are used to make projections of total ozone amounts that account for the effects of ozone-depleting substances (ODSs) and 

climate change. Regional and global projections are shown for total ozone and ESC for the period 1960–2100, referenced 

to 1960 values. The globe in the center shows average total ozone projections for the last decade of the 21st century. 

Total ozone depletion increased after 1960 as ESC values steadily increased throughout the stratosphere. ESC values have 

peaked and are now in a slow decline. All the projections show maximum total ozone depletion around 2000, coincident 

with the highest abundances of ESC. Thereafter, total ozone increases, except in the tropics, as ESC slowly declines. In all 

the projections except the Antarctic and the tropics, total ozone returns to 1960 values by midcentury, which is earlier than 

expected from the decrease in ESC alone. The earlier returns are attributable to climate change, which influences total 

ozone through changes in stratospheric transport and temperatures. In the tropics, in contrast, climate change causes total 

ozone to remain below 1960 values throughout the century. In the Antarctic, the effect of climate change is smaller than 

in other regions. As a result, Antarctic total ozone in springtime mirrors the changes in ESC, with both closely approaching 

1960 values at century’s end. The dots on each curve mark the occurrences of 1980 values of total ozone and ESC. Note that 

the equal vertical scales in each panel allow direct comparisons of ozone and ESC changes between regions.

smaller than the springtime losses in polar regions. Both 

midlatitude regions resemble the Arctic with total ozone 

returning to 1960 and 1980 values much sooner than ESC. 

In the northern midlatitudes, the models predict a return 

to 1960 values in 2030 whereas ESC requires the full century to return close to 1960 values. Total ozone in southern 

midlatitudes changes in a manner similar to the northern 

midlatitudes. Differences are that in southern midlatitudes the 1960 return date for total ozone is somewhat later 

(2055) and the maximum ozone depletion observed near 

2000 is greater. Both aspects reflect the influence of Antarctic stratospheric air, depleted in ozone, being routinely 

transported to southern midlatitudes in spring (see Q11). 

The more rapid return of total ozone in both regions compared with ESC derives again from the influence of climateinduced changes in transport and upper stratospheric temperatures in the model projections. The model range is 

significantly less in midlatitude annual averages than in the 

polar springtime averages. After returning to either 1960 or 

1980 values, total ozone continues to increase in both hemispheres and significantly exceeds those values by 2100.

▶ Tropics. Total ozone changes in the tropics are smaller 

than in any other region. Ozone is less sensitive to ESC in 

the tropical stratosphere because of the dominant roles of 

production and transport in controlling ozone. In contrast 

to other regions, chemistry-climate models project total 

ozone to remain below 1960 values throughout the 21st century. Total ozone gradually returns to 1980 values in 2040, 

peaks around 2060 and decreases again until the end of the 

century. The different total ozone behavior in the tropics 

is due to the changing balance between ozone increases in 

the upper stratosphere and decreases in the lower stratosphere. The ozone increases in the upper stratosphere 

occur from declining ODS abundances and decreases in 

temperatures caused by increasing greenhouse gases, primarily CO2. In the lower stratosphere, ozone is reduced 

because the strengthening of the stratospheric circulation 

reduces the time for ozone production before air leaves the 

tropics. These circulation changes also influence the Arctic 

and midlatitude regions as noted above.

▶ The globe. The annual averages of global total ozone 

are projected to return to 1960 levels around the middle 

of the century (2040 to 2080) while ESC returns to 1960 

values near century’s end. The comparable return dates for 

1980 values are substantially earlier. Chemistry-climate 

model analysis suggests that this early return of total ozone 

is primarily a result of upper stratosphere cooling and a 

strengthened circulation as noted for midlatitude ozone. 

The larger total ozone changes shown for polar regions in 

springtime do not significantly influence global total ozone 

values because these values are annual averages and polar 

regions are a small geographical fraction of the globe.

Future ultraviolet radiation. Projections of long-term 

changes in total ozone can be used to estimate long-term 

changes in solar ultraviolet radiation reaching Earth’s surface (see Q17). The UV-B component of ultraviolet radiation 

increases as total ozone decreases. Based on the global total 

ozone projections, clear-sky UV-B radiation is expected to 

decrease below 1960 values by the end of the century because 

ozone is above 1960 values in the later decades. The latitude 
Q.68

20 Questions: 2010 Update Section VI: STRATOSPHERIC OZONE IN THE FUTURE

regions likely to have some increase in ultraviolet radiation 

are the Antarctic and the tropics, where total ozone remains 

lower than 1960 values until the end of the century. Large 

increases or decreases in surface erythemal radiation away 

from 1960 values are expected to lead to adverse effects on 

human and ecosystem health. 

Volcanoes and geoengineering. Other factors not 

included in chemistry-climate models can potentially affect 

future total ozone amounts. Explosive volcanic eruptions 

have temporarily reduced global total ozone in the past (see 

Q14). Similar eruptions, especially in the early decades of 

this century when ESC values are highest, are also expected 

to reduce total ozone for a few years. Volcanic eruptions are 

an additional source of uncertainty not included in the ozone 

projections in Figure Q20-2.

Several geoengineering methods have been proposed to 

reduce climate forcing from human activities. A widely 

discussed proposal is the enhancement of sulfate aerosols 

in the stratosphere from direct injections of sulfur or sulfuric acid. The expected response is a cooling of the climate 

system from increased aerosol scattering of sunlight, similar 

to that observed after some explosive volcanic eruptions. The 

required injections, if sustained over many years, are likely 

to have unintended consequences such as reductions in total 

ozone amounts and changes in stratospheric temperatures 

and circulation.
Appendices

Appendix A

List of International Authors, Contributors, and Reviewers

Appendix B

Major Acronyms and Abbreviations

Appendix C

Major Chemical Formulae and Nomenclature

from this Assessment

A.1

Appendix A

List of International Authors,

Contributors, and Reviewers

Cochairs

Ayité-Lô Nohende Ajavon Université de Lomé Togo

Paul A. Newman NASA Goddard Space Flight Center USA

John A. Pyle University of Cambridge, National Centre for Atmospheric Science UK

A.R. Ravishankara NOAA ESRL Chemical Sciences Division USA

Authors and Contributors

CHAPTER 1

OZONE-DEPLETING SUBSTANCES (ODSs) AND RELATED CHEMICALS

Coordinating Lead Authors

Stephen A. Montzka NOAA ESRL Global Monitoring Division USA

Stefan Reimann Swiss Federal Laboratories for Materials Science and Technology (Empa) Germany

Lead Authors

Andreas Engel Goethe University, Frankfurt am Main Germany

Kirstin Krüger IFM-GEOMAR / Leibniz-Institute of Marine Sciences at Kiel University Germany

Simon O’Doherty University of Bristol UK

William T. Sturges University of East Anglia UK

Coauthors

Donald R. Blake University of California, Irvine USA

Marcel Dorf University of Heidelberg Germany

Paul J. Fraser CSIRO Division of Marine and Atmospheric Research Australia

Lucien Froidevaux Jet Propulsion Laboratory, California Institute of Technology USA

Kenneth Jucks NASA Headquarters USA

Karin Kreher National Institute of Water and Atmospheric Research (NIWA) New Zealand

Michael J. Kurylo Goddard Earth Sciences and Technology Center / USA

 University of Maryland Baltimore County

Abdelwahid Mellouki Institut de Combustion Aérothermique Réactivité et France

 Environnement (ICARE), CNRS

John Miller CIRES-University of Colorado / NOAA ESRL Global Monitoring Division USA

Ole John Nielsen University of Copenhagen Denmark

Vladimir L. Orkin National Institute of Standards and Technology USA

Ronald G. Prinn Massachusetts Institute of Technology, Center for Global Change Science USA

Robert Rhew University of California, Berkeley USA

Michelle L. Santee Jet Propulsion Laboratory, California Institute of Technology USA

Andreas Stohl Norwegian Institute for Air Research (NILU) Norway

Daniel P. Verdonik Hughes Associates, Inc. USA
A.2

Appendix A

Contributors

Elliot Atlas University of Miami USA

Peter F. Bernath University of York UK

Thomas Blumenstock Karlsruhe Institute of Technology Germany

James H. Butler NOAA ESRL Global Monitoring Division USA

André Butz Netherlands Institute for Space Research (SRON) The Netherlands

Brian J. Connor BC Consulting Limited New Zealand

Pierre Duchatelet University of Liège Belgium

Geoffrey S. Dutton CIRES-University of Colorado / NOAA ESRL Global Monitoring Division USA

François Hendrick Belgian Institute for Space Aeronomy Belgium

Paul B. Krummel Centre for Australian Weather & Climate Research / Australia

 CSIRO Marine & Atmospheric Research

Lambert Kuijpers Technical University, Eindhoven Center for Sustainability The Netherlands

Emmanuel Mahieu University of Liège Belgium

Alistair J. Manning Met Office UK

Jens Mühle Scripps Institution of Oceanography, University of California, San Diego USA

Klaus Pfeilsticker University of Heidelberg Germany

Birgit Quack IFM-GEOMAR / Leibniz-Institute of Marine Sciences at Kiel University Germany

Martin N. Ross The Aerospace Corporation USA

Ross J. Salawitch University of Maryland USA

Sue Schauffler National Center for Atmospheric Research, Atmospheric Chemistry Division USA

Isobel J. Simpson University of California, Irvine USA

Darin W. Toohey University of Colorado USA

Martin K. Vollmer Swiss Federal Laboratories for Materials Science and Technology (Empa) Switzerland

Timothy J. Wallington Ford Motor Company USA

Hsiang J. (Ray) Wang Georgia Institute of Technology USA

Ray F. Weiss Scripps Institution of Oceanography, University of California, San Diego USA

Masaaki Yamabe National Institute of Advanced Industrial Science and Technology; TEAP Japan

Yoko Yokouchi National Institute for Environmental Studies Japan

Shari A. Yvon-Lewis Texas A&M University USA

CHAPTER 2

STRATOSPHERIC OZONE AND SUFACE ULTRAVIOLET RADIATION

Coordinating Lead Authors

Anne Douglass NASA Goddard Space Flight Center USA

Vitali Fioletov Environment Canada Canada

Lead Authors

Sophie Godin-Beekmann Laboratoire Atmosphères, Milieux, Observations Spatiales France

 ( LATMOS)/UPMC, CNRS

Rolf Müller Forschungszentrum Jülich Germany

Richard S. Stolarski NASA Goddard Space Flight Center USA

Ann R. Webb University of Manchester, School of Earth, Atmospheric, and UK

 Environmental Sciences

Coauthors

Antti Arola Finnish Meteorological Institute Finland

James B. Burkholder NOAA ESRL Chemical Sciences Division USA

John P. Burrows University of Bremen, Institute of Environmental Physics Germany/UK

Martyn P. Chipperfield University of Leeds UK

Raul Cordero Universidad de Santiago de Chile Chile

Christine David Laboratoire Atmosphères, Milieux, Observations Spatiales France

 (LATMOS)/UPMC, CNRS
A.3

Authors, Contributors, and Reviewers

Peter N. den Outer National Institute for Public Health and the Environment The Netherlands

Susana B. Diaz INGEBI / CADIC / Consejo Nacional de Investigaciones Cientificas Argentina

Lawrence E. Flynn NOAA National Environmental Satellite, Data, and Information Service USA

Michaela I. Hegglin University of Toronto Canada

Jay R. Herman University of Maryland-JCET / Goddard Space Flight Center USA

Petra Huck National Institute of Water and Atmospheric Research (NIWA) New Zealand

Serm Janjai Silpakorn University Thailand

Imre M. Jánosi Eötvös Loránd University, Budapest Hungary

Janusz W. Krzyścin Polish Academy of Sciences Poland

Yi Liu Chinese Academy of Sciences, Institute of Atmospheric Physics China

Jennifer Logan Harvard University USA

Katja Matthes Helmholtz Centre Potsdam (GFZ) / Freie Universität Berlin Germany

Richard L. McKenzie National Institute of Water and Atmospheric Research (NIWA) New Zealand

Nzioka John Muthama University of Nairobi Kenya

Irina Petropavlovskikh CIRES-University of Colorado / NOAA ESRL Global Monitoring Division USA

Michael Pitts NASA Langley Research Center USA

S. Ramachandran Physical Research Laboratory India

Markus Rex Alfred Wegener Institute for Polar and Marine Research, Potsdam Germany

Ross J. Salawitch University of Maryland USA

Björn-Martin Sinnhuber University of Bremen Germany

Johannes Staehelin Swiss Federal Institute of Technology – Zurich Switzerland

Susan Strahan NASA Goddard Space Flight Center USA

Kleareti Tourpali Aristotle University of Thessoloniki Greece

Jéssica Valverde-Canossa Universidad Nacional Costa Rica

Corinne Vigouroux Belgian Institute for Space Aeronomy Belgium

Contributors

Gregory E. Bodeker Bodeker Scientific New Zealand

Timothy Canty University of Maryland USA

Hugo De Backer Royal Meteorological Institute of Belgium Belgium

Philippe Demoulin University of Liège Belgium

Uwe Feister Deutscher Wetterdienst Germany

Stacey M. Frith Science Systems and Applications, Inc. USA

Jens-Uwe Grooß Forschungszentrum Jülich Germany

Frank Hase Karlsruhe Institute of Technology Germany

Jon Klyft Chalmers University of Technology Sweden

Takashi Koide Japan Meteorological Agency Japan

Michael J. Kurylo Goddard Earth Sciences and Technology Center / USA

 University of Maryland Baltimore County

Diego Loyola Deutsches Zentrum für Luft- und Raumfahrt, Germany

 Institut für Methodik der Fernerkundung

Chris A. McLinden Environment Canada Canada

Inna A. Megretskaia Harvard University USA

Prijitha J. Nair Laboratoire Atmosphères, Milieux, Observations Spatiales France

 (LATMOS)/UPMC, CNRS

Mathias Palm University of Bremen Germany

Dimitrios Papanastasiou CIRES-University of Colorado / NOAA ESRL Chemical Sciences Division USA

Lamont R. Poole Science Systems and Applications, Inc. USA

Matthias Schneider Karlsruhe Institute of Technology Germany

Robyn Schofield Alfred Wegener Institute for Polar and Marine Research, Potsdam Germany

Harry Slaper National Institute for Public Health and the Environment The Netherlands

Wolfgang Steinbrecht Deutscher Wetterdienst Germany

Susann Tegtmeier IFM-GEOMAR / Leibniz-Institute for Marine Sciences at Kiel University Germany

Yukio Terao National Institute for Environmental Studies Japan
A.4

Appendix A

Simone Tilmes National Center for Atmospheric Research USA

Dmitry I. Vyushin University of Toronto Canada

Mark Weber University of Bremen, Institute of Environmental Physics Germany

Eun-Su Yang University of Alabama in Huntsville USA

CHAPTER 3

FUTURE OZONE AND ITS IMPACT ON SURFACE UV

Coordinating Lead Authors

Slimane Bekki Laboratoire Atmosphères, Milieux, Observations Spatiales France

 (LATMOS)/IPSL, CNRS

Gregory E. Bodeker Bodeker Scientific New Zealand

Lead Authors

Alkiviadis F. Bais Aristotle University of Thessaloniki Greece

Neal Butchart Met Office Hadley Centre UK

Veronika Eyring Deutsches Zentrum für Luft- und Raumfahrt, Germany

 Institut für Physik der Atmosphäre

David W. Fahey NOAA ESRL Chemical Sciences Division USA

Douglas E. Kinnison National Center for Atmospheric Research, Atmospheric Chemistry Division USA

Ulrike Langematz Freie Universität Berlin Germany

Bernhard Mayer Ludwig-Maximilians-University Munich Germany

Robert W. Portmann NOAA ESRL Chemical Sciences Division USA

Eugene Rozanov World Radiation Center/Institute for Atmospheric and Switzerland

 Climate Science-Zürich (ETHZ)

Coauthors

Peter Braesicke University of Cambridge, National Centre for Atmospheric Science UK

Andrew J. Charlton-Perez University of Reading UK

Natalia E. Chubarova Moscow State University Russia

Irene Cionni Deutsches Zentrum für Luft- und Raumfahrt, Germany

 Institut für Physik der Atmosphäre

Susana B. Diaz INGEBI / CADIC / Consejo Nacional de Investigaciones Cientificas Argentina

Nathan P. Gillett Environment Canada, Canadian Centre for Climate Modelling and Analysis Canada

Marco A. Giorgetta Max-Planck-Institut für Meteorologie – Hamburg Germany

Ninong Komala National Institute of Aeronautics and Space Indonesia

Franck Lefèvre Laboratoire Atmosphères, Milieux, Observations Spatiales France

 (LATMOS)/IPSL, CNRS

Charles McLandress University of Toronto Canada

Judith Perlwitz CIRES-University of Colorado / NOAA ESRL Physical Sciences Division USA

Thomas Peter Swiss Federal Institute of Technology, Switzerland

 Institute for Atmospheric and Climate Science

Kiyotaka Shibata Meteorological Research Institute Japan

Contributors

Hideharu Akiyoshi National Institute for Environmental Studies Japan

John Austin University Corporation for Atmospheric Research / USA

 NOAA Geophysical Fluid Dynamics Laboratory

Martyn P. Chipperfield University of Leeds UK

Martin Dameris Deutsches Zentrum für Luft- und Raumfahrt, Germany

 Institut für Physik der Atmosphäre

Sandip Dhomse University of Leeds UK

Stacey M. Frith Science Systems and Applications, Inc. USA

Rolando R. Garcia National Center for Atmospheric Research, Atmospheric Chemistry Division USA
A.5

Authors, Contributors, and Reviewers

Hella Garny Deutsches Zentrum für Luft- und Raumfahrt, Germany

 Institut für Physik der Atmosphäre Germany

Andrew Gettelman National Center for Atmospheric Research, Atmospheric Chemistry Division USA

Steven C. Hardiman Met Office Hadley Centre UK

Patrick Jöckel Max Planck Institute for Chemistry, Mainz; now at DLR-IPA Germany

Andreas I. Jonsson University of Toronto Canada

Andreas Kazantzidis University of Patras Greece

Anne Kubin Freie Universität Berlin Germany

Jean-François Lamarque National Center for Atmospheric Research, Atmospheric Chemistry Division USA

Eva Mancini Università degli Studi di L’Aquila, Dipartimento di Fisica Italy

Marion Marchand Laboratoire Atmosphères, Milieux, Observations Spatiales France

 (LATMOS)/IPSL, CNRS

Martine Michou GAME-CNRM/Météo-France, CNRS France

Olaf Morgenstern National Institute of Water and Atmospheric Research (NIWA) New Zealand

Luke D. Oman NASA Goddard Space Flight Center USA

Steven Pawson NASA Goddard Space Flight Center USA

Giovanni Pitari Università degli Studi di L’Aquila, Dipartimento di Fisica Italy

David Plummer Environment Canada, Canadian Centre for Climate Modelling and Analysis Canada

John A. Pyle University of Cambridge, National Centre for Atmospheric Science UK

David Saint-Martin GAME-CNRM/Météo-France, CNRS France

John F. Scinocca Environment Canada, Canadian Centre for Climate Modelling and Analysis Canada

Theodore G. Shepherd University of Toronto Canada

Dan Smale National Institute of Water and Atmospheric Research (NIWA) New Zealand

Richard S. Stolarski NASA Goddard Space Flight Center USA

Hubert Teyssèdre GAME-CNRM/Météo-France, CNRS France

Simone Tilmes National Center for Atmospheric Research, USA

 Climate and Global Dynamics Division

CHAPTER 4

STRATOSPHERIC CHANGES AND CLIMATE

Coordinating Lead Authors

Piers M. Forster University of Reading UK

David W.J. Thompson Colorado State University USA

Lead Authors

Mark P. Baldwin Northwest Research Associates USA

Martyn P. Chipperfield University of Leeds UK

Martin Dameris Deutsches Zentrum für Luft- und Raumfahrt, Germany

 Institut für Physik der Atmosphäre

Joanna D. Haigh Imperial College London UK

David J. Karoly University of Melbourne Australia

Paul J. Kushner University of Toronto, Department of Physics Canada

William J. Randel National Center for Atmospheric Research, Atmospheric Chemistry Division USA

Karen H. Rosenlof NOAA ESRL Chemical Sciences Division USA

Dian J. Seidel NOAA Air Resources Laboratory USA

Susan Solomon NOAA ESRL Chemical Sciences Division USA

Coauthors

Gufran Beig Indian Institute of Tropical Meteorology India

Peter Braesicke University of Cambridge, National Centre for Atmospheric Science UK

Neal Butchart Met Office Hadley Centre UK

Nathan P. Gillett Environment Canada, Canadian Centre for Climate Modelling and Analysis Canada

Kevin M. Grise Colorado State University USA
A.6

Appendix A

Daniel R. Marsh National Center for Atmospheric Research, Atmospheric Chemistry Division USA

Charles McLandress University of Toronto Canada

T. Narayan Rao National Atmospheric Research Laboratory India

Seok-Woo Son McGill University Canada

Georgiy L. Stenchikov King Abdullah University of Science and Technology Saudi Arabia

Shigeo Yoden Kyoto University Japan

Contributors

Eugene C. Cordero San Jose State University USA

Melissa P. Free NOAA Air Resources Laboratory USA

Andreas I. Jonsson University of Toronto Canada

Jennifer Logan Harvard University USA

David Stevenson University of Edinburgh UK

CHAPTER 5

A FOCUS ON INFORMATION AND OPTIONS FOR POLICYMAKERS

Coordinating Lead Authors

John S. Daniel NOAA ESRL Chemical Sciences Division USA

Guus J.M. Velders Netherlands Environmental Assessment Agency (PBL) The Netherlands

Lead Authors

Olaf Morgenstern National Institute of Water and Atmospheric Research (NIWA) New Zealand

Darin W. Toohey University of Colorado USA

Timothy J. Wallington Ford Motor Company USA

Donald J. Wuebbles University of Illinois USA

Coauthors

Hideharu Akiyoshi National Institute for Environmental Studies Japan

Alkiviadis F. Bais Aristotle University of Thessaloniki Greece

Eric L. Fleming Science Systems and Applications, Inc. USA

Charles H. Jackman NASA Goddard Space Flight Center USA

Lambert Kuijpers Technical University, Eindhoven Center for Sustainability The Netherlands

Mack McFarland DuPont Chemicals and Fluoroproducts USA

Stephen A. Montzka NOAA ESRL Global Monitoring Division USA

Martin N. Ross The Aerospace Corporation USA

Simone Tilmes National Center for Atmospheric Research, USA

 Climate and Global Dynamics Division

Matthew B. Tully Australian Bureau of Meteorology Australia

Contributors

Stephen O. Andersen Co-Chair, Montreal Protocol Technology and USA

 Economic Assessment Panel (TEAP)

Ulrike Langematz Freie Universität Berlin Germany

Pauline M. Midgley University of Bern Switzerland

TWENTY QUESTIONS AND ANSWERS ABOUT THE OZONE LAYER: 2010 UPDATE

Coordinating Lead Authors

David W. Fahey NOAA ESRL Chemical Sciences Division USA

Michaela I. Hegglin University of Toronto Canada
A.7

Authors, Contributors, and Reviewers

Chapter Editorial Contributors

CHAPTER 1: OZONE-DEPLETING SUBSTANCES (ODSs) AND RELATED CHEMICALS

Nada Derek CSIRO Division of Atmospheric Research Australia

CHAPTER 3: FUTURE OZONE AND ITS IMPACT ON SURFACE UV

Stefanie Kremser Bodeker Scientific New Zealand

TWENTY QUESTIONS AND ANSWERS ABOUT THE OZONE LAYER: 2010 UPDATE

Jennifer Fox NOAA ESRL Chemical Sciences Division USA

Scientific Review and Adisory Group

Malcolm K.W. Ko NASA Langley Research Center USA

Theodore G. Shepherd University of Toronto Canada

Susan Solomon NOAA ESRL Chemical Sciences Division USA

Reviewers

Patricio Aceituno Universidad de Chile Chile

Ayité-Lô Nohende Ajavon Université de Lomé Togo

Hideharu Akiyoshi National Institute for Environmental Studies Japan

Daniel L. Albritton NOAA ESRL Chemical Sciences Division (retired) USA

Taofiki Aminou Université Nationale du Bénin Departement de Chimie/FAST Benin

Stephen O. Andersen Co-Chair, Montreal Protocol Technology and USA 

 Economic Assessment Panel (TEAP)

Julie M. Arblaster Australian Bureau of Meteorology / Australia/USA

 National Center for Atmospheric Research

Ghassem Asrar World Climate Research Programme / World Meteorological Organization Switzerland

Pieter J. Aucamp Ptersa South Africa

John Austin University Corporation for Atmospheric Research / USA

 NOAA Geophysical Fluid Dynamics Laboratory

Alkiviadis F. Bais Aristotle University of Thessaloniki Greece

Mark P. Baldwin Northwest Research Associates USA

Dimitris Balis Aristotle University of Thessaloniki Greece

Slimane Bekki Laboratoire Atmosphères, Milieux, Observations Spatiales France

 (LATMOS)/IPSL, CNRS

Peter F. Bernath University of York UK

Gregory E. Bodeker Bodeker Scientific New Zealand

Rumen D. Bojkov University of Dresden Germany

Janet F. Bornman University of Waikato New Zealand

Geir O. Braathen World Meteorological Organization Switzerland

Peter Braesicke University of Cambridge, National Centre for Atmospheric Science UK

Christoph Brühl Max Planck Institute for Chemistry, Mainz Germany

Claus Brüning European Commission Belgium
A.8

Appendix A

Dominik Brunner Swiss Federal Laboratories for Materials Science and Technology (Empa) Switzerland

John P. Burrows University of Bremen, Institute of Environmental Physics Germany/UK

Pablo O. Canziani Pontificia Universidad Católica Argentina / CONICET Argentina

Bruno Carli Consiglio Nazionale delle Ricerche, Institute of Applied Physics Italy

Lucy Carpenter University of York UK

Ken Carslaw University of Leeds, School of Earth and Environment UK

Marie-Lise Chanin Laboratoire Atmosphères, Milieux, Observations Spatiales France

 (LATMOS)/IPSL/UVSQ, CNRS

Martyn P. Chipperfield University of Leeds UK

Natalia E. Chubarova Moscow State University Russia

Irene Cionni Deutsches Zentrum für Luft- und Raumfahrt, Germany

 Institut für Physik der Atmosphäre

Hans Claude Deutscher Wetterdienst Germany

Cathy Clerbaux Laboratoire Atmosphères, Milieux, Observations Spatiales France

 (LATMOS)/IPSL, CNRS-INSU

Gerrie Coetzee South African Weather Service South Africa

William J. Collins Met Office Hadley Centre UK

Derek M. Cunnold Georgia Institute of Technology USA

Martin Dameris Deutsches Zentrum für Luft- und Raumfahrt, Germany 

 Institut für Physik der Atmosphäre

John S. Daniel NOAA ESRL Chemical Sciences Division USA

Martine De Mazière Belgian Institute for Space Aeronomy Belgium

Dick Derwent rdscientific UK

Panuganti Devara Indian Institute of Tropical Meteorology India

Roseanne Diab Academy of Science of South Africa South Africa

Susana B. Diaz INGEBI / CADIC / Consejo Nacional de Investigaciones Cientificas Argentina

Anne R. Douglass NASA Goddard Space Flight Center USA

Ellsworth G. Dutton NOAA ESRL Global Monitoring Division USA

Kalju Eerme Tartu Observatory Estonia

James William Elkins NOAA ESRL Global Monitoring Division USA

Andreas Engel Goethe University, Frankfurt am Main Germany

Christine A. Ennis CIRES-University of Colorado / NOAA ESRL Chemical Sciences Division USA

Veronika Eyring Deutsches Zentrum für Luft- und Raumfahrt, Germany

 Institut für Physik der Atmosphäre

David W. Fahey NOAA ESRL Chemical Sciences Division USA

Vitali E. Fioletov Environment Canada Canada

Ian Folkins Dalhousie University Canada

Piers M. Forster University of Leeds UK

James Franklin CLF-Chem Consulting Belgium

John C. Fyfe Environment Canada, Canadian Centre for Climate Modeling and Analysis Canada

Annie Gabriel Department of the Environment, Water, Heritage and the Arts Australia

Rolando R. Garcia National Center for Atmospheric Research, Atmospheric Chemistry Division USA

Marvin A. Geller Stony Brook University USA

Andrew Gettelman National Center for Atmospheric Research, Atmospheric Chemistry Division USA

Manuel Gil Instituto Nacional de Técnica Aeroespacial Spain

Sophie Godin-Beekmann Laboratoire Atmosphères, Milieux, Observations Spatiales France

 (LATMOS)/UPMC, CNRS

Marco González United Nations Environment Programme, Ozone Secretariat Kenya

Hans-F Graf University of Cambridge, National Centre for Atmospheric Science UK

Lesley Gray Reading University, National Centre for Atmospheric Science UK
A.9

Authors, Contributors, and Reviewers

Ebrahim Hajizadeh Department of Environment Iran

Neil R.P. Harris University of Cambridge, European Ozone Research Coordinating Unit UK

Dennis L. Hartmann University of Washington, Department of Atmospheric Sciences USA

Birgit Hassler CIRES-University of Colorado / NOAA ESRL Chemical Sciences Division USA

Michaela I. Hegglin University of Toronto Canada

Ernest Hilsenrath University of Maryland Baltimore County USA

David J. Hofmann NOAA ESRL Global Monitoring Division USA

Paul Horwitz United Nations Environment Programme, Ozone Secretariat USA

Robert D. Hudson University of Maryland, Department of Atmospheric and Oceanic Science USA

Mohammad Ilyas Albiruni Environment and Science Development Centre (EnviSC) Malaysia

Takashi Imamura National Institute for Environmental Studies Japan

Ivar S.A. Isaksen University of Oslo Norway

Kenneth Jucks NASA Headquarters USA

David J. Karoly University of Melbourne Australia

Philippe Keckhut Laboratoire Atmosphères, Milieux, Observations Spatiales France

 (LATMOS)/IPSL/UVSQ, CNRS

Douglas E. Kinnison National Center for Atmospheric Research, Atmospheric Chemistry Division USA

Malcolm K.W. Ko NASA Langley Research Center USA

Kunihiko Kodera Nagoya University / Meteorological Research Institute, Japan

 Japan Meteorological Agency

Takashi Koide Japan Meteorological Agency Japan

Ninong Komala National Institute of Aeronautics and Space Indonesia

Yutaka Kondo University of Tokyo Japan

Karin Kreher National Institute of Water and Atmospheric Research (NIWA) New Zealand

Mark Kroon Royal Netherlands Meteorological Institute (KNMI) The Netherlands

Kirstin Krüger IFM-GEOMAR / Leibniz-Institute of Marine Sciences at Kiel University Germany

Lambert Kuijpers Technical University, Eindhoven Center for Sustainability The Netherlands

Michael J. Kurylo Goddard Earth Sciences and Technology Center / USA

 University of Maryland Baltimore County

Paul J. Kushner University of Toronto, Department of Physics Canada

Esko Kyrö Finnish Meteorological Institute Finland

Shyam Lal Physical Research Laboratory India

Tom Land U.S. Environmental Protection Agency USA

Igor Larin Russian Academy of Sciences Russia

Katharine Law Laboratoire Atmosphères, Milieux, Observations Spatiales France

 (LATMOS)/UPMC/UVSQ, CNRS

Jos Lelieveld Max Planck Institute for Chemistry, Mainz Germany

Sasha Madronich National Center for Atmospheric Research, Atmospheric Chemistry Division USA

Gloria L. Manney Jet Propulsion Laboratory, California Institute of Technology / USA

 New Mexico Institute of Mining and Technology

Alistair J. Manning Met Office UK

Elisa Manzini Max-Planck-Institut für Meteorologie – Hamburg Germany

John C. McConnell York University, Department of Earth and Space Science and Engineering Canada

C. Thomas McElroy Environment Canada Canada

Mack McFarland DuPont Chemicals and Fluoroproducts USA

Norman McFarlane SPARC International Project Office Canada

Danny McKenna Oak Ridge National Laboratory USA

Richard L. McKenzie National Institute of Water and Atmospheric Research (NIWA) New Zealand

Charles McLandress University of Toronto Canada
A.10

Appendix A

Mario J. Molina University of California, San Diego and Mario Molina Center USA/Mexico

 for Strategic Studies in Energy and the Environment

Stephen A. Montzka NOAA ESRL Global Monitoring Division USA

Rolf Müller Forschungszentrum Jülich Germany

Cathrine Lund Myhre Norwegian Institute for Air Research (NILU) Norway

Hideaki Nakane National Institute for Environmental Studies Japan

Eric R. Nash Science Systems and Applications, Inc. USA

Cindy Newberg U.S. Environmental Protection Agency USA

Mike Newchurch University of Alabama in Huntsville USA

Paul A. Newman NASA Goddard Space Flight Center USA

Ole John Nielsen University of Copenhagen Denmark

Alan O’Neill National Centre for Earth Observation / University of Reading UK

Samuel J. Oltmans NOAA ESRL Global Monitoring Division USA

Luke D. Oman NASA Goddard Space Flight Center USA

Edward A. Parson University of Michigan USA

Nigel D. Paul Lancaster University UK

Stuart A. Penkett University of East Anglia, Norwich UK

Judith Perlwitz CIRES-University of Colorado / NOAA ESRL Physical Sciences Division USA

Thomas Peter Swiss Federal Institute of Technology, Switzerland

 Institute for Atmospheric and Climate Science

Giovanni Pitari Università degli Studi di L’Aquila, Dipartimento di Fisica Italy

R. Alan Plumb Massachusetts Institute of Technology USA

David Plummer Environment Canada, Canadian Centre for Climate Modelling and Analysis Canada

Jean-Pierre Pommereau Laboratoire Atmosphères, Milieux, Observations Spatiales (LATMOS), CNRS France

Michael Ponater Deutsches Zentrum für Luft- und Raumfahrt, Germany

 Institut für Physik der Atmosphäre

Lamont R. Poole Science Systems and Applications, Inc. USA

Robert W. Portmann NOAA ESRL Chemical Sciences Division USA

Michael J. Prather University of California, Irvine / Earth System Science Department USA

Ronald G. Prinn Massachusetts Institute of Technology, Center for Global Change Science USA

John A. Pyle University of Cambridge, National Centre for Atmospheric Science UK

V. Ramaswamy NOAA Geophysical Fluid Dynamics Laboratory USA

William J. Randel National Center for Atmospheric Research, Atmospheric Chemistry Division USA

A.R. Ravishankara NOAA ESRL Chemical Sciences Division USA

Claire E. Reeves University of East Anglia / National Centre for Atmospheric Sciences UK

Stefan Reimann Swiss Federal Laboratories for Materials Science and Technology (Empa) Switzerland

Markus Rex Alfred Wegener Institute for Polar and Marine Research, Potsdam Germany

Martin Riese Research Centre Jülich, Institute for Chemistry and Germany

 Dynamics of the Geosphere

Vincenzo Rizi Università degli Studi di L’Aquila Italy

Alan Robock Rutgers University USA

Howard K. Roscoe British Antarctic Survey UK

Vladimir Ryabinin World Climate Research Programme/World Meteorological Organization Switzerland

Ross J. Salawitch University of Maryland USA

Michelle L. Santee Jet Propulsion Laboratory, California Institute of Technology USA

K. Madhava Sarma Consultant India

Robert Sausen Deutsches Zentrum für Luft- und Raumfahrt, Germany

 Institut für Physik der Atmosphäre

Adam A. Scaife Met Office Hadley Centre UK
A.11

Authors, Contributors, and Reviewers

Ulrich Schmidt Goethe University, Institute for Atmosphere and Environment Germany

Ulrich Schumann Deutsches Zentrum für Luft- und Raumfahrt, Germany

 Institut für Physik der Atmosphäre

Dian J. Seidel NOAA Air Resources Laboratory USA

Megumi Seki United Nations Environment Programme, Ozone Secretariat Kenya

Jonathan Shanklin British Antarctic Survey UK

Wafik M. Sharobiem Egyptian Meteorological Authority, Scientific Research Department Egypt

Theodore G. Shepherd University of Toronto Canada

Keith P. Shine University of Reading, Department of Meteorology UK

Masato Shiotani Kyoto University, Research Institute for Sustainable Humanosphere Japan

Michael Sigmond University of Toronto Canada

Peter Simmonds University of Bristol, School of Chemistry UK

Anne Smith National Center for Atmospheric Research, Atmospheric Chemistry Division USA

Susan Solomon NOAA ESRL Chemical Sciences Division USA

Johannes Staehelin Swiss Federal Institute of Technology – Zurich Switzerland

Wolfgang Steinbrecht Deutscher Wetterdienst Germany

Georgiy L. Stenchikov King Abdullah University of Science and Technology Saudi Arabia

Richard S. Stolarski NASA Goddard Space Flight Center USA

Frode Stordal University of Oslo, Department of Geosciences Norway

Fred Stroh Forschungszentrum Jülich GmbH, Institute for Chemistry and Germany

 Dynamics of the Geosphere

William T. Sturges University of East Anglia UK

Kenshi Takahashi Kyoto University Japan

David W. Tarasick Environment Canada, Air Quality Research Division Canada

Said Ali Thaoubane Université des Comores Comores

David W.J. Thompson Colorado State University USA

Simone Tilmes National Center for Atmospheric Research, USA

 Climate and Global Dynamics Division

Matthew B. Tully Australian Bureau of Meteorology Australia

Ronald Van der A Royal Netherlands Meteorological Institute (KNMI) The Netherlands

Karel Vanicek Czech Hydrometeorological Institute Czech Republic

Guus J.M. Velders Netherlands Environmental Assessment Agency (PBL) The Netherlands

Marc von Hobe Forschungszentrum Jülich Germany

Darryn W. Waugh Johns Hopkins University USA

Elizabeth C. Weatherhead CIRES-University of Colorado / NOAA ESRL Global Systems Division USA

Ann R. Webb University of Manchester, School of Earth, Atmospheric, and UK

 Environmental Sciences

Mark Weber University of Bremen, Institute of Environmental Physics Germany

Ray F. Weiss Scripps Institution of Oceanography, USA

 University of California, San Diego

Donald J. Wuebbles University of Illinois USA

Shigeo Yoden Kyoto University Japan

Yoko Yokouchi National Institute for Environmental Studies Japan

Durwood Zaelke Institute for Governance and Sustainable Development USA

Rodolphe Zander University of Liège Belgium

Christos S. Zerefos International Ozone Commission; Academy of Athens Greece

Lingxi Zhou Chinese Meteorological Administration, China

 Chinese Academy of Meteorological Sciences
A.12

Appendix A

Ozone Peer-Review Meeting

Les Diablerets, Switzerland

28 June – 2 July 2010

Patricio Aceituno Universidad de Chile 

Chile

Ayité-Lô Nohende Ajavon Université de Lomé Togo

Stephen O. Andersen Co-Chair, Montreal Protocol Technology and USA

 Economic Assessment Panel (TEAP) 

Pieter J. Aucamp Ptersa South Africa

Alkiviadis F. Bais Aristotle University of Thessaloniki Greece

Slimane Bekki Laboratoire Atmosphères, Milieux, Observations Spatiales France

 (LATMOS)/IPSL, CNRS

Peter F. Bernath University of York UK

Gregory E. Bodeker Bodeker Scientific New Zealand

Janet F. Bornman University of Waikato New Zealand

Geir O. Braathen World Meteorological Organization Switzerland

Peter Braesicke University of Cambridge, National Centre for Atmospheric Science UK

John P. Burrows University of Bremen, Institute of Environmental Physics Germany/UK

Lucy Carpenter University of York UK

Martyn P. Chipperfield University of Leeds UK

Martin Dameris Deutsches Zentrum für Luft- und Raumfahrt, Germany

 Institut für Physik der Atmosphäre

John S. Daniel NOAA ESRL Chemical Sciences Division USA

Susana B. Diaz INGEBI / CADIC / Consejo Nacional de Investigaciones Cientificas Argentina

Anne R. Douglass NASA Goddard Space Flight Center USA

Andreas Engel Goethe University, Frankfurt am Main Germany

Christine A. Ennis CIRES-University of Colorado / NOAA ESRL Chemical Sciences Division USA

Veronika Eyring Deutsches Zentrum für Luft- und Raumfahrt, Germany

 Institut für Physik der Atmosphäre

David W. Fahey NOAA ESRL Chemical Sciences Division USA

Vitali E. Fioletov Environment Canada Canada

Piers M. Forster University of Leeds UK

Marvin A. Geller Stony Brook University USA

Sophie Godin-Beekmann Laboratoire Atmosphères, Milieux, Observations Spatiales France

 (LATMOS)/UPMC, CNRS

Marco González United Nations Environment Programme, Ozone Secretariat Kenya

Neil R.P. Harris University of Cambridge, European Ozone Research Coordinating Unit UK

by Jean-Pierre Pommereau, June 2010
A.13

Authors, Contributors, and Reviewers

Michaela I. Hegglin University of Toronto Canada

Kenneth Jucks NASA Headquarters USA

David J. Karoly University of Melbourne Australia

Malcolm K.W. Ko NASA Langley Research Center USA

Takashi Koide Japan Meteorological Agency Japan

Kirstin Krüger IFM-GEOMAR / Leibniz-Institute of Marine Sciences at Kiel University Germany

Lambert Kuijpers Technical University, Eindhoven Center for Sustainability The Netherlands

Michael J. Kurylo Goddard Earth Sciences and Technology Center / USA

 University of Maryland Baltimore County

Paul J. Kushner University of Toronto, Department of Physics Canada

Igor Larin Russian Academy of Sciences Russia

Elisa Manzini Max-Planck-Institut für Meteorologie – Hamburg Germany

Mack McFarland DuPont Chemicals and Fluoroproducts USA

Stephen A. Montzka NOAA ESRL Global Monitoring Division USA

Rolf Müller Forschungszentrum Jülich Germany

Hideaki Nakane National Institute for Environmental Studies Japan

Paul A. Newman NASA Goddard Space Flight Center USA

Alan O’Neill National Centre for Earth Observation / University of Reading UK

Judith Perlwitz CIRES-University of Colorado / NOAA ESRL Physical Sciences Division USA

Thomas Peter Swiss Federal Institute of Technology, Institute for Switzerland

 Atmospheric and Climate Science

Jean-Pierre Pommereau Laboratoire Atmosphères, Milieux, Observations Spatiales (LATMOS), CNRS France

Robert W. Portmann NOAA ESRL Chemical Sciences Division USA

Michael J. Prather University of California, Irvine / Earth System Science Department USA

John A. Pyle University of Cambridge, National Centre for Atmospheric Science UK

V. Ramaswamy NOAA Geophysical Fluid Dynamics Laboratory USA

A.R. Ravishankara NOAA ESRL Chemical Sciences Division USA

Claire E. Reeves University of East Anglia / National Centre for Atmospheric Sciences UK

Stefan Reimann Swiss Federal Laboratories for Materials Science and Technology (Empa) Switzerland

Ross J. Salawitch University of Maryland USA

Dian J. Seidel NOAA Air Resources Laboratory USA

Wafik M. Sharobiem Egyptian Meteorological Authority, Scientific Research Department Egypt

Theodore G. Shepherd University of Toronto Canada

Peter Simmonds University of Bristol, School of Chemistry UK

Anne Smith National Center for Atmospheric Research, Atmospheric Chemistry Division USA

Susan Solomon NOAA ESRL Chemical Sciences Division USA

Georgiy L. Stenchikov King Abdullah University of Science and Technology Saudi Arabia

Richard S. Stolarski NASA Goddard Space Flight Center USA

William T. Sturges University of East Anglia UK

Said Ali Thaoubane Université des Comores Comores

David W.J. Thompson Colorado State University USA

Matthew B. Tully Australian Bureau of Meteorology Australia

Guus J.M. Velders Netherlands Environmental Assessment Agency (PBL) The Netherlands

Darryn W. Waugh Johns Hopkins University USA

Ann R. Webb University of Manchester, School of Earth, Atmospheric, and UK

 Environmental Sciences

Ray F. Weiss Scripps Institution of Oceanography, University of California, San Diego USA

Donald J. Wuebbles University of Illinois USA

Lingxi Zhou Chinese Meteorological Administration, China

 Chinese Academy of Meteorological Sciences
A.14

Sponsoring Organizations Liaisons

Geir O. Braathen World Meteorological Organization Switzerland

Marco González United Nations Environment Programme Kenya

A.R. Ravishankara National Oceanic and Atmospheric Administration USA

Kenneth Jucks National Aeronautics and Space Administration USA

Claus Brüning European Commission Belgium

Assessment Coordinator and Technical Editor

Christine A. Ennis CIRES/NOAA ESRL Chemical Sciences Division USA

Publication/Graphics Design and Layout

Debra Dailey-Fisher (Publication, Design, Layout, & Graphics Lead) NOAA ESRL CSD USA

Dennis Dickerson (Graphics & Layout Design, “Twenty Questions”) Respond Grafiks USA

Albert D. Romero (Consulting and Support) NOAA Boulder Facilities Operations Division USA

Editorial Assistance

Debra Dailey-Fisher NOAA ESRL Chemical Sciences Division USA

Jessica N. Lucas NOAA Earth System Research Laboratory USA

Reference Research and Editing

Barbara Keppler NOAA ESRL Chemical Sciences Division USA

Jessica N. Lucas NOAA Earth System Research Laboratory USA

Debra R. Wilson NOAA ESRL Chemical Sciences Division USA

Kristen McCormack Boulder High School USA

Mary Gutierrez NOAA ESRL Chemical Sciences Division USA

Debra Dailey-Fisher NOAA ESRL Chemical Sciences Division USA

William L. Fisher Fisher Enterprises USA

Christine A. Ennis CIRES/NOAA ESRL Chemical Sciences Division USA

Conference Coordination and Documentation

Christine A. Ennis CIRES/NOAA ESRL Chemical Sciences Division USA

Geir O. Braathen World Meteorological Organization Switzerland

Debra Dailey-Fisher NOAA ESRL Chemical Sciences Division USA

John A. Pyle University of Cambridge UK

Norman McFarlane SPARC International Project Office Canada

Theodore G. Shepherd University of Toronto Canada

Conference Support

Jeanne S. Waters NOAA ESRL Chemical Sciences Division USA

Kathy A. Thompson CSC USA

Debra Dailey-Fisher NOAA ESRL Chemical Sciences Division USA

Jennifer Fox CIRES/NOAA ESRL Chemical Sciences Division USA

Chantal Renaudot World Meteorological Organization Switzerland

Rose M. Kendall CSC USA

Ruth Batten UNEP Ozone Secretariat Kenya

Ann Gachingiri UNEP Ozone Secretariat Kenya

Peter Braesicke University of Cambridge UK

Tina Jost University of Cambridge UK

Victoria De Luca SPARC International Project Office UK

David Dokken IPCC Working Group II Technical Support Unit USA

Kyle Terran University Corporation for Atmospheric Research USA

Eric Kissel IPCC Working Group II Technical Support Unit USA 

Computing and Networking Support

Richard J. Tisinai and Jennifer Fox CIRES/NOAA ESRL Chemical Sciences Division USA

Document Distribution

Jeanne S. Waters NOAA ESRL Chemical Sciences Division USA

Debra Dailey-Fisher NOAA ESRL Chemical Sciences Division USA

Chantal Renaudot World Meteorological Organization Switzerland

Ruth Batten UNEP Ozone Secretariat Kenya
B.1

Appendix B

Major Acronyms and Abbreviations

A1 baseline (or most likely) halocarbon scenario of the 2006 Ozone Assessment 

A1B scenario of the IPCC Special Report on Emissions Scenarios (SRES)

A5 Article 5 countries of the Montreal Protocol

AAE Absorption Angstrom Exponent

AASE Airborne Arctic Stratospheric Expedition (NASA)

Ab baseline halocarbon scenario of the 2002 Ozone Assessment

ACE-FTS Atmospheric Chemistry Experiment Fourier Transform Spectrometer

AERONET AErosol RObotic NETwork

AGAGE Advanced Global Atmospheric Gases Experiment

AGWP Absolute Global Warming Potential

AMA average mean age 

AMSU Advanced Microwave Sounding Unit

AMTRAC Atmospheric Model with TRansport and Chemistry (Table 3-1)

AO Arctic Oscillation

AOGCM atmosphere-ocean general circulation model

AOT aerosol optical thickness

AR4 IPCC Fourth Assessment Report

ARCTAS Arctic Research of the Composition of the Troposphere from Aircraft and Satellites

ARCPAC Aerosol, Radiation, and Cloud Processes affecting Arctic Climate

ATLAS Atmospheric Laboratory for Applications and Science

ATMOS Atmospheric Trace Molecule Spectroscopy

AVE Aura Validation Experiment

B1 a lower-emissions scenario of the IPCC Special Report on Emissions Scenarios (SRES)

B2 scenario of the IPCC Special Report on Emissions Scenarios (SRES)

BC black carbon

BCE Before the Common Era

BDC Brewer-Dobson circulation

BL boundary layer

BP British Petroleum

BUV Backscatter (or Backscattered) Ultraviolet (spectrometer)

C Celsius (unit of temperature)

CADIC Centro Austral de Investigaciones Cientificas (Argentina)

CALIOP Cloud-Aerosol Lidar with Orthogonal Polarization

CALIPSO Cloud-Aerosol Lidar and Infrared Pathfinder Satellite Observation

CAM Community Atmosphere Model (Table 3-1)

CATO Candidoz Assimilated Three-dimensional Ozone

CCM chemistry-climate model

CCMVal Chemistry-Climate Model (CCM) Validation Activity

CCSP Climate Change Science Program (United States)

CCSR-NIES Center for Climate-Systems Research – National Institute for Environmental Studies CCM (Table 

3-1)

CDM Clean Development Mechanism

CE Common Era
B.2

Appendix B

CFC chlorofluorocarbon

CHAMP Challenging Minisatellite Payload

CIE Commission Internationale de l’Éclairage (France)

CIRES Cooperative Institute for Research in Environmental Sciences (United States)

cm centimeters (unit of length)

CMAM Canadian Middle Atmosphere Model (Table 3-1)

CMF cloud modification factor

CMIP Coupled Model Intercomparison Project

CNRM Centre National de Recherches Météorologiques (France)

CNRM-ACM CNRM ARPEGE-Climate model (Table 3-1)

CNRS Centre National de la Recherche Scientifique (France)

CONICET Consejo de Investigaciones Cientificas y Técnicas (Argentina)

CO2-eq carbon dioxide equivalents

COS carbonyl sulfide (also OCS)

CPT cold point tropopause

CR-AVE Costa Rica-Aura Validation Experiment

CSIRO Commonwealth Scientific and Industrial Research Organisation (Australia)

CTM chemistry and transport model

CUE critical-use exemption

CUSUM Cumulative Sum of Residuals

DFA difluoroacetic acid

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

DU Dobson unit

E39CA a coupled chemistry-climate model of DLR (Table 3-1)

EC European Commission

ECl equivalent chlorine

ECMWF European Centre for Medium-Range Weather Forecasts (United Kingdom)

EEAP Environmental Effects Assessment Panel

EECl equivalent effective chlorine

EESC equivalent effective stratospheric chlorine

EMAC ECHAM/MESSy Atmospheric Chemistry model of FUB (Table 3-1)

ENSO El Niño-Southern Oscillation

Envisat Environmental Satellite

EP Eliassen-Palm

EPA Environmental Protection Agency (United States)

ER-2 Earth Resources-2 (aircraft)

ERA ECMWF Re-Analysis

ERA-40 ECMWF 40-year reanalysis

ERAI ECMWF Interim Re-Analysis

ERS European Remote Sensing Satellite

ESC equivalent stratospheric chlorine

ESRL Earth System Research Laboratory (NOAA)

EU European Union

EUPLEX European Polar Stratospheric Cloud and Lee Wave Experiment
B.3

Acronyms

fGHG sensitivity simulation with fixed greenhouse gas levels (see Chapter 3 and Table 3-2)

fODS sensitivity simulation with fixed ozone-depleting substance levels (see Chapter 3 and Table 3-2)

FT free troposphere

FTIR Fourier transform infrared

FUB Freie Universität Berlin (Germany)

GCM general circulation model

GEOS Goddard Earth Observing System

GEOSCCM Goddard Earth Observing System Chemistry-Climate Model (Table 3-1)

GFDL Geophysical Fluid Dynamics Laboratory (NOAA)

Gg Gigagrams (109 grams) (unit of mass)

GHG greenhouse gas

GHG-x a greenhouse gas sensitivity simulation (see Chapter 3 and Table 3-2)

GISS Goddard Institute for Space Studies (NASA)

GJ gigajoule (109 joules) (unit of energy)

GMD Global Monitoring Division (NOAA/ESRL)

GOME Global Ozone Monitoring Experiment

GREET Greenhouse Gases, Regulated Emissions, and Energy Use in Transportation (model)

GSFC Goddard Space Flight Center (NASA)

Gt gigatonnes

GtCO2-eq gigatonnes of carbon dioxide equivalents

GWP Global Warming Potential

HadAT Hadley Centre radiosonde temperature product

HALOE Halogen Occultation Experiment

HAMMONIA HAMburg Model of the Neutral and Ionized Atmosphere

HC hydrocarbon

HCFC hydrochlorofluorocarbon

HFC hydrofluorocarbon

HFE hydrofluorinated ether or hydrofluoroether

HFO hydrofluoro-olefin

hPa hectoPascal (102 Pascal) (unit of pressure)

IGRA Integrated Global Radiosonde Archive (NOAA)

INGEBI Instituto de Investigaciones en Ingeniería Genética y Biología Molecular (Argentina)

IPA Institut für Physik der Atmosphäre (DLR) (Germany)

IPCC Intergovernmental Panel on Climate Change

IPSL Institut Pierre-Simon Laplace (France)

IR infrared

IUK Iterative Universal Kriging

IUP Institute of Environmental Physics, University of Bremen (Germany)

IUPAC International Union of Pure and Applied Chemistry

J joule (unit of energy)

JCET Joint Center for Earth Systems Technology (United States)

JJA June-July-August

JMA Japan Meteorological Agency (Japan)

JPL Jet Propulsion Laboratory (NASA)

JRA25 Japanese 25-year Reanalysis Project
B.4

Appendix B

K Kelvin (unit of temperature)

kcal kilocalories (103 calories) (unit of energy)

kg kilogram (103 grams) (unit of mass)

km kilometer (103 meters) (unit of length)

Kt kilotons (103 tons) (unit of mass)

LLGHG long-lived greenhouse gases

LMDZrepro general circulation model of the Laboratory of Dynamic Meteorology (IPSL) (Table 3-1)

LMS lowermost stratosphere

LS lower stratosphere

LZRH level of zero clear-sky radiative heating

m meter (unit of length)

MAC mobile air conditioning

MAM March-April-May

MATCH Model for Atmospheric Transport and Chemistry

MBL marine boundary layer

MetOp Meteorological Operational satellite

MFA monofluoroacetic acid

MFRSR Multi-Filter Rotating Shadowband Radiometer

mg milligram (10-3 grams) (unit of mass)

MILAGRO Megacity Initiative: Local and Global Research Observations

MIPAS Michelson Interferometer for Passive Atmospheric Sounding 

MIT Massachusetts Institute of Technology

MLD mixed layer depth

MJ megajoule (106 joules) (unit of energy)

MLR multiple linear regression

MLS Microwave Limb Sounder

mm millimeters (10-3 meters) (unit of length)

μm micrometer; micron (10-6 meters) (unit of length)

MMBtu million British thermal units (1MMBtu = 1.055 gigajoules)

MOD merged ozone data set

MODIS Moderate Resolution Imaging Spectroradiometer

mol mole (unit, amount of substance)

molec molecule

MOZAIC Measurement of Ozone and Water Vapor by Airbus In-Service Aircraft

MOZART Model for Ozone and Related Chemical Tracers

MRI Meteorological Research Institute chemistry-climate model (Japan) (Table 3-1)

MSG Meteosat Second Generation

MSU Microwave Sounding Unit

MtCFC-11-eq megatonnes of CFC-11 equivalents

Mt megatonne

mW milliWatt (10-3 Watts)

NAM Northern Annular Mode

NAO North Atlantic Oscillation

NASA National Aeronautics and Space Administration (United States)

NAT nitric acid trihydrate

NCAR National Center for Atmospheric Research (United States)

NCEP National Centers for Environmental Prediction (United States)

ng nanogram (10-9 grams) (unit of mass)
B.5

Acronyms

NH Northern Hemisphere

NIES National Institute for Environmental Studies (Japan)

NIWA National Institute of Water and Atmospheric Research (New Zealand)

nm nanometers (10-9 meters) (unit of length)

NOAA National Oceanic and Atmospheric Administration (United States)

NPLS nonparametric least-squares fit

1-D one-dimensional

OCS carbonyl sulfide (also COS)

ODP Ozone Depletion Potential

ODS ozone-depleting substance

OMI Ozone Monitoring Instrument

OSIRIS Optical Spectrograph and InfraRed Imager System

PBL planetary boundary layer

PEM Pacific Exploratory Mission

PFC perfluorocarbon

Pg petagram (1015 grams) (unit of mass)

PG product gas

PGI product gas injection

PhotoComp photolysis intercomparison

POAM Polar Ozone and Aerosol Measurement

POCP Photochemical Ozone Creation Potential

ppb part per billion

ppbv part per billion by volume

ppm part per million

ppmv part per million by volume

ppt part per trillion

pptv part per trillion by volume

PSC polar stratospheric cloud

PSS photochemical steady state

PTFE polytetrafluoroethylene

PWLT piecewise linear trend

QBO quasi-biennial oscillation

QPS quarantine and pre-shipment

RAF radiation amplification factor

RAOB RAwinsonde OBservation

RAOBCORE Radiosonde Observation Correction using Reanalyses

RATPAC Radiosonde Atmospheric Temperature Products for Assessing Climate

RCP Representative Concentration Pathway

RECON a near-surface temperature reconstruction

REF-B2 reference “future” simulation of SPARC CCMVal-2 (see Table 3-2)

RF radiative forcing

RICH Radiosonde Innovation Composite Homogenization

RO radio occultation

RSS Remote Sensing Systems Inc.
B.6

Appendix B

s second

SAD surface area density

SAGE Stratospheric Aerosol and Gas Experiment

SAM Southern Annular Mode

SAM II Stratospheric Aerosol Measurement II

SAOZ Système d’Analyse par Observation Zénithale

SAP Scientific Assessment Panel

SBUV/SBUV2 Solar Backscatter (or Backscattered) Ultraviolet (spectrometer)

SCIAMACHY Scanning Imaging Absorption Spectrometer for Atmospheric Chartography

SCISAT a Canadian satellite also known as Atmospheric Chemistry Experiment (ACE)

SE southeast

SG source gas

SGI source gas injection

SH Southern Hemisphere

SIC sea ice concentration

SLIMCAT Single-Layer Isentropic Model of Chemistry and Transport

SOCOL modeling tool for studies of Solar-Climate-Ozone Links (Table 3-1)

SOGE System for Observation of Halogenated Greenhouse Gases in Europe

SOLVE SAGEIII (Stratospheric Aerosol and Gas Experiment III) Ozone Loss and Validation Experiment 

(NASA)

SON September-October-November

SPARC Stratospheric Processes and Their Role in Climate (WCRP)

sr steradian (unit of solid angle)

SRES Special Report on Emissions Scenarios (IPCC)

SRM solid rocket motor

SSA stratospheric sulfate aerosol

SSA single scattering albedo 

SST sea surface temperature

SSU Stratospheric Sounding Unit

SSW sudden stratospheric warming

STAR System for Transfer of Atmospheric Radiation

STE stratosphere-troposphere exchange

STS supercooled ternary solution 

SZA solar zenith angle

2-D two-dimensional

3-D three-dimensional

TC4 Tropical Composition, Cloud and Climate Coupling mission

TCO tropospheric column ozone

TEAP Technology and Economic Assessment Panel (UNEP)

TFA trifluoroacetic acid

Tg teragrams (1012 grams) (unit of mass; equivalent to megatonne)

THESEO Third European Stratospheric Experiment on Ozone

T-M tropical-midlatitude

TOMCAT Toulouse Off-line Model of Chemistry and Transport

TOMS Total Ozone Mapping Spectrometer

TSAM time series additive model

TST troposphere-to-stratosphere transport

TTL tropical tropopause layer

TUV Tropospheric Ultraviolet-Visible model
B.7

Acronyms

UAH University of Alabama–Huntsville

UARS Upper Atmosphere Research Satellite

UBA Umweltbundesamt (Germany Federal Environment Agency)

UCI University of California, Irvine

UDMH unsymmetrical dimethylhydrazine

UEA University of East Anglia (United Kingdom)

UIUC University of Illinois at Urbana-Champaign

UK United Kingdom

UKCA U.K. Chemistry and Aerosols chemistry-climate model (Table 3-1)

ULAQ University of L’Aquila chemistry-climate model (Italy) (Table 3-1)

UMSLIMCAT Unified Model Single-Layer Isentropic Model of Chemistry and Transport CCM (Table 3-1)

UNEP United Nations Environment Programme

UNFCCC United Nations Framework Convention on Climate Change

UPMC Université Pierre et Marie Curie (France)

US, USA United States of America

USDA United States Department of Agriculture

UT upper troposphere

UTLS upper troposphere/lower stratosphere

UV ultraviolet

UV-A ultraviolet-A (315–400 nm)

UV-B ultraviolet-B (280–315 nm)

UVSQ Université Versailles Saint-Quentin (France)

VSL very short-lived

VSLS very short-lived substance

W Watt (unit of energy)

WACCM Whole-Atmosphere Community Climate Model (Table 3-1)

WCRP World Climate Research Programme

WFDOAS Weighting Function Differential Optical Absorption Spectroscopy

WMO World Meteorological Organization

C.1

Appendix C

Major Chemical Formulae and Nomenclature

from this Assessment

Halogen-Containing Species

Cl atomic chlorine Br atomic bromine

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 ([ClO] + 2×[ClOOCl]) BrOx bromine radicals

OClO chlorine 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

F atomic fluorine I atomic iodine

F2 molecular fluorine I2 molecular iodine

Fy total inorganic fluorine Iy total inorganic iodine

HF hydrogen fluoride (hydrofluoric acid) IO iodine monoxide

 IOx iodine radicals

OIO iodine dioxide

HOI hypoiodous acid

 

SF6 sulfur hexafluoride

SO2F2 sulfuryl fluoride

NF3 nitrogen trifluoride

PBr3 phosphorus tribromide
C.2

Appendix C

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-123a CHClFCF2Cl

CFC-114 CClF2CClF2 HCFC-123b CHF2CCl2F

CFC-114a CCl2FCF3 HCFC-124 CHClFCF3

CFC-115 CClF2CF3 HCFC-124a CHF2CClF2

 HCFC-133a CH2ClCF3

Halons HCFC-141b CH3CCl2F

halon-1202 CBr2F2 HCFC-142b CH3CClF2

halon-1211 CBrClF2 HCFC-225ca CHCl2CF2CF3

halon-1301 CBrF3 HCFC-225cb CHClFCF2CClF2

halon-2402 CBrF2CBrF2 HCFC-234fb CF3CH2CCl2F

halon-2311 CHBrClCF3 HCFC-243cc CH3CF2CCl2F

 (Halothane)

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

Chlorocarbons romocarbons

CH3Cl methyl chloride, chloromethane CH3Br methyl bromide, bromomethane

CH2Cl2 methylene chloride, dichloromethane CH2Br2 methylene bromide, dibromomethane

CHCl3 chloroform, trichloromethane CHBr3 bromoform, tribromomethane

CCl4 carbon tetrachloride CH3CH2Br, C2H5Br ethyl bromide, bromoethane

CHClCCl2 trichloroethylene, trichloroethene CH2BrCH2Br 1,2 dibromoethane

CCl2CCl2 tetrachloroethene, perchloroethene CH3CH2CH2Br, n-propyl bromide, n-PB, 

CH3CH2Cl, C2H5Cl ethyl chloride, chloroethane n-C3H7Br 1-bromopropane

CH2ClCH2Cl 1,2 dichloroethane COBr2 carbonyl bromide

CH3CCl3 methyl chloroform

CH3CHClCH3 isopropylchloride, 2-chloropropane

CH3CH2CH2Cl n-propyl chloride, 1-chloropropane

COCl2, Cl2C(O) phosgene, carbonyl chloride
C.3

Chemical Formulae

Iodocarbons Fluorocarbons

CH3I methyl iodide, iodomethane CF4 (PFC-14) perfluoromethane, 

CH2I2 diiodomethane carbon tetrafluoride

CH3CH2I, C2H5I ethyl iodide, iodoethane C2F6, CF3CF3 (PFC-116) perfluoroethane

CH3CHICH3 isopropyl iodide, 2-iodopropane C3F8, CF3CF2CF3 (PFC-218) perfluoropropane

CH3CH2CH2I, n-propyl iodide, 1-iodopropane c-C3F6 (PFC-C216) perfluorocyclopropane

n-C3H7I C4F10 (PFC-31-10) perfluorobutane

 c-C4F8 (PFC-C318) perfluorocyclobutane

Others C5F12 (PFC-41-12) perfluoropentane

CHBr2Cl dibromochloromethane C6F14 (PFC-51-14) perfluorohexane

CH2BrCl bromochloromethane C7H16 (PFC-61-16) perflouoroheptane

CHBrCl2 bromodichloromethane C10F18 perfluorodecalin

CH2BrI bromoiodomethane COF2 carbonyl fluoride

CHBrF2 bromodifluoromethane CH2FC(O)OH monofluoroacetic acid (MFA)

CH2ClI chloroiodomethane CHF2C(O)OH difluoroacetic acid (DFA)

CF3I trifluoroiodomethane CF3C(O)OH trifluoroacetic acid (TFA)

CF3CF2CF2I 1-iodo-heptafluoropropane

COClF chlorofluorocarbonyl 

SF5CF3 trifluoromethylsulfurpentafluoride

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(1D), O3) or HOx odd hydrogen (H, OH, HO2, H2O2)

oxidant (O3 + NO2) 

N atomic nitrogen HNO2, HONO nitrous acid

N2 molecular nitrogen HOONO pernitrous acid

N2O nitrous oxide HNO3 nitric acid

NO nitric oxide NH3 ammonia

NO2 nitrogen dioxide NH4NO3 ammonium nitrate

NO3 nitrogen trioxide, nitrate radical 

N2O5 dinitrogen pentoxide

NOx nitrogen oxides (NO + NO2)

NOy total reactive nitrogen 

 (usually includes NO, NO2, NO3,

 N2O5, ClONO2, HNO4, HNO3)

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 CO2 carbon dioxide

CO carbon monoxide

CH4 methane CH3OH methyl alcohol, methanol

CH3CH3 ethane

CH3CH2CH3 propane