The First State of the 

Carbon Cycle Report

The North American Carbon Budget 

and Implications for the

Global Carbon Cycle

U.S. Climate Change Science Program

Synthesis and Assessment Product 2.4

November 2008

Trends in Emissions of 

Ozone-Depleting Substances, 

Ozone Layer Recovery, and 

Implications for Ultraviolet 

Radiation Exposure
EDITORIAL TEAM

Chief Editor: ..................................................................................XXX

Associate Editors: .........................................................................xxx

Graphic Design: ............................................................................xxx

Technical Support: ........................................................................xxx

FEDERAL EXECUTIVE TEAM

Director, Climate Change Science Program: ................................t

Acting Director, Climate Change Science Program Office: .........Peter A. Schultz

NOAA Assistant Administrator, 

Program Planning and Integration Office: ....................................Mary M. Glackin

Director, NOAA Climate Goal: .....................................................Chester J. Koblinsky

Director, NOAA National Climatic Data Center;

NOAA Program Manager for Climate 

Observations and Analysis: ............................................................Thomas R. Karl

Program Manager of the NOAA Climate Program 

Office Climate Change Data and Detection Program; 

Federal Advisory Committee Designated Federal Official:............Christopher D. Miller

This document, the first of the Synthesis and Assessment Products described in the U.S. Climate Change Science 

Program (CCSP) Strategic Plan, was prepared in accordance with Section 515 of the Treasury and General 

Government Appropriations Act for Fiscal Year 2001 (Public Law 106-554) and the information quality act 

guidelines issued by the Department of Commerce and NOAA pursuant to Section 515 <http://www.noaanews.

noaa.gov/stories/iq.htm>). The CCSP Interagency Committee relies on Department of Commerce and NOAA 

certifications regarding compliance with Section 515 and Department guidelines as the basis for determining 

that this product conforms with Section 515. For purposes of compliance with Section 515, this CCSP Synthesis 

and Assessment Product is an “interpreted product” as that term is used in NOAA guidelines and is classified as 

“highly influential”. This document does not express any regulatory policies of the United States or any of its 

agencies, or provide recommendations for regulatory action.

Lisa Dilling to 

provide

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FEDERAL EXECUTIVE TEAM

Director, Climate Change Science Program .............................................William J. Brennan

Director, Climate Change Science Program Office ..................................Peter A. Schultz

Lead Agency Principal Representative to CCSP;

Deputy Under Secretary of Commerce for Oceans and Atmosphere,

National Oceanic and Atmospheric Administration .................................Mary M. Glackin

Product Lead, Earth System Research Laboratory,

National Oceanic and Atmospheric Administration .................................A.R. Ravishankara

Synthesis and Assessment Product Advisory 

Group Chair; Associate Director, EPA National 

Center for Environmental Assessment......................................................Michael W. Slimak

Synthesis and Assessment Product Coordinator, 

Climate Change Science Program Office .................................................Fabien J.G. Laurier

Special Advisor, National Oceanic 

and Atmospheric Administration ..............................................................Chad A. McNutt

EDITORIAL AND PRODUCTION TEAM

Co-Chair....................................................................................................A.R. Ravishankara, NOAA

Co-Chair....................................................................................................Michael J. Kurylo, NASA

Scientific and Technical Editor .................................................................Christine A. Ennis, NOAA/CIRES

Scientific Editor ........................................................................................Jessica Blunden, STG, Inc.

Scientific Editor ........................................................................................Anne M. Waple, STG, Inc.

Scientific Editor ........................................................................................Christian Zamarra, STG, Inc.

Technical Advisor .....................................................................................David J. Dokken, USGCRP

Graphic Design Lead ............................................................................... Sara W. Veasey, NOAA

Graphic Design Co-Lead ......................................................................... Deborah B. Riddle, NOAA

Designer................................................................................................... Brandon Farrar, STG, Inc.

Designer................................................................................................... Glenn M. Hyatt, NOAA

Designer................................................................................................... Deborah Misch, STG, Inc.

Copy Editor.............................................................................................. Anne Markel, STG, Inc.

Copy Editor.............................................................................................. Lesley Morgan, STG, Inc.

Copy Editor.............................................................................................. Susan Osborne, STG, Inc.

Copy Editor.............................................................................................. Susanne Skok, STG, Inc.

Copy Editor.............................................................................................. Mara Sprain, STG, Inc.

Copy Editor.............................................................................................. Brooke Stewart, STG, Inc.

Graphics Support ..................................................................................... Debra A. Dailey-Fisher, NOAA

Technical Support..................................................................................... Jesse Enloe, STG, Inc.

This Synthesis and Assessment Product, described in the U.S. Climate Change Science Program (CCSP) Strategic 

Plan, was prepared in accordance with Section 515 of the Treasury and General Government Appropriations Act 

for Fiscal Year 2001 (Public Law 106-554) and the information quality act guidelines issued by the Department 

of Commerce and NOAA pursuant to Section 515 <http://www.noaanews.noaa.gov/stories/iq.htm>. The CCSP 

Interagency Committee relies on Department of Commerce and NOAA certifications regarding compliance with 

Section 515 and Department guidelines as the basis for determining that this product conforms with Section 515. 

For purposes of compliance with Section 515, this CCSP Synthesis and Assessment Product is an “interpreted 

product” as that term is used in NOAA guidelines and is classified as “highly influential.” This document does 

not express any regulatory policies of the United States or any of its agencies, or provide recommendations for 

regulatory action. 
The First State of the 

Carbon Cycle Report

The North American Carbon Budget 

and Implications for the

Global Carbon Cycle

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, Michael J. Kurylo, and Christine A. Ennis

Trends in Emissions of 

Ozone-Depleting Substances, 

Ozone Layer Recovery, and 

Implications for Ultraviolet 

Radiation Exposure
Transmittal letter to Congress will go here.
TABLE OF CONTENTS

Synopsis.................................................................................................................VII

Preface/Motivation for Report............................................................................ XI

Executive Summary.............................................................................................15

Contents:

1...........................................................................................................................23

Introduction

2...........................................................................................................................29

Current Trends, Mixing Ratios, and Emissions of Ozone-Depleting

Substances and Their Substitutes

Introduction...................................................................................................................................... 33

2.1 Production and Consumption of Ozone-Depleting Chemicals and Their 

Substitutes Derived From Industry Estimates......................................................................... 35

2.1.1 Production and Consumption: Global Trends................................................................... 35

2.1.2 Production and Consumption: Comparing UNEP and AFEAS Compilations................ 38

2.1.3 Production and Consumption of ODSs and Substitutes Not Reported by AFEAS 

 or in UNEP Compilations.................................................................................................... 41

2.1.4 Production and Consumption: United States Trends for ODSs and Substitutes.......... 42

2.1.5 United States Production and Consumption of ODSs and Substitutes Not Included 

 in Published UNEP Compilations...................................................................................... 46

2.2 Emissions: Ozone-Depleting Chemicals and Their Substitutes................................... 46

2.2.1 Global Emissions: Estimates Derived From Atmospheric Observations and Weighted 

 by Ozone Depletion Potentials.......................................................................................... 47

2.2.2 Global Emissions: Estimates Derived From Atmospheric Observations and Weighted 

 by Global Warming Potentials........................................................................................... 48

2.2.3 Global Emissions: The Contribution of Banks and Bank Sizes...................................... 49

2.2.4 Global Emissions: The Influence of Non-restricted Uses and Other Factors .............. 53

2.2.5 United States Emissions and Banks: Estimates Derived by U.S. EPA

 Vintaging Models................................................................................................................. 53

2.2.6 United States Emissions: Derived From Atmospheric Data in Non-remote Areas..... 55

2.3 Changes in the Atmospheric Abundance of Ozone-Depleting Chemicals

 and Their Substitutes.............................................................................................................. 56

2.3.1 Global Atmospheric Abundances....................................................................................... 56

2.3.2 The Untied States Contribution to Global Atmospheric Abundances........................... 61

2.4 The Atmospheric Abundance of Aggregated Chlorine and Bromine From 

 Long-Lived ODSs..................................................................................................................... 63

2.4.1 Atmospheric Chlorine.......................................................................................................... 63

2.4.2 Atmospheric Bromine ......................................................................................................... 65

2.4.3 Equivalent Effective Stratospheric Chlorine and Equivalent Effective Chlorine .......... 67

I
2.5 Changes in Radiative Forcing Arising From Ozone-Depleting Chemicals 

 and Substitutes......................................................................................................................... 68

2.5.1 Changes in Direct Radiative Forcing ................................................................................. 68

2.5.2 Changes in Net Radiative Forcing..................................................................................... 70

2.6 Summary of Findings Related to the Role of the United States in Influencing 

 Past Changes in Production, Consumption, Emissions, and Mixing Ratios of 

 Ozone-Depleting Substances and Their Substitutes........................................................71

Appendix 2A..................................................................................................................................... 73

3...........................................................................................................................79

Ozone and UV Observations

3.1 Introduction ............................................................................................................................... 81

3.2 Ozone.......................................................................................................................................... 84

3.2.1 Total Ozone Observations.................................................................................................. 84

3.2.2 Vertical Distribution of Ozone........................................................................................... 89

3.2.3 Processes That Affect Ozone............................................................................................. 92

3.3 Ultraviolet Radiation at the Earth’s Surface ...................................................................... 97

3.3.1 Background (Factors Controlling UV Surface Irradiance) .............................................. 97

3.3.2 UV in the Polar Regions....................................................................................................106

3.3.3 Human Exposure to UV................................................................................................... 106

3.3.4 UV Summary......................................................................................................................108

Appendix 3A................................................................................................................................... 109

Appendix 3B ................................................................................................................................... 109

4......................................................................................................................... 111

How Do Climate Change and Stratospheric Ozone Loss Interact?

4.1 Introduction ..............................................................................................................................114

4.2 Radiative Forcing of Climate by Ozone-Depleting Substances and 

 Ozone Changes........................................................................................................................115

4.2.1 Direct Radiative Forcing by Ozone-Depleting Substances............................................115

4.2.2 Radiative Forcing From Ozone Changes .........................................................................119

4.3 The Response of Ozone to Climate Change Parameters............................................ 120

4.3.1 Calculating the Response of Ozone to Climate Change Parameters

 with CCMs......................................................................................................................... 121

4.3.2 Stratospheric Temperature Changes.............................................................................. 122

4.3.3 Stratospheric Water Vapor Changes.............................................................................. 125

4.3.4 Changes in Ozone From Increases in Long-Lived Gases in the Stratosphere............ 126

4.4 The Effect of Ozone Changes on Climate Parameters ................................................ 127

4.4.1 Response of Stratospheric and Tropospheric Temperatures to Ozone Depletion .... 128

4.4.2 Response of Surface Temperature to Antarctic Ozone Depletion.............................. 128

4.5 Importance of Volcanoes..................................................................................................... 129

4.5.1 The Effect of Volcanic Aerosol on Ozone....................................................................... 129

4.6 Summary................................................................................................................................... 130

4.6.1 Relevance for the United States................................................................................. 131

TABLE OF CONTENTS

II
5......................................................................................................................... 133

The Future and Recovery

5.1 Introduction ............................................................................................................................. 136

5.2 Model Simulations and Analyses of Ozone Behavior.................................................... 136

5.2.1 Processes and Scenarios Used in Model Simulations ................................................... 137

5.2.2 Results From Model Simulations..................................................................................... 138

5.2.3 Stages of Ozone Recovery From ODSs .......................................................................... 141

5.3 Expected Response in Surface UV..................................................................................... 142

5.4 Future Scenarios for ODSs and their Replacements.................................................... 142

5.4.1 Baseline Scenario............................................................................................................... 142

5.4.2 Alternate Scenarios........................................................................................................... 144

5.4.3 Time Series of Source Gases............................................................................................ 144

5.5 Changes in Integrated EESC and Radiative Forcing....................................................... 145

5.5.1 Time Series of EESC ......................................................................................................... 145

5.5.2 EESC and Mean Age of Air.............................................................................................. 150

5.5.3 Time Series of Radiative Forcing ..................................................................................... 151

5.6 United States Contributions to EESC and to Radiative Forcing

 by ODSs and Their Replacements..................................................................................... 152

5.6.1 Contribution to EESC........................................................................................................ 152

5.6.2 Contribution to Radiative Forcing by ODSs and Their Replacements........................ 153

5.6.3 Options for United States ODS Banks........................................................................... 154

6......................................................................................................................... 155

Implications for the United States

6.1 Introduction............................................................................................................................. 155

6.2 Impacts...................................................................................................................................... 156

6.2.1 Changes in Ozone over the United States..................................................................... 156

6.2.2 Changes in UV over the United States........................................................................... 156

6.2.3 Changes in Radiative Forcing........................................................................................... 157

6.2.4 Future Ozone and UV Changes over the United States............................................... 157

6.2.5 Future Changes in Radiative Forcing .............................................................................. 158

6.3 Accountability ......................................................................................................................... 158

6.3.1 Contribution of the United States to the Global Abundance of ODSs........................ 159

6.3.2 Contribution of the United States to Climate Change via Emission of 

 Ozone-Depleting Substances and the Resulting Ozone Changes .............................. 160

6.4 Potential Management Options.......................................................................................... 161

6.4.1 The World Avoided............................................................................................................ 162

Appendix A..................................................................................................................................... 167

Twenty Questions and Answers About the Ozone Layer: 2006 Update

Glossary and Acronyms............................................................................................................... 215

References....................................................................................................................................... 219

III
AUTHOR TEAM FOR THIS REPORT

Preface Authors: A.R. Ravishankara, NOAA; Michael J. Kurylo, NASA

Executive Summary Convening Lead Authors: A.R. Ravishankara, NOAA; Michael J. Kurylo, NASA

 Lead Authors: Richard Bevilacqua, Naval Research Laboratory; Jeff Cohen, U.S. EPA; 

 John S. Daniel, NOAA; Anne R. Douglass, NASA; David W. Fahey, NOAA; Jay R. 

Herman, NASA; Terry Keating, U.S. EPA; Malcolm Ko, NASA; Stephen A. Montzka, 

 NOAA; Paul A. Newman, NASA; V. Ramaswamy, NOAA; Anne-Marie Schmoltner, 

NSF; Richard Stolarski, NASA; Kenneth Vick, USDA 

Chapter 1 Convening Lead Authors: A.R. Ravishankara, NOAA; Michael J. Kurylo, NASA; 

 Anne-Marie Schmoltner, NSF

Chapter 2 Convening Lead Author: Stephen A. Montzka

 Lead Authors: John S. Daniel, NOAA; Jeff Cohen, U.S. EPA; Kenneth Vick, USDA

Chapter 3 Convening Lead Authors: Paul A. Newman, NASA; Jay R. Herman, NASA

 Lead Authors: Richard Bevilacqua, Naval Research Laboratory; Richard Stolarski, 

NASA; Terry Keating, U.S. EPA

Chapter 4 Convening Lead Author: David W. Fahey, NOAA

 Lead Authors: Anne R. Douglass, NASA; V. Ramaswamy, NOAA; Anne-Marie 

 Schmoltner, NSF

Chapter 5 Convening Lead Author: Malcolm Ko, NASA

 Lead Authors: John S. Daniel, NOAA; Jay R. Herman, NASA; Paul A. Newman, 

 ASA; V. Ramaswamy, NOAA

Chapter 6 Convening Lead Authors: A.R. Ravishankara, NOAA; Michael J. Kurylo, NASA

 Lead Authors: John S. Daniel, NOAA; David W. Fahey, NOAA; Jay R. Herman, NASA; 

 Stephen A. Montzka, NOAA; Malcolm Ko, NASA; Paul A. Newman, NASA; Richard 

 Stolarski, NASA

IV
ACKNOWLEDGEMENT

V

CCSP Synthesis and Assessment Product 2.4 (SAP 2.4) was developed with the benefit of a 

scientifically rigorous, first draft peer review conducted by a committee appointed by the National 

Research Council (NRC). Prior to their delivery to the SAP 2.4 Author Team, the NRC review 

comments, in turn, were reviewed in draft form by a second group of highly qualified experts to 

ensure that the review met NRC standards. The resultant NRC Review Report was instrumental 

in shaping the final version of SAP 2.4, and in improving its completeness, sharpening its focus, 

communicating its conclusions and recommendations, and improving its general readability. 

We thank the members of the NRC Review Committee: M. Joan Alexander (Chair), NorthWest 

Research Associates, Boulder, Colorado; Derek Cunnold, Georgia Institute of Technology, Atlanta; 

Terry Deshler, University of Wyoming, Laramie; Steven Lloyd, The Johns Hopkins Applied Physics 

Laboratory, Laurel, Maryland; Mack McFarland, DuPont Fluoroproducts, Wilmington, Delaware; 

Michelle Santee, Jet Propulsion Laboratory, Pasadena, California; Theodore G. Shepherd, University 

of Toronto, Ontario, Canada; Margaret Tolbert, University of Colorado, Boulder; and Donald 

Wuebbles, University of Illinois, Urbana; and also the NRC Staff members who coordinated the 

process: Chris Elfring, Director, Board on Atmospheric Sciences and Climate; Leah Probst, Study 

Director; and Katherine Weller, Senior Program Assistant.

We also thank the individuals who reviewed the NRC Report in its draft form: James G. Anderson, 

Harvard University, Cambridge, Massachusetts; Greg Bodeker, National Institute of Water and 

Atmospheric Research, Lauder, New Zealand; Mary Anne Carroll, University of Michigan, Ann 

Arbor; Veronika Eyring, Institut für Physik der Atmosphäre, Wessling, Germany; Vitali Fioletov, 

Environment Canada, Downsview, Ontario, Canada; and Ross J. Salawitch, University of Maryland, 

College Park; and also Marvin Geller, The State University of New York, Stony Brook, the overseer 

of the NRC review.

We also thank the NOAA Research Council for coordinating a review conducted in preparation 

for the final clearance of this report. This review provided valuable comments from the following 

internal NOAA reviewers: 

Craig Long (Climate Prediction Center)

Robert Portmann (Earth System Research Laboratory)

Susan Solomon (Earth System Research Laboratory)

The review process for SAP 2.4 also included a public review of the Second Draft, and we thank 

the individuals who participated in this cycle. The Author Team carefully considered all comments 

submitted, and a substantial number resulted in further improvements and clarity of SAP 2.4.
VI

We are grateful to Daniel L. Albritton (NOAA) and Philip L. DeCola (NASA), who led the initial 

discussions that resulted in the inclusion of this topic in the suite of Synthesis and Assessment 

Products of the U.S. Climate Change Science Program. We also thank the following individuals for 

their insightful comments during the initial information-gathering stage: Alkiviadis Bais, Aristotle 

University of Thessaloniki; Donald Blake, University of California at Irvine; Greg Bodeker, 

National Institute of Water and Atmospheric Research, New Zealand; Martin Dameris, Institut für 

Physik der Atmosphäre, Germany; Veronika Eyring, Institut für Physik der Atmosphäre, Germany; 

Vitali Fioletov, Environment Canada; Claire Granier, NOAA/Cooperative Institute for Research in 

Environmental Sciences; Neil Harris, University of Cambridge, United Kingdom; Randy Kawa, 

NASA; Kathy Lantz, NOAA/Cooperative Institute for Research in Environmental Sciences; Rolf 

Müller, Forschungszentrum Jülich, Germany; Ronald Prinn, Massachusetts Institute of Technology; 

Ross Salawitch, University of Maryland; Michelle Santee, Jet Propulsion Laboratory; Guus 

Velders, Netherlands Environmental Assessment Agency; and Ray Weiss, Scripps Institution of 

Oceanography.

We also greatly appreciate the assistance of Krisa Arzayus (NOAA) and Fabien Laurier (CCSP) for 

expert guidance and coordination throughout the process of preparing SAP 2.4, and Chad McNutt 

(NOAA) and Rebecca Feldman (NOAA) for assistance with the approval processes. 

Finally, it should be noted that the respective review bodies were not asked to endorse the final 

version of SAP 2.4, as this was the responsibility of the National Science and Technology Council.
copy editor to

 provide

content for

this page

Depletion of the stratospheric ozone layer by human-produced ozone-depleting substances has been recognized 

as a global environmental issue for more than three decades, and the international effort to address the issue 

via the United Nations Montreal Protocol marked its 20-year anniversary in 2007. Scientific understanding 

underpinned the Protocol at its inception and ever since. As scientific knowledge advanced and evolved, the 

Protocol evolved through amendment and adjustment. Policy-relevant science has documented the rise, and 

now the beginning decline, of the atmospheric abundances of many ozone-depleting substances in response to 

actions taken by the nations of the world. Projections are for a return of ozone-depleting chemicals (compounds 

containing chlorine and bromine) to their “pre-ozone-depletion” (pre-1980) levels by the middle of this century 

for the midlatitudes; the polar regions are expected to follow suit within 20 years after that. Since the 1980s, 

global ozone sustained a depletion of about 5 percent in the midlatitudes of both the Northern Hemisphere 

and Southern Hemisphere, where most of the Earth’s population resides; it is now showing signs of turning 

the corner towards increasing ozone. The large seasonal depletions in the polar regions are likely to continue 

over the next decade but are expected to subside over the next few decades. Ozone-depleting substances 

should have a negligible effect on ozone in all regions beyond 2070, assuming continued compliance with the 

Montreal Protocol.

Large increases in surface ultraviolet (UVB; 280-315 nm) radiation and the associated impacts on human health 

and ecosystems would likely have occurred if atmospheric abundances of ozone-depleting substances had 

continued to grow. Scientific findings regarding the role of ozone-depleting chemicals, projected ozone losses, 

and the potential UV impacts galvanized international decision making in the 1980s. As a result of the worldwide 

adherence to the 1987 Montreal Protocol and its Amendments and Adjustments, the large impacts were 

avoided, and future trends in UVB and UVA (315-400 nm) at the surface are expected to be more influenced 

by factors other than stratospheric ozone depletion (such as changes in clouds, atmospheric fine particles, and 

air quality in the lower atmosphere).

Emissions of ozone-depleting substances by the United States have been significant throughout the history 

of the ozone depletion issue. At the same time, the United States has played a leading role in advancing the 

scientific understanding, leading the international decision making, and leading industry’s actions to reduce 

usage of ozone-depleting substances. Continued future declines in emissions of ozone-depleting substances 

from the United States, along with those from other nations, will play a key role in ensuring the ozone layer’s 

recovery.

Projections of a changing climate have added a new dimension to the issue of the stratospheric ozone layer and 

its recovery, and scientific knowledge is emerging on the interconnections between these two global issues. 

Climate change is expected to alter the timing of the recovery of the ozone layer. Ozone-depleting chemicals 

and ozone depletion are known to influence climate change. The curtailment of the ozone-depleting substances 

not only helped the ozone layer but also very likely lessened the forcing of climate (i.e., how it alters climate). 

Climate change and ozone layer depletion are coupled; this has led to new scientific and decision-making 

challenges. The recovery of the ozone layer will occur in an atmosphere that is different from where we started 

roughly three decades back. Our scientific understanding of the connections between climate change and 

ozone layer depletion is at an early but rapidly advancing stage. That topic will remain a focus for the scientific 

community’s efforts over the next few decades.

SYNOPSIS

VII
RECOMMENDED CITATIONS

For the Report as a whole:

CCSP, 2008: Trends in Emissions of Ozone-Depleting Substances, Ozone Layer Recovery, and Implications for Ultraviolet 

Radiation Exposure. A Report by the U.S. Climate Change Science Program and the Subcommittee on Global Change 

Research. [Ravishankara, A.R., M.J. Kurylo, and C.A. Ennis (eds.)]. Department of Commerce, NOAA’s National Climatic 

Data Center, Asheville, NC, 240 pp.

For the Preface:

Ravishankara, A.R. and M.J. Kurylo, 2008: Preface. In: Trends in Emissions of Ozone-Depleting Substances, Ozone Layer 

Recovery, and Implications for Ultraviolet Radiation Exposure. A Report by the U.S. Climate Change Science Program and 

the Subcommittee on Global Change Research. [Ravishankara, A.R., M.J. Kurylo, and C.A. Ennis (eds.)]. Department of 

Commerce, NOAA’s National Climatic Data Center, Asheville, NC, pp.X1–XIII.

For the Executive Summary:

Ravishankara, A.R., M.J. Kurylo, R. Bevilacqua, J. Cohen, J.S. Daniel, A.R. Douglass, D.W. Fahey, J.R. Herman, T. 

Keating, M. Ko, S.A. Montzka, P.A. Newman, V. Ramaswamy, A-M. Schmoltner, R. Stolarski, and K. Vick, 2008: Executive 

Summary. In: Trends in Emissions of Ozone-Depleting Substances, Ozone Layer Recovery, and Implications for Ultraviolet 

Radiation Exposure. A Report by the U.S. Climate Change Science Program and the Subcommittee on Global Change 

Research. [Ravishankara, A.R., M.J. Kurylo, and C.A. Ennis (eds.)]. Department of Commerce, NOAA’s National Climatic 

Data Center, Asheville, NC, pp. 15–22.

For Chapter 1:

Ravishankara, A.R., M.J. Kurylo, and A-M. Schmoltner, 2008: Introduction. In: Trends in Emissions of Ozone-Depleting 

Substances, Ozone Layer Recovery, and Implications for Ultraviolet Radiation Exposure. A Report by the U.S. Climate 

Change Science Program and the Subcommittee on Global Change Research. [Ravishankara, A.R., M.J. Kurylo, and C.A. 

Ennis (eds.)]. Department of Commerce, NOAA’s National Climatic Data Center, Asheville, NC, pp. 23–28.

For Chapter 2:

Montzka, S.A., J.S. Daniel, J. Cohen, and K. Vick, 2008: 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. A Report by the U.S. Climate Change Science Program and the 

Subcommittee on Global Change Research. [Ravishankara, A.R., M.J. Kurylo, and C.A. Ennis (eds.)]. Department of 

Commerce, NOAA’s National Climatic Data Center, Asheville, NC, pp. 29–78.

For Chapter 3:

Newman, P.A., J.R. Herman, R. Bevilacqua, R. Stolarski, and T. Keating, 2008: Ozone and UV Observations. In: Trends in 

Emissions of Ozone-Depleting Substances, Ozone Layer Recovery, and Implications for Ultraviolet Radiation Exposure. A 

Report by the U.S. Climate Change Science Program and the Subcommittee on Global Change Research. [Ravishankara, 

A.R., M.J. Kurylo, and C.A. Ennis (eds.)]. Department of Commerce, NOAA’s National Climatic Data Center, Asheville, 

NC, pp. 79–110.

For Chapter 4:

Fahey, D.W., A.R. Douglass, V. Ramaswamy, and A-M. Schmoltner, 2008: How Do Climate Change and Stratospheric 

Ozone Loss Interact? In: Trends in Emissions of Ozone-Depleting Substances, Ozone Layer Recovery, and Implications for 

Ultraviolet Radiation Exposure. A Report by the U.S. Climate Change Science Program and the Subcommittee on Global 

Change Research. [Ravishankara, A.R., M.J. Kurylo, and C.A. Ennis (eds.)]. Department of Commerce, NOAA’s National 

Climatic Data Center, Asheville, NC, pp.111–132.

VIII
RECOMMENDED CITATIONS

For Chapter 5:

Ko, M., J.S. Daniel, J.R. Herman, P.A. Newman, and V. Ramaswamy, 2008: The Future and Recovery. In: Trends in 

Emissions of Ozone-Depleting Substances, Ozone Layer Recovery, and Implications for Ultraviolet Radiation Exposure. A 

Report by the U.S. Climate Change Science Program and the Subcommittee on Global Change Research. [Ravishankara, 

A.R., M.J. Kurylo, and C.A. Ennis (eds.)]. Department of Commerce, NOAA’s National Climatic Data Center, Asheville, 

NC, pp. 133–154.

For Chapter 6:

Ravishankara, A.R., M.J. Kurylo, J.S. Daniel, D.W. Fahey, J.R. Herman, S.A. Montzka, M. Ko, P.A. Newman, and R. 

Stolarski, 2008: Implications for the United States. In: Trends in Emissions of Ozone-Depleting Substances, Ozone Layer 

Recovery, and Implications for Ultraviolet Radiation Exposure. A Report by the U.S. Climate Change Science Program and 

the Subcommittee on Global Change Research. [Ravishankara, A.R., M.J. Kurylo, and C.A. Ennis (eds.)]. Department of 

Commerce, NOAA’s National Climatic Data Center, Asheville, NC, pp. 155–166.

For Appendix A:

Fahey, D.W. (Lead Author), Twenty Questions and Answers About the Ozone Layer: 2006 Update, 50 pp., World 

Meteorological Organization, Geneva, Switzerland, 2007. [Reprinted from Scientific Assessment of Ozone Depletion: 2006, 

Global Ozone Research and Monitoring Project—Report no. 50, 572 pp., World Meteorological Organization, Geneva, 

Switzerland, 2007.]

IX
The U.S. Climate Change Science Program Preface

X
Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

XI

A primary objective of the U.S. Climate Change Science Program (CCSP) is to provide the best possible 

scientific information to support public discussion, 

as well as government and private sector decision 

making, on key climate-related issues. To help meet 

this objective, the CCSP has identified an initial set 

of 21 Synthesis and Assessment Products (SAPs) that 

address its highest priority research, observation, and 

decision support needs.

This report, CCSP SAP 2.4, addresses Goal 2 of the 

CCSP Strategic Plan: Improve quantification of the 

forces bringing about changes in the Earth’s climate 

and related systems. The Atmospheric Composition 

chapter of the CCSP Strategic Plan describes a vision to produce a Synthesis and Assessment Product 

(SAP) on “Trends in emissions of ozone-depleting 

substances, ozone layer recovery, and implications 

for ultraviolet radiation (UV) exposure–SAP 2.4.” 

The report provides a synthesis and integration of the 

current knowledge of the stratospheric ozone layer, 

ozone-depleting substances, and ultraviolet radiation 

reaching the Earth’s surface.

P.1 CONTEXT FOR THIS SYNTHESIS 

AND ASSESSMENT PRODUCT

SAP 2.4 contributes to the ongoing and iterative 

international process of producing and refining 

climate-related assessments and decision support 

tools. SAP 2.4 integrates findings from the World 

Meteorological Organization (WMO) / United Nations 

Environment Programme (UNEP) 2006 assessment 

on the ozone layer (Scientific Assessment of Ozone 

Depletion: 2006) and the 2005 Special Report of the 

Intergovernmental Panel on Climate Change (IPCC) 

and the Technology and Economic Assessment Panel 

(TEAP) on Safeguarding the Ozone Layer and the 

Global Climate System – Issues Related to Hydrofluorocarbons and Perfluorocarbons. Both of these 

assessments have been extensively reviewed prior 

to their publication. SAP 2.4 discusses these assessments from both the global perspective and in the specific 

context of the United States of America; this SAP 2.4 gives 

the United States-specific perspective of a global issue for 

decision-makers in the United States. The SAP discusses 

ozone changes over North America, the contributions of 

the United States to ozone-depleting substances, and the 

UV changes due to the ozone layer changes over the North 

American continent. This SAP takes advantage of these 

thoroughly vetted scientific assessments to prepare a product 

that can be used to inform domestic and international decision makers in government and industry, scientists, and the 

public. This SAP was planned and initiated in August 2005, 

before the release of the Fourth Assessment Report (AR4) 

of the Intergovernmental Panel on Climate Change (Climate 

Change 2007: The Physical Science Basis). Therefore, this 

report does not rely on the IPCC AR4; however, some key 

pertinent issues from the IPCC report are used in a few 

instances where updated information was essential. They 

are noted as such in the chapters.

P.2 AUDIENCE AND INTENDED USE

The audience for SAP 2.4 includes scientists, decision 

makers in the public sector (federal, state, and local governments), the private sector (chemical industry, transportation, and agriculture; and climate policy and health-related 

interest groups), the international community, and the 

general public. This broad audience is indicative of the 

diversity of stakeholder groups interested in knowledge of 

the stratospheric ozone layer, ozone-depleting substances, 

and ultraviolet radiation, and of how such knowledge might 

be used to inform decisions. The primary users of SAP 

2.4 are intended to include, but are not limited to, officials 

involved in formulating climate and environmental policy, 

individuals responsible for managing emissions of ozonedepleting substances, and scientists involved in assessing 

and/or advancing the frontier of knowledge. The plan for this 

SAP was presented at the CCSP workshop, “U.S. Climate 

Change Science Program, Climate Sciences in Support of 

Decision Making,” held in Arlington, Virginia, during 14-16 

November 2005.

PREFACE

Report Motivation and Guidance for Using 

this Synthesis/Assessment Product

Authors: A.R. Ravishankara, NOAA; Michael J. Kurylo, NASA
The U.S. Climate Change Science Program Preface

XII

SAP 2.4 is intended to be used:

• as a state-of-the-art assessment of our knowledge of the 

stratospheric ozone layer, ozone-depleting substances, 

and ultraviolet radiation at the surface; 

• to provide the scientific basis for decision support to 

guide management and policy decisions that affect 

the ozone layer and emissions of ozone-depleting 

substances; 

• as a means of informing policymakers and the public 

concerning the general state of our knowledge of the 

stratospheric ozone layer and emissions of ozonedepleting substances with respect to the contributions 

of and impacts on the United States; and 

• to provide scientific information on the ozone layer to 

inform important stakeholder groups. Examples of these 

groups include: the chemical industry that produces 

ozone-depleting substances and substitutes for ozonedepleting substances; agencies in the United States 

and sectors of the United States economy that request 

exemptions from emissions of substances banned by 

the Montreal Protocol and its Amendments; and the 

climate-science community. 

Senior managers and the general public may use the Executive Summary of SAP 2.4 to improve their overall understanding of what is known and unknown about the effects 

of United States emissions on the stratospheric ozone layer 

and ultraviolet radiation at the surface. It will also provide 

an estimate of the impacts of the ozone layer changes on 

the country.

P.3 TOPICS AND CONTENT

The focus of this Report follows the Prospectus guidelines 

developed by the Climate Change Science Program and 

posted on its website at (http://www.climatescience.gov). 

SAP 2.4 addresses key issues related to the stratospheric 

ozone layer, including its changes in the past and expected 

levels in the future. Also, it takes account of the current 

abundances and emissions of ozone-depleting substances. 

Further, it synthesizes the best available information on the 

past and future levels of ultraviolet radiation at the Earth’s 

surface. Lastly, it explores the interactions between climate 

change and stratospheric ozone changes. The discussion of 

these topics is carried out within the context of both the globe 

and the United States to distill a regional assessment from 

the global assessments. More specifically, SAP 2.4:

• Quantifies current information on sources, sinks, and 

abundances of ozone-depleting substances and associated uncertainties. 

• Discusses levels of ozone in various regions of the 

stratosphere, including the polar regions. It pays special 

attention to the Antarctic ozone hole and to ozone above 

the continental United States.

• Provides information on the past, current, and future 

levels of ultraviolet radiation, both generally and for the 

continental United States.

• Provides an assessment of the impact of climate and 

compositional changes on the future of the ozone layer, 

and provides some qualitative discussion of the impacts 

of the ozone layer on climate.

• Describes how these findings relate especially to the 

United States.

• Identifies the gaps in understanding where research is 

critical for future assessments of the ozone layer.

The questions addressed by this report include:

• What is the current state of the stratospheric ozone 

layer?

• What are the recorded changes in the emissions and 

concentrations of ozone-depleting substances?

• What do the observations indicate about the abundances 

and trends of stratospheric ozone?

• What is the trend in the occurrence, depth, duration, 

and extent of the Antarctic ozone hole?

• What is the state of ozone depletion in the Arctic 

region?

• When can one expect recovery of the global ozone layer 

and of the Antarctic ozone hole?

• What are the influences of climate change on the 

recovery of the ozone layer?

• How has surface ultraviolet radiation changed in the 

past and what is expected for the future?

• What are the findings specific to the United States on 

the topics of ozone-depleting substances, stratospheric 

ozone depletion, surface ultraviolet radiation changes, 

and expectations for the future ozone layer? 

• What are the various possible emission scenarios that 

can be considered for any further policy actions on 

emissions of ozone-depleting gases?

P.4 OUTLINE OF THE REPORT

The above questions provide the basis for information 

presented in the six chapters of SAP 2.4. The chapters are 

written in a style consistent with major authoritative international scientific assessments (e.g., IPCC assessments, and 

the reports of the Global Ozone Research and Monitoring 

Project of WMO). However, additional explanatory material 

is included both within the Chapters and as an Appendix to 

aid the diverse readership of this SAP. The Executive Summary, which presents the key findings from the main body 

of the Report, as well as Chapters 1 and 6, are intended to 

be useful especially for those involved with policy-related 

ozone layer issues. Chapter 1 is intended as a background 

“primer” for those less familiar with the topic of stratospheric ozone depletion. Chapters 2 through 5 provide the 

detailed material that supports the findings of the Executive 
Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

XIII

Summary. Though they are written at a more technical level, 

they incorporate material to aid their accessibility to the 

broad readership of this SAP. The chapters of SAP 2.4 are:

• Chapter 1: Introduction

• Chapter 2: Current Trends, Mixing Ratios, and 

Emissions of Ozone-Depleting Substances and Their 

Substitutes

• Chapter 3: Ozone and UV Observations

• Chapter 4: How Do Climate Change and Stratospheric 

Ozone Loss Interact?

• Chapter 5: The Future and Recovery

• Chapter 6: Implications for the United States

For those interested readers who are not specialists on the 

ozone-layer issue, an Appendix gives additional scientific 

background on the topics of this SAP. A glossary and a list 

of acronyms are included at the end of the report.

P.5 THE SYNTHESIS AND ASSESSMENT 

PRODUCT TEAM

The authors for this SAP were chosen based on their expertise and participation in the international assessments from 

which this product derives a great deal of information. The 

SAP 2.4 Author Team and their roles are:

• Dr. A. R. Ravishankara, NOAA, Overall Lead

• Dr. Michael J. Kurylo, NASA, Overall Lead

• Dr. Richard Bevilacqua, NRL/DoD, Scientific

Content

• Dr. Jeff Cohen, USEPA, Scientific Content

• Dr. John Daniel, NOAA, Scientific Content

• Dr. Anne Douglass, NASA, Scientific Content

• Dr. David Fahey, NOAA, Scientific Content

• Dr. Jay Herman, NASA, Scientific Content 

• Dr. Terry Keating, USEPA, Scientific Content

• Dr. Malcolm Ko, NASA, Scientific Content

• Dr. Stephen Montzka, NOAA, Scientific Content

• Dr. Paul Newman, NASA, Scientific Content

• Dr. V. Ramaswamy, NOAA, Scientific Content

• Dr. Anne-Marie Schmoltner, NSF, Scientific Content

• Dr. Richard Stolarski, NASA, Scientific Content

• Dr. Kenneth Vick, USDA, Scientific Content

Those who served as Convening Lead Authors (CLAs) and 

Lead Authors (LAs) are shown at the beginning of each 

chapter. An Editorial Staff managed the assembly, formatting, and preparation of the Report. 
The U.S. Climate Change Science Program Preface

XIV
EXECUTIVE SUMMARY

15

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

Synopsis

Depletion of the stratospheric ozone layer by human-produced ozonedepleting substances has been recognized as a global environmental issue 

for more than three decades, and the international effort to address 

the issue via the United Nations Montreal Protocol marked its 20-year 

anniversary in 2007. Scientific understanding underpinned the Protocol 

at its inception and ever since. As scientific knowledge advanced and 

evolved, the Protocol evolved through amendment and adjustment. 

Policy-relevant science has documented the rise, and now the beginning decline, of the atmospheric abundances of many 

ozone-depleting substances in response to actions taken by the nations of the world. Projections are for a return of 

ozone-depleting chemicals (compounds containing chlorine and bromine) to their “pre-ozone-depletion” (pre-1980) levels 

by the middle of this century for the midlatitudes; the polar regions are expected to follow suit within 20 years after 

that. Since the 1980s, global ozone sustained a depletion of about 5 percent in the midlatitudes of both the Northern 

Hemisphere and Southern Hemisphere, where most of the Earth’s population resides; it is now showing signs of turning 

the corner towards increasing ozone. The large seasonal depletions in the polar regions are likely to continue over the 

next decade but are expected to subside over the next few decades. Ozone-depleting substances should have a negligible 

effect on ozone in all regions beyond 2070, assuming continued compliance with the Montreal Protocol.

Large increases in surface ultraviolet (UVB; 280-315 nm) radiation and the associated impacts on human health and 

ecosystems would likely have occurred if atmospheric abundances of ozone-depleting substances had continued to grow. 

Scientific findings regarding the role of ozone-depleting chemicals, projected ozone losses, and the potential UV impacts 

galvanized international decision making in the 1980s. As a result of the worldwide adherence to the 1987 Montreal 

Protocol and its Amendments and Adjustments, the large impacts were avoided, and future trends in UVB and UVA 

(315-400 nm) at the surface are expected to be more influenced by factors other than stratospheric ozone depletion 

(such as changes in clouds, atmospheric fine particles, and air quality in the lower atmosphere).

Emissions of ozone-depleting substances by the United States have been significant throughout the history of the ozone 

depletion issue. At the same time, the United States has played a leading role in advancing the scientific understanding, 

leading the international decision making, and leading industry’s actions to reduce usage of ozone-depleting substances. 

Continued future declines in emissions of ozone-depleting substances from the United States, along with those from 

other nations, will play a key role in ensuring the ozone layer’s recovery.

Projections of a changing climate have added a new dimension to the issue of the stratospheric ozone layer and its 

recovery, and scientific knowledge is emerging on the interconnections between these two global issues. Climate change 

is expected to alter the timing of the recovery of the ozone layer. Ozone-depleting chemicals and ozone depletion are 

known to influence climate change. The curtailment of the ozone-depleting substances not only helped the ozone layer 

but also very likely lessened the forcing of climate (i.e., how it alters climate). 

Climate change and ozone layer depletion are coupled; this has led to new scientific and decision-making challenges. 

The recovery of the ozone layer will occur in an atmosphere that is different from where we started roughly three 

decades back. Our scientific understanding of the connections between climate change and ozone layer depletion is at 

an early but rapidly advancing stage. That topic will remain a focus for the scientific community’s efforts over the next 

few decades.

Convening Lead Authors: A.R. Ravishankara, NOAA; Michael J. 

 Kurylo, NASA

Lead Authors: Richard Bevilacqua, NRL; Jeff Cohen, U.S. EPA; John 

 S. Daniel, NOAA; Anne R. Douglass, NASA; David W. Fahey, 

 NOAA; Jay R. Herman, NASA; Terry Keating, U.S. EPA; Malcolm 

 Ko, NASA; Stephen A. Montzka, NOAA; Paul A. Newman, NASA; 

 V. Ramaswamy, NOAA; Anne-Marie Schmoltner, NSF; Richard 

 Stolarski, NASA; Kenneth Vick, USDA 
The U.S. Climate Change Science Program Executive Summary

16

Research in the 1970s and early 1980s showed 

that the ozone-depleting substances (ODSs), 

mainly chlorofluorocarbons (CFCs) and certain 

compounds containing bromine, would deplete 

stratospheric ozone. The discovery of the 

springtime Antarctic ozone hole in 1985 showed 

that ozone depletion was real and occurring at 

that time, and was not just a prediction for the 

future. 

Faced with the scientific consensus that ozone 

depletion was real and due to human-produced 

ozone-depleting substances, nations throughout 

the world agreed to the Montreal Protocol and 

its subsequent Amendments and Adjustments. 

The United States is a signatory to this protocol. 

The Protocol and its Amendments were 

successfully implemented starting in the late 

1980s. Thus, this Protocol was one of the first 

international agreements to address a global 

environmental problem. The Montreal Protocol 

has had clear benefits in reducing ozonedepleting substances, placing the ozone layer 

on a path to recovery, and protecting human 

health (Figure ES.1). 

Ozone layer depletion, like climate change, 

is a global issue with regional impacts. The 

depletion of the ozone layer is caused by the 

collective emissions of human-produced ozonedepleting substances at Earth’s surface from 

various regions and countries. These ozonedepleting substances persist long enough in 

the atmosphere to be quite well mixed in the 

lower atmosphere and then be transported to 

the stratosphere, where their interaction with 

the harsh UV radiation releases chlorine and 

bromine. Thus, they pose a global threat, 

regardless of where on Earth’s surface they 

are emitted. Emissions of ozone-depleting 

substances arise from their use as coolants, 

fire-extinguishing chemicals, electronics 

cleaning agents, and in foam blowing and other 

applications. The contributions to the global 

atmospheric burden of these ozone-depleting 

substances vary by regions and countries. There 

are large variations in the extent and timing 

of ozone depletion in various regions, and 

the impacts are also different. Consequently, 

the impacts of ozone layer depletion can be 

different in different regions of the world.

ES.1 What is Ozone Layer

Depletion and Why Is it a

Concern?

The stratospheric ozone layer lies in a region 

of the atmosphere approximately 15 to 45 

kilometers (roughly 9 to 28 miles) above Earth’s 

surface. The ozone layer acts as a protective 

shield, preventing most of the Sun’s harmful 

ultraviolet (UV) radiation from reaching the 

surface. The depletion of the ozone layer can 

therefore lead to enhancements of the UV 

radiation that reaches Earth’s surface, with 

consequences for human health, the Earth’s 

ecosystems, and physical materials. The 

ozone layer and its changes can also alter 

the atmosphere’s temperature structure and 

weather/climate-related circulation patterns.

Figure ES.1 Effect of the Montreal Protocol. The top 

panel gives a measure of the projected future abundance 

of ozone-depleting substances in the stratosphere, without 

and with the Protocol and its various Amendments. The 

bottom panel shows similar projections for how excess skin 

cancer cases might have increased (adapted from Appendix 

A of this Report).

The depletion of the 

ozone layer can lead 

to enhancements 

of ultraviolet (UV) 

radiation that 

reaches Earth’s 

surface, with 

consequences for 

human health, 

the Earth’s 

ecosystems, and 

physical materials.
17

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

The f indings from this Synthesis and 

Assessment Product are summarized in three 

parts. Section ES.2 of this Executive Summary 

lists the findings to inform the public in 

general nontechnical terms, and Section ES.3 

summarizes findings for those involved in 

potential policy formulation. The Executive 

Summary findings are backed up by a more 

technical set of findings, primarily for scientists 

and secondarily for those who want to delve more 

into the details. These technical findings are 

listed near the beginning of Chapters 2 through 

5, and in Chapter 6 on Policy Implications for 

the United States. Appendix A of this Synthesis 

and Assessment Product provides extensive 

background material on the science regarding 

the ozone layer, ozone-depleting substances, 

surface ultraviolet radiation, and connections 

to climate change.

ES.2 Key Findings about

the ozone layer, surface

UV, ozone-depleting 

substances, and

connections to climate

change

ES.2.1 The Ozone Layer, OzoneDepleting Substances, and Climate 

Change: What Are the Connections?

Ozone layer changes caused by ozone-depleting 

substances are intertwined with the issue of 

climate change, even though the two issues have 

been distinct in most policy formulations. 

Over the course of the past 20 years, the close 

connections between stratospheric ozone 

depletion and climate change issues have 

become clearer (Figure ES.2). 

• Ozone-depleting substances and many 

of the chemicals being used to replace 

them are potent greenhouse gases that 

influence the Earth’s climate by trapping 

terrestrial infrared (heat) radiation that 

would otherwise escape to space. 

• Ozone is itself a greenhouse gas. The 

stratospheric ozone layer heats the 

stratosphere and, indirectly, the lower 

at mos phe r e (t rop os phe r e). T hu s , 

stratospheric ozone is a key component 

that affects climate. Depletion of the ozone 

layer has a cooling effect on climate, 

though large uncertainties exist regarding 

this effect, which is a combination of 

multiple contributing factors. 

• The recovery of the ozone layer is 

inf luenced not only by the decreases 

in ozone-depleting substances required 

by the Montreal Protocol, but also by 

changes to climate and Earth’s atmospheric 

composition.

Ozone-depleting substances are continuing to 

make a significant contribution to global climate 

change, but in the future ODSs are expected to 

make a smaller and smaller contribution. The 

direct ODS contribution to global climate 

change between 1750 (pre-industrial times) 

and 2005, as measured by a quantity called 

radiative forcing that is a metric for the ability 

to force climate change, is approximately 20% 

of that from carbon dioxide (CO2), the largest 

human-caused contributor to global radiative 

forcing (Figure ES.3). The combined radiative 

forcing from ODSs and substitutes including 

hydrofluorocarbons (HFCs) is still increasing, 

but at a much slower rate than in the 1980s. The 

total contribution of human-produced ODSs 

and substitutes in 2005 was about 15% of the 

contribution from the major greenhouse gases 

(CO2, methane [CH4], nitrous oxide [N2O]). 

The ODS contribution is expected to decline in 

coming decades as ODS emissions decline and 

CO2 emissions continue to rise. 

The recovery of 

the ozone layer is 

influenced not only 

by the decreases 

in ozone-depleting 

substances required 

by the Montreal 

Protocol, but 

also by changes 

to climate and 

Earth’s atmospheric 

composition.

Figure ES.2 Simplified schematic of some of the processes that interconnect the issues of ozone layer depletion and climate change (adapted from 

Chapter 4 of this Report).
The U.S. Climate Change Science Program Executive Summary

18

Depletion of stratospheric ozone since about 

1980 is estimated to have caused a slight 

negative (cooling) radiative forcing of climate 

(approximately –0.05 Watts per meter squared 

[W per m2] with a range of −0.15 to +0.05 W per 

m2) (Figure ES.3). While this forcing is likely to 

be a cooling term (i.e., in the opposite direction 

to climate forcing by the ODSs that caused the 

depletion) it has large uncertainties. Globally 

averaged, it may even represent a warming 

within the error bars, or it could offset a large 

portion (up to 44%) of the ODS warming, 

while the current best estimate is an offset of 

approximately 15%. This estimate is based on 

observed ozone changes and assumes that they 

are due entirely to ODSs. Recent research has 

shown that ozone cooling and ODS warming 

often occur in different places and times, 

making it less appropriate to consider the two 

terms as offsetting one another than previously 

thought.

Climate change will lead to either increases 

or decreases in ozone abundances depending 

on the location in the atmosphere and the 

magnitude of climate change. While the 

surface temperature has increased, observed 

stratospheric temperature decreased starting 

in the 1960s and it is expected to continue 

to decrease. The global average trend is 

attributed mainly to ozone depletion, increased 

CO2, and changes in water vapor. Dynamical 

changes are also likely to be important for local 

temperature changes, but are not significant 

for global mean stratospheric temperature 

trends. Stratospheric temperatures influence 

ozone amounts through chemical and transport 

processes. Stratospheric water vapor influences 

stratospheric ozone through chemistry, 

formation of polar stratospheric clouds, and 

changes in temperature.

ES.2.2 Ozone-Depleting Substances: 

Past, Present, and Future

The Montreal Protocol has been effective in 

reducing the use of ozone-depleting substances. 

Assuming continued compliance with the 

Protocol, the atmospheric abundance of ODSs 

is expected to decline back to its pre-1980 level 

by the middle of this century.

Total global production and consumption of 

ODSs have declined substantially since the late 

1980s in response to the Montreal Protocol. 

By 2005, the annual aggregated production 

and consumption magnitudes of the ODSs, 

after accounting for their differences in ozone 

depletion capabilities, had declined 95 percent 

from peak amounts produced and consumed in 

the late 1980s.

In response to these global production and 

consumption changes, global ODS emissions 

have declined. Hence, the total amount of 

ODSs in the atmosphere, as measured by their 

combined ability to deplete the ozone layer, is 

now decreasing both in the troposphere and 

stratosphere.

In this Report, future halocarbon emissions 

are derived using a new bottom-up approach 

for estimating emissions from the sizes of the 

banks (ODSs produced but not yet released). 

The new method gives future CFC emissions 

that are higher than previously estimated in 

WMO (2003). There are still some uncertainties 

in the future abundances of ODSs.

Total global 

production and 

consumption of 

ozone-depleting 

substances (ODSs) 

have declined 

substantially since 

the late 1980s in 

response to the 

Montreal Protocol. 

Hence, the total 

amount of ODSs in 

the atmosphere is 

now decreasing both 

in the troposphere 

and stratosphere.

Figure ES.3 Radiative forcing values for the principal contributions 

to climate change from atmospheric gas changes since preindustrial 

times, including halogen-containing gases such as ODSs, and the 

cooling caused by depletion of stratospheric ozone. These climate 

influences are expressed as radiative forcings, a metric for the ability 

to force climate change (adapted from IPCC, 2007).
19

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

The effective sum of chlorine and bromine in 

the stratosphere, with bromine weighted by its 

larger per-atom efficiency in depleting ozone, is 

estimated to recover to the 1980 value between 

2040 and 2050 in the midlatitudes (Figure 

ES.4) and between 2060 and 2070 in the polar 

regions. 

ES.2.3 Ozone in the Stratosphere: 

Past, Present, and Future

Total global ozone, as well as seasonal 

springtime ozone in both southern and northern 

polar regions, exhibited declines since the 

early 1980s, but recent observations show 

that ozone depletion is not worsening and in 

some atmospheric regions is showing signs 

that recovery has started. Ozone in the future 

is projected to recover as the atmospheric 

amounts of ODSs decline over the next few 

decades (with recovery above midlatitudes and 

the Arctic preceding Antarctic recovery). With 

continued adherence to the Montreal Protocol, 

ozone-depleting substances identified in the 

Protocol should have a negligible effect on 

ozone in all regions beyond 2070.

Total global ozone declined by roughly 5 

percent since the early 1980s but has remained 

relatively constant over the last four years 

(2002 to 2006). Northern midlatitude ozone 

reached a minimum in 1993, and has increased 

somewhat since then. The 1993 minimum 

largely resulted from the increase of particles 

in the stratosphere caused by the eruption of 

Mt. Pinatubo. Southern midlatitude ozone 

decreased until the late 1990s, and has been 

constant since. There are no significant total 

ozone trends over the tropics.

Ozone depletion in the upper stratosphere, where 

the influence of chlorine is easiest to detect, has 

slowed and has closely followed the trends in 

the sum of total chlorine. Although bromine 

plays a lesser role than chlorine in controlling 

ozone in the upper stratosphere, it too shows 

signs of leveling off in the stratosphere (see 

Section 2.4.2).

Antarctic ozone depletion can be measured 

in different ways, such as the total amount of 

ozone lost (called mass deficit), the minimum 

values of ozone observed, and the geographical 

area of the ozone hole. Over the last decade 

(1995 to 2006), the Antarctic ozone depletion by 

all these measures has not worsened. The ozone 

hole area and ozone mass deficit were observed 

to be below average in some recent winter years 

while higher minimum column amounts have 

also been recorded. This variability results 

from the strong influence of meteorological 

variability on ozone amounts, and not from any 

changes in the amounts of chlorine and bromine 

available for ozone depletion. Declines in the 

amounts of chlorine and bromine available for 

ozone depletion are likely quite small in this 

region.

Arctic spring total ozone values over the last 

decade were lower than values observed in 

the 1980s. In addition, spring Arctic ozone is 

highly variable depending on meteorological 

conditions. For current halogen levels, humancaused chemical loss and variability in ozone 

transport are about equally important for 

year-to-year Arctic ozone variability. Colderthan-average vortex conditions result in larger 

halogen-driven chemical ozone losses.

If volcanic eruptions that inject material into 

the stratosphere were to occur in the coming 

Recent observations 

show that ozone 

depletion is not 

worsening and in 

some atmospheric 

regions is showing 

signs that recovery 

has started.

Figure ES.4 Estimates (presented in parts per trillion, [ppt]) of the effective sum 

of ozone-depleting chlorine and bromine in the stratosphere (called Equivalent 

Effective Stratospheric Chlorine, [EESC]), a metric that accounts for the differences in ozone depletion capabilities of chlorine and bromine. Estimates in the 

past are based upon observations, and estimates in the future are based upon a 

baseline scenario and three comparative test cases. The horizontal line represents 

the 1980 (“pre-ozone-depletion”) level of EESC (adapted from WMO, 2007).
The U.S. Climate Change Science Program Executive Summary

20

decades, they are expected to cause major 

temperature and circulation changes in the 

stratosphere as have occurred after past 

eruptions. The changes are caused by the 

large increases in fine particles formed from 

sulfur dioxide injected into the stratosphere 

following such eruptions. The increases result 

in a transient shift in stratospheric ozone levels 

and climate because natural processes gradually 

remove the additional sulfate particles after the 

eruption.

Assuming an absence of volcanic injections 

into the stratosphere, and based on the projected 

changes in ozone-depleting substances and 

changes in the major climate-relevant trace 

gases, modeling calculations predict the 

following for the future of the ozone layer 

(Figure ES.5). 

• The ozone content between 60°N and 60°S, 

between now and 2020, will increase in 

response to decreases in halogen loading.

• Global ozone is expected to return to its 

1980 value up to 15 years earlier than 

the halogen recovery date because of 

stratospheric cooling and changes in 

circulation associated with greenhouse 

gas emissions. 

• Global ozone abundances (from 60°N to 

60°S) are expected to be 2 percent above 

the 1980 values by 2100 for the assumed 

scenario for greenhouse gases noted in this 

report. Values at midlatitudes could be as 

much as 5 percent higher.

The minimum ozone value for Antarctic ozone 

is projected to start increasing after 2010 in 

several model calculations, while another 

measure of ozone depletion (the ozone mass 

deficit, the total amount lost in a season) begins 

decreasing around 2005 in most models. 

• Model simulations show that the ozone 

amount in the Antarctic will reach the 1980 

values 10 to 20 years earlier than the 2060 

to 2070 time frame of when the ODSs reach 

their 1980 levels in polar regions.

• Ozone in the Arctic region is expected to 

increase as ODSs decline in the atmosphere. 

Because of large interannual variability, 

the simulated results do not show a smooth 

monotonic recovery of Arctic ozone. The 

dates of the minimum ozone from different 

models occur between 1997 and 2015. 

• Most climate chemistry models show 

Arctic ozone values by 2050 larger than 

the 1980 values, with the recovery date 

between 2020 and 2040.

The above projections are based on currently 

available models. As our scientific understanding 

and modeling capabilities continue to evolve, 

our best predictions of the timing and extent of 

ozone layer recovery will also evolve.

ES.2.4 Surface Ultraviolet Radiation: 

Past, Present, and Future

The Montreal Protocol and its Amendments 

have prevented large increases in global 

surface UVB radiation. As the stratospheric 

ozone layer recovers over the next few decades, 

factors such as changes in clouds, atmospheric 

fine particles, and air quality in the lower 

atmosphere will be the dominant factors 

influencing future UV changes.

Surface UVB changes resulting from ozone 

depletion over Antarctica in early austral spring 

have been very large. Changes in the surface 

The Montreal 

Protocol and its 

Amendments have 

prevented large 

increases in global 

surface ultraviolet 

radiation.

Figure ES.5 Global ozone recovery predictions (derived 

from Fahey, 2007).
21

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

UVB due to ozone depletion in most other 

locations of the world have not been clearly 

discernable, because the effects have been much 

smaller compared with changes due to other 

factors. For example, trends in UV exposure 

changes at ground level in the midlatitude 

United States attributable to ozone changes 

are difficult to discern from ground-based 

observations, since the observations are also 

dependent on changes in clouds and pollution 

from suspended fine particles in the air. What 

is clear is that in the absence of the Montreal 

Protocol, ozone depletion would have caused 

increases in surface UV by 2010 over most of 

the world, to such an extent that other factors 

(e.g., clouds, atmospheric fine particles, air 

quality) would have been of relatively minor 

importance.

Possible future UV trends at the surface are 

likely to be influenced more by changes in 

clouds, atmospheric fine particles, and lower 

atmosphere air quality than by ozone layer 

depletion. 

ES.3 Implications for the

United States: Impacts, 

Accountability, and

Potential Management

Options

It is not possible to make a simple connection 

between emissions of ozone-depleting 

substances from the United States and the 

depletion of ozone above the country. This 

is because ODSs persist long enough in the 

atmosphere to be quite well mixed in the global 

lower atmosphere, before transport to the 

stratosphere occurs. Thus, ODSs pose a global 

threat, regardless of where on Earth’s surface 

they are emitted. However, the depletion of 

stratospheric ozone over the various regions 

of the United States, and the contribution of 

emissions from the United States to the global 

burden of ozone-depleting substances, can be 

quantified. 

Impacts: Changes in Ozone and 

Surface Ultraviolet Radiation 

over the United States

Ozone depletion above the continental United 

States (i.e., the midlatitudes) has essentially 

followed the depletion occurring over the 

northern midlatitude regions: a decrease to a 

minimum around the mid-1990s and a slight 

increase since that time. The minimum total 

column ozone amounts over the continental 

United States, reached in 1993, were about 5 

to 8 percent below the amounts present prior 

to 1980. The ozone increase since 1993 has 

diminished the ozone deficit to about 2 to 5 

percent below the pre-1980 amounts. These 

midlatitude ozone changes are estimated to 

contain a significant contribution from the 

ozone depletion that occurs in the Arctic during 

springtime. 

Ozone over Northern high latitudes, such as over 

northern Alaska, is most influenced by Arctic 

springtime total ozone values, which in recent 

years have been lower than those observed in 

the 1980s. The springtime ozone depletions are 

highly variable from year to year. 

Calculations based on satellite observations of 

column ozone and surface reflectivity suggest 

that the averaged erythemal irradiance (which 

is a weighted combination of UVA and UVB 

based on skin sensitivity) over the United States 

had increased roughly by about 7 percent at the 

time when the ozone minimum was reached in 

1993 and is now about 4 percent higher than 

in 1979. Direct surface-based observations do 

not show significant trends in UV levels over 

the United States over the past three decades 

because effects of clouds and atmospheric fine 

particles have likely masked the increase in UV 

due to ozone depletion over this region.

Accountability: United States 

Contributions to OzoneDepleting Substances 

The contributions of the United States to the 

emission of ODSs to date have been significant. 

For example, in terms of dispersive uses of 

ODSs regulated and restricted by the Montreal 

Protocol, emissions from the United States 

accounted for between 15 and 39 percent of 

the overall atmospheric abundance of ODSs 

measured between 1994 and 2004. The United 

States has also contributed significantly to 

emission reductions of ODSs, thereby helping 

efforts to achieve the expected recovery of 

the ozone layer and prevent large surface UV 

changes. 

Emissions from 

the United States 

accounted for 

between 15 and 

39% of the overall 

atmospheric 

abundance of ODSs 

measured between 

1994 and 2004. The 

United States has 

also contributed 

significantly to 

emission reductions 

of ODSs, thereby 

helping efforts to 

achieve the expected 

recovery of the 

ozone layer.
The U.S. Climate Change Science Program Executive Summary

22

Future Options

United States emissions of ODSs in the future, 

like those from other developed nations, will be 

determined largely by the size of ODS “banks,” 

i.e., those ODSs that are already produced 

but not yet released to the atmosphere. While 

global ODS banks are estimated to have been 

2960 ODP-kilotons (Kt) in 2005, ODS banks 

in the United States then were 830 ODP-Kt. 

Of this U.S. bank, approximately 210 ODPKt has been classified as accessible by the 

U.S. Environmental Protection Agency. The 

expected future declining emissions of ODSs 

from the United States and throughout the globe 

will also aid in reducing the climate forcing from 

these substances. While global banks amounted 

to between 5 and 24 Gt CO2-equivalents, the 

accessible bank of hydrochlorofluorocarbons 

(HCFCs) in the United States, for example, 

amounted to between 0.9 and 1.1 Gt CO2-

equivalents.

While the Montreal Protocol has had a large 

beneficial effect on current and projected 

ozone depletion, options remain for the United 

States, and other countries as well, to reduce 

ozone depletion arising from ozone-depleting 

substances over the coming decades. The 

greatest reduction possible would be obtained 

from the hypothetical cessation of all future 

emissions of ozone-depleting substances 

(including emissions from banks and future 

production). If such a cessation had been 

implemented globally in 2007, the anticipated 

return of the ozone-depleting substances to 

their 1980 level would be advanced by about 

15 years.

Methyl bromide is a potent ODS that has 

significant quarantine and pre-shipment (QPS) 

uses that are not restricted by the Montreal 

Protocol, and Critical Use Exemptions (CUEs) 

that are currently large compared to restricted 

uses. The importance of human-emitted 

methyl bromide to future ozone depletion will 

depend on the future magnitude of emissions 

from these unrestricted uses and from CUEs. 

Reducing these emissions would benefit the 

ozone layer.

The World Avoided

Without the Montreal Protocol regulations, the 

levels of ODSs around 2010 likely would have 

been more than 50 percent larger than currently 

predicted (Figure ES.1). The abundances in 

the remaining twenty-first century would have 

depended on the specific actions taken by 

humankind. The increases in ODSs would have 

caused a corresponding substantially greater 

global ozone depletion. The Antarctic ozone 

hole would have persisted longer and may have 

been even larger than what has been observed 

to date.

Without the 

Montreal Protocol 

regulations, the levels 

of ODSs around 2010 

likely would have 

been more than 50% 

larger than currently 

predicted. The 

increases in ODSs 

would have caused 

a corresponding 

substantially 

greater global 

ozone depletion.
Introduction

CHAPTER

1

23

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

Convening Lead Authors: A.R. Ravishankara, NOAA; 

Michael J. Kurylo, NASA; Anne-Marie Schmoltner, NSF

Ozone (O3) is the triatomic form of oxygen. It is a key atmospheric trace gas that is present 

everywhere in the atmosphere and is most abundant in the stratosphere. The abundance of 

ozone in the stratosphere is largest in the region that is roughly between 15 and 35 kilometers 

(km) height above the Earth’s surface, which is referred to as the stratospheric ozone layer. This 

stratospheric ozone layer (Box 1.1) plays many important roles in the Earth system: 

• It protects the lower part of the atmosphere (the troposphere) and the Earth’s surface 

from damaging, or “harsh” ultraviolet1 (UV) radiation from the sun; 

• It influences the chemical composition of the lower atmosphere by altering the amount and 

type (wavelength distribution) of solar radiation passing through it; 

• It changes the temperature structure of the stratosphere and thus influences atmospheric 

transport and mixing; and 

• It contributes ozone to the upper troposphere, where ozone is an important greenhouse 

gas.

Because of many of the above contributions, ozone in the stratosphere and its changes also 

play a significant role in the Earth’s climate system; changes in the ozone layer are influenced 

by climate change and also contribute to climate change. Appendix A of this Product contains 

background information and answers to some of the most frequently asked questions about 

the stratospheric ozone layer (Fahey, 2007).

The focus of this Product is on key issues related to: (1) the stratospheric ozone layer, including 

its changes in the past, its current abundances, and expected levels in the future; (2) emissions of 

ozone-depleting substances (ODSs) and their influences on the ozone layer and climate; and (3) 

the changes in the ground-level UV radiation associated with stratospheric ozone changes.

1 ‘Harsh’ UV radiation indicates the higher energy portion of the UV spectrum.

The potential for 

human-produced 

chemicals, such as 

chlorofluorocarbons 

(CFCs), to deplete 

the stratospheric 

ozone layer has 

received a great deal 

of attention since 

the early 1970s. 
The U.S. Climate Change Science Program Chapter 1

24

The chemical processes that lead to the formation of ozone, as well as those that remove or 

destroy it, are distinctly different in the stratosphere from those in the troposphere (Box 1.2). 

The ever-present balance in the stratosphere 

between production, removal, and transport 

determines the abundance of ozone in any given 

part of the ozone layer. The majority of the 

removal processes in the stratosphere involve 

catalytic cycles in which ozone-destroying 

chemicals are re-formed after destroying ozone. 

This catalytic capability is a key reason why 

very small amounts of ozone-destroying chemicals introduced into the atmosphere can vastly 

influence the ozone layer (Box 1.2).

The potential for human-produced chemicals, 

such as chlorofluorocarbons (CFCs), to deplete 

the stratospheric ozone layer has received a 

great deal of attention since the early 1970s. 

The depletion by chlorine released from CFCs 

in the stratosphere was expected to be catalytic 

in nature, meaning that small amounts of CFCs 

could destroy vast amounts of ozone. The ozone 

depletion was predicted to lead to changes in 

UV radiation at the Earth’s surface, with potentially major environmental consequences. The 

anticipated effects of increased UV radiation 

included: increased incidence of skin cancer 

and cataracts in humans; detrimental effects 

on ecosystems including the aquatic system; 

and harmful effects on materials, such as rubber and plastics. These potential effects were 

debated and the nations of the world agreed 

to protect the ozone layer through the 1985 

Vienna Convention. Then the ozone hole in 

Antarctica was discovered in 1985. Investigation of the causes of this annually recurring 

The nations of the 

world agreed to 

protect the ozone 

layer through 

the 1985 Vienna 

Convention. That 

same year the ozone 

hole in Antarctica 

was discovered.

About 90% of the atmospheric ozone resides in 

the stratosphere, in a region between roughly 

15 and 35 km above the Earth’s surface, as 

indicated by the red line in Box Figure 1.1. 

This region is referred to as the stratospheric 

ozone layer. The remainder of the atmospheric 

ozone resides in the troposphere, the lower 

layer of the atmosphere. Stratospheric ozone 

is formed and destroyed by chemical reactions, 

as shown in Box 1.2. Of particular note are the 

need for higher-energy UV radiation for the 

formation of ozone and the catalytic nature 

of the ozone removal processes. The ozone 

layer in turn shields the lower part of the 

atmosphere and the surface from damaging 

UV radiation because ozone itself absorbs UV 

radiation. Depletion of the ozone layer allows 

more UVB radiation (wavelength 280 to 315 

nanometers) to reach the Earth’s surface. 

This radiation is harmful to humans and many 

other biological systems and causes damage to 

materials. The ozone in the lower atmosphere, 

the troposphere, is formed by methods 

different from those in the stratosphere, as 

shown in Box 1.2. Further, the contribution of this lower atmospheric ozone to the total in the atmosphere is 

small, of the order of a few percent in the Southern Hemisphere to about 10% in the Northern Hemisphere. The 

ozone in the lower atmosphere is harmful because, in direct contact, ozone is toxic to biological systems and can 

deteriorate many materials. It can cause respiratory and other health problems for humans. In addition, ozone 

and its changes in both the stratosphere and the lower atmosphere are important greenhouse gases and thus 

their changes influence climate. See Appendix A of this Synthesis and Assessment Product for further background 

information about ozone.

BOX 1.1: The Stratospheric Ozone Layer and Its Role in the Atmosphere

Box Figure 1.1 This figure shows the distribution of ozone in the 

atmosphere (adapted from Fahey, 2007; see Appendix A of this Report).
25

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

polar springtime ozone depletion indicated 

that CFCs and other ozone-depleting chemicals 

were involved in additional catalytic ozone 

destruction pathways unique to the extremely 

cold polar stratosphere. It was also discovered 

that small particles containing water and nitric 

and/or sulfuric acid that are found in polar 

stratospheric clouds (PSCs) play a crucial role 

in these processes by converting chemically 

less reactive halogen-containing chemicals 

into more reactive chemicals, which are more 

effective in ozone depletion, and involved some 

catalytic cycles unique to this region.

The Montreal Protocol, a sequel to the Vienna 

Convention, was agreed to in 1987 in the setting of the scientific knowledge at that time. 

First, the agreements of the Protocol were to 

reduce CFC emissions by replacing CFCs with 

less harmful substances, if 

possible, that could be used 

in existing devices for most 

applications. A few applications utilized not-in-kind (alternative) processes that did 

not require ozone-depleting 

chemicals. Many of the replacement chemicals, such 

as hydrochlorofluorocarbons (HCFCs), still contained 

chlorine but overall were less 

harmful to the stratospheric 

ozone layer than CFCs. Eventually, even the chlorine-containing substitutes were to be 

replaced by non-chlorine or 

bromine containing replacements such as hydrofluorocarbons (HFCs).

Over the past three decades 

of ozone-layer research, it has 

become clearer that ODSs, as 

well as many of the CFC-substitutes introduced to comply 

with the Montreal Protocol, 

are also potent greenhouse 

gases. Ozone depletion and 

climate change are distinct 

issues but are inextricably 

linked because ozone itself 

is a greenhouse gas and many 

of the ozone-depleting gases 

are potent greenhouse gases. To add to the 

complexity, changes in the major greenhouse 

gases such as carbon dioxide (CO2), methane 

(CH4), and nitrous oxide (N2O) also influence 

ozone depletion. Increases in CO2 lead to a 

cooling of the stratosphere, which increases 

ozone in the upper stratosphere in non-polar 

regions, but decreases ozone in the polar lower 

stratosphere. The influence of CH4 and N2O 

on the stratospheric ozone layer is dominated 

by their chemical interactions. Figure 1.1 captures this influence in a schematic form. An 

assessment of the climate effects of ODSs must 

consider both of their roles: as chemicals that 

deplete ozone, and as greenhouse gases that 

alter climate.

Figure 1.1 The two faces of ozone-depleting substances: their roles as depleting agents of 

stratospheric ozone, and as greenhouse gases that influence climate. The two roles are further 

interconnected because ozone itself is a greenhouse gas and because climate change can lead to 

changes in the ozone layer. The various connections between these two phenomena are shown. A 

plot of the changes in the observed global ozone illustrates the stratospheric ozone depletion issue. 

The radiative forcing due to various greenhouse gases, including ODSs, depicts the greenhouse 

gas issue and stratospheric ozone changes.

Because many ozonedepleting substances 

have lifetimes of 

many years in 

the atmosphere, 

the depletion of 

stratospheric ozone 

is a global problem. 

The observed ozone 

depletion above a 

given region will not 

be directly related 

to the emissions 

from that region.
The U.S. Climate Change Science Program Chapter 1

26

Since 1987, there have been many amendments 

and adjustments to the Montreal Protocol to accelerate efforts to curtail the emissions of ODSs. 

These actions have come about in response to 

our evolving knowledge of the ozone layer 

and its changes, and have led to a reduction in 

the emissions and, subsequently, in the atmospheric abundances of most ozone-depleting 

substances. Thus, the projected extremely high 

atmospheric abundances of ODSs and the associated larger-scale stratospheric ozone depletions were prevented from occurring. However, 

many key questions remain: 

• Are the emission controls working as anticipated, i.e., are the atmospheric abundances 

of ODSs declining as expected? 

• Is the ozone layer recovering due to 

decreases in emissions of ODSs as predicted?

• Are the changes in UV occurring as expected with changes in ozone?

• What are the influences of other Earth 

system changes, e.g., climate and atmospheric composition, on the ozone layer 

and its recovery from the ODS-induced 

depletion?

• What are the influences of ODSs, and their 

substitutes, on other aspects of the Earth 

system, especially climate?

The extent of the ozone layer depletion for a given emission differs depending on the location 

(e.g., latitude) and time (e.g., season). Because 

many ODSs have lifetimes of many years in 

the atmosphere, the depletion of stratospheric 

ozone is a global problem, and emissions of 

ODSs anywhere on the globe contribute to the 

ozone layer depletion. Therefore, the observed 

ozone depletion in a given region will not be directly related to the emissions from that region. 

Yet, it is appropriate to ask: what is the contribution of one nation, or region, to the depletion of 

the global ozone layer? And, how do the ODSs 

influence stratospheric ozone, and hence UV, 

in a specific region or over a specific nation? Of 

course, it may not be feasible to answer these 

questions completely at the present time, given 

our current (and evolving) state of knowledge.

This Synthesis and Assessment Product (SAP) 

of the Climate Change Science Program 

(CCSP), SAP 2.4, addresses key issues related 

to the stratospheric ozone layer, including its 

changes in the past and its expected evolution in 

the future. Also, it takes account of the current 

abundances and emissions of ozone-depleting 

substances. Further, it synthesizes the best 

available information on the past and future 

levels of UV radiation at the Earth’s surface. 

Lastly, it explores the interactions between climate change and stratospheric ozone changes as 

well as the ODS changes, and briefly recounts 

the influence of stratospheric ozone changes 

on climate change. All of these topics are carried out within the context of the United States 

in order to distill a regional assessment from 

current global assessments. More specifically, 

this document:

• Summarizes current quantitative information on sources (i.e., emissions), sinks (i.e., 

the removal pathways and their speed), 

and abundances of ozone-depleting substances as well as associated uncertainties; 

describes how the combined influence of 

chlorinated and brominated ODSs in the 

stratosphere can be quantified, and how all 

these are likely to change in the future.

• Discusses levels of ozone in various regions of the stratosphere, including the 

polar regions, paying special attention to 

the Antarctic ozone hole.

• Provides information on the past, current, 

and anticipated future levels of ultraviolet 

radiation.

• Provides an assessment of the impact of 

changes in both climate and atmospheric 

composition on the future of the ozone 

layer.

• Provides a brief assessment of the contribution of ozone-depleting substances on 

forcing of climate because these chemicals 

are also greenhouse gases.

• Describes how these findings relate to human activities, with a particular emphasis 

on the United States. Special emphasis has 

been placed on quantifying the contributions of the United States to the global 

The issues of this 

Synthesis and 

Assessment Product 

are addressed within 

the context of the 

United States in 

order to distill a 

regional assessment 

from current global 

assessments.
27

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

BOX 1.2: A Simplified Representation of the Production and Removal of Ozone in the 

Atmosphere - the Processes that Determine the Abundance of Ozone

Oxygen molecules (O2) are broken apart by the harsh UV radiation in the stratosphere to produce atomic oxygen, 

which reacts further with oxygen molecules to make ozone (O3). The ozone in the stratosphere is removed 

predominantly via catalytic chemical reactions that regenerate the catalysts. The catalysts include atoms and 

radicals produced in the stratosphere from the breakdown of various chemicals emitted at the Earth’s surface. 

They include naturally occurring chemicals such as nitrogen oxides and hydrogen oxides, as well as humanemitted chemicals containing chlorine and bromine atoms, such as chlorofluorocarbons (CFCs) and brominecontaining halons that are used as fire extinguishants. These human-emitted species, referred to as ozonedepleting substances (ODSs), are 

of concern for the depletion of 

the ozone layer. The destruction 

pathway marked “Pathway 1” in 

Box Figure 1.2a is predominant 

outside of the springtime polar 

regions, while the pathway marked 

“Pathway 2” is dominant in the 

springtime polar ozone depletion 

including the Antarctic ozone hole. 

Because of the nature of these 

chemical processes, as discussed 

above, a very small amount of the 

catalyst (for example, chlorine 

atoms from CFCs) can destroy 

a large amount of stratospheric 

ozone. In addition to these 

chemical processes, transport 

of ozone (redistribution) is key 

to determining the abundance of 

ozone in a given location.

In contrast to the stratosphere, 

in the troposphere ozone is made 

using near UV and visible radiation (i.e., longer wavelength) because the higher energy, harsh UV 

(shorter wavelength) is screened 

out by the stratospheric ozone 

layer. This tropospheric ozone production process requires nitrogen 

oxides, mostly from combustion, 

and volatile organic compounds. 

Unlike stratospheric ozone, tropospheric ozone is removed not only 

by chemical reactions but also by 

other processes including contact 

with the surface. The transport of 

ozone from the stratosphere to 

the troposphere is important as an 

ozone source in certain regions.

Box Figure 1.2a Highly simplified schematic representation of the chemical 

processes that lead to the production and removal of stratospheric ozone. The 

catalysts include both natural and human-emitted species, including chlorine and 

bromine from ozone-depleting substances. See Chapter 3 and Appendix A for 

further details.

Box Figure 1.2b Schematic representations of the chemical processes that 

lead to the production and removal of ozone in the troposphere. The VOCs 

include methane. Other species such as hydrogen also can act as ingredients 

for ozone production in place of VOCs and CO.
The U.S. Climate Change Science Program Chapter 1

28

amounts of ODSs. Further, given the influence that ODSs and substitute chemicals 

have on climate, this Product attempts to 

calculate the contributions to the relief of 

climate change via reductions in the emissions of ODSs and switching over to more 

climate-friendly and ozone-friendly CFC 

substitutes.

The primary sources of information for this 

report are the World Meteorological Organization (WMO) / United Nations Environment 

Programme (UNEP) 2006 assessment on the 

ozone layer Scientific Assessment of Ozone 

Depletion: 2006 (WMO, 2007), and the 2005 

Special Report of the Intergovernmental Panel 

on Climate Change (IPCC) on Safeguarding 

the Ozone Layer and the Global Climate System – Issues Related to Hydrofluorocarbons 

and Perfluorocarbons (IPCC/TEAP, 2005) 

and references therein. In addition, this report 

bases some findings on a few peer-reviewed 

publications of direct import to this issue that 

have become available since the finalization of 

the two international assessments. The report 

was initiated before the release of the IPCC 

Fourth Assessment Report (AR4) (IPCC, 2007). 

Therefore, this report does not rely on the IPCC 

AR4; however, some key pertinent issues from 

the IPCC report are used in a few instances 

where updated information was essential. They 

are noted as such in those chapters.
Current Trends, Mixing Ratios, 

and Emissions of Ozone-Depleting 

Substances and Their Substitutes

CHAPTER

2

29

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

KEY FINDINGS

Convening Lead Author: Stephen A. Montzka, NOAA

Lead Authors: John S. Daniel, NOAA; Jeff Cohen, U.S. EPA; 

Kenneth Vick, USDA

Measures of production, consumption, emission, and atmospheric abundances of ozone-depleting 

substances (ODSs) and their substitute chemicals provide a coherent picture of how the Montreal 

Protocol has brought about substantial changes in the chemical composition of the atmosphere. 

All measures point to a shift away from ozone-depleting substances and toward increases in 

substitute chemicals. This shift will continue to reduce stratospheric ozone depletion and has had 

notable climate benefits. These different measures, some of which are independent, are discussed 

separately here:

From data reported by industry for the globe and for the United States1:

• Owing to the Montreal Protocol, by 2005, the summed, global annual production and 

consumption of ozone-depleting substances for uses regulated by this Protocol had decreased 

95% from peak amounts reported during the late 1980s2. Summed U.S. production and 

consumption of these substances for these regulated uses declined by 97-98% over this same 

period2.

• Use of substitutes for the more potent ozone-depleting gases has increased over time, but 

these chemicals are much less efficient at depleting stratospheric ozone than the chemicals 

they replace3. 

• Declines in overall U.S. consumption of ozone-depleting substances and substitute chemicals 

through 2005 for uses regulated by the Montreal Protocol have been more rapid than total 

global declines. When ozone-depletion influences are considered4

, the fractional contribution 

1

 Production and consumption amounts for uses of ODSs regulated by the Montreal Protocol were obtained from 

United Nations Environment Programme (UNEP) compilations of data reported to them (UNEP, 2007) and, for 

magnitudes of unregulated and unrestricted uses, from UNEP Technical Option committee reports (UNEP/MBTOC, 

2007; UNEP/CTOC, 2007); global production data for HFCs was taken from IPCC-TEAP (2005).

2

 Consumption is defined here and in the Montreal Protocol as amounts produced plus imports minus exports 

of a substance or group of substances. Production is defined as amounts produced minus the sum of amounts 

destroyed or used in feedstock (non-dispersive) applications. Consumption should equal production on a global 

scale averaged over time. In this Key Finding, magnitudes of production and consumption have been multiplied by 

weighting factors that are Ozone Depletion Potentials (see footnote #4).

3

 The more potent and abundant ozone-depleting gases referred to here include chlorofluorocarbons, or CFCs, 

halons, methyl chloroform, and carbon tetrachloride. Chemicals considered to be substitutes include the hydrochlorofluorocarbons, or HCFCs, and the hydrofluorocarbons, or HFCs.

4

 Weighting factors are applied to consumption, production, emission and banks of ODSs throughout this document 

that approximate the ozone depletion influences and the direct or indirect climate effects of these chemicals so as 

to allow consideration of them on an equivalent basis and as sums. These weighting factors account for the wide 

range of influences different chemicals have on ozone and climate. In the case of ozone, the weighting factors are 
The U.S. Climate Change Science Program Chapter 2

30

of the United States to annual global consumption of ozone-depleting substances (ODSs) in 

data reported to UNEP5 for all regulated, dispersive uses decreased by more than half, from 

a mean of 24 (±2)% in 1986-1994 to 10 (±2)%, on average, during 2001-2005. This decline is 

noted despite an increase in U.S. consumption of methyl bromide (CH3Br) relative to global 

consumption in recent years. When direct and indirect climate effects of these chemicals are 

considered4, the contribution of the United States to total global consumption of ODSs for 

regulated, dispersive uses also decreased from the late 1980s to 2005, though the precise 

magnitude of this decline is sensitive to our understanding of the indirect climate forcing from 

ODSs related to stratospheric ozone changes.

• Declines in U.S. consumption for uses regulated by the Montreal Protocol have been slightly 

faster than phase-out schedules for all developed countries in the adjusted and amended 

Montreal Protocol for most ODSs. Consumption for methyl bromide was notably larger than 

this scheduled allotment in 2005 and 2006 (by 4.3 and 4.1 ODP-kilotons [Kt], respectively) 

because of Critical Use Exemptions (CUEs)6.

• Global consumption of methyl bromide for all fumigation-related uses declined by a factor 

of two from 1997 to 2005 despite substantial consumption in applications not regulated or 

restricted by the Montreal Protocol. Nearly half (43%) of the global, industrially-derived 

emissions of CH3Br during 2005 arose from QPS6 consumption not restricted or limited by 

the Montreal Protocol. 

• U.S. consumption of CH3Br for all fumigation uses declined 40% from 1997 to 2005 despite 

enhanced Critical Use Exemptions6 and QPS6 consumption since 2001. Enhanced Critical Use 

Exemptions caused the annual U.S. contribution to global CH3Br consumption for regulated 

and restricted uses in data reported to UNEP to increase from 23 (±4)% during 2000-2003 to 

36 (±1)% during 2004-2005. In the United States during 2001-2006, the additional consumption 

of methyl bromide for fumigation not restricted or limited by the Protocol (QPS use) was, on 

average, 57 (±20)% of the amounts used and reported to UNEP5 for regulated applications and 

had increased by 13% per year, on average, during 2001-2005.

• The mix of ozone-depleting chemicals produced throughout the globe has changed over time in 

response to the Montreal Protocol. In 2005, global production weighted by Ozone Depletion 

Potentials (ODPs) for relevance to ozone depletion was dominated by chlorofluorocarbons, 

or CFCs, (50%), hydrochlorofluorocarbons or HCFCs (33%), and CH3Br (11%); in the United 

States, ODP-weighted consumption was dominated by HCFCs (54%) and CH3Br (34%). 

When weighted by overall climate influences [net Global Warming Potentials (GWPs)], 

global production in 2005 was accounted for primarily by HCFCs and hydrofluorocarbons 

(HFCs) in similar amounts, a somewhat lesser contribution from CFCs, and very small or 

negative contributions from halons, CH3Br, and other chemicals. In the United States, direct 

and net GWP-weighted consumption was dominated by HFCs and HCFCs with only small 

contributions from CFCs and other ODSs. Current estimates of global HFC production have 

large uncertainties owing to restrictions on reporting production magnitudes when less than 

three manufacturers produce a given chemical.

• Future emission rates from banks7 will play a substantial role in determining future mixing ratios 

Ozone Depletion Potentials (ODPs) with units of ODP-Tons or ODP-Kt; 1 ODP-Kt=1 billion grams multiplied 

by the ODP of a given chemical. In the case of climate, the weighting factors are 100-year direct or net Global 

Warming Potentials (GWPs), where net GWPs include the radiative influence of stratospheric ozone depletion. 

Units for quantities weighted by 100-year GWPs are expressed equivalently by, for example, GWP-Tons or gigatons 

(Gt) CO2-equivalents. Additional descriptions of these weighting factors appear in Box 2.2 and the main chapter 

text, and tables of the weighting factors used here appear in Appendix 2A.

5

 The United Nations Environment Programme (UNEP) compiles and publishes global and national statistics on 

production and consumption of ODSs based upon data reported to them in order to monitor compliance with 

the adjusted and amended Montreal Protocol (UNEP, 2007).

6

 QPS refers to quarantine and pre-shipment use of an ODS, specifically CH3Br. Though reporting requirements 

exist for this dispersive use, it is not restricted or scheduled for phase-out in the Montreal Protocol and these 

use magnitudes are not included in data published by UNEP to assess compliance with the Protocol. CUEs refer 

to Critical Use Exemptions for consumption of an ODS above existing Montreal Protocol allotments; they are 

approved only on a case-by-case basis and are included in amounts reported to and published by UNEP.

7

 Banks represent the amount of a chemical that has been produced but not yet emitted or chemically altered. 

They exist either in reserve storage or in current applications. Bank magnitudes are derived from the U.S. EPA’s 

Vintaging Model analysis of sales and use of ODSs and substitutes in the U.S. (U.S. EPA, 2007). Owing to a lack of 

available data at this time, U.S. bank estimates presented here do not include stockpiles of halons.
31

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

for some ODSs. Banks in the U.S. and throughout the globe in 2005 are estimated to have been 

7 to 16 times larger than emissions during this year, when weighted by their potential influence 

on climate or ozone depletion. CFCs accounted for the largest fraction of 2005 banks in the 

United States and throughout the globe. The U.S. EPA has classified approximately one-quarter 

of U.S. banks in 2005 as being accessible (210 ODP-Kt8, roughly two-thirds, i.e., approximately 

140 ODP-Kt, accounted for by halons; and 1.9 (0.9-2.2) gigatons (Gt) carbon dioxide (CO2)-

equivalents, of which HCFCs account for approximately 1.0 (0.9-1.1) Gt CO2-equivalents)8. 

Additional halon is likely present in stockpiles, but these amounts are not included in these 

estimates of U.S. banks owing to a lack of available data at this time.

• Emission histories derived from global ODS production and consumption data and assumed 

release functions have large uncertainties but suggest strong declines in global emissions of 

most ODSs other than HCFCs. 

From national data quantifying applications that use ozone-depleting substances and 

substitutes the U.S. EPA has derived U.S. emission histories starting in 1985. Though these 

emissions estimates are recognized to have substantial uncertainties, they suggest that:

• Total emissions of ODSs and substitutes from the United States have declined substantially 

since the late 1980s. By 2005, U.S. emissions are estimated by the U.S. EPA to have declined by 

81%, when emissions are weighted with factors relevant to ozone depletion. When weighted 

with factors relevant to climate, annual U.S. emissions of ODSs and substitutes including HFCs 

declined 74% (between 63 and 76% when indirect climate influences associated with ozone 

depletion are also included) over this same period.

• The United States accounted for a substantial fraction of global atmospheric mixing ratios of 

individual ODSs and HFCs measured in 2005, though precise quantification of these contributions 

is difficult owing to incomplete emission histories for most ODSs. The results suggest that U.S. 

emissions accounted for between 10 and 50% of the global atmospheric abundances of most 

ODSs and substitute chemicals measured in 2005, 17-42% of the tropospheric chlorine, 17-35% 

of the tropospheric bromine, and 15-36% of the tropospheric Equivalent Effective Chlorine 

(EECl)9

 arising from these chemicals in that year.

• Changes in atmospheric chlorine and bromine inferred from U.S. emissions estimates of 

chemicals regulated by the Montreal Protocol have less uncertainty than absolute amounts. The 

data suggest that atmospheric chlorine from U.S. emissions has declined steadily since 1995, 

but atmospheric bromine from U.S. emissions in 2005 was similar to 1998 levels primarily as a 

result of recent increases in exempted critical uses6

 and for QPS6 uses of CH3Br.

• Atmospheric changes derived from U.S. emissions of chlorinated and brominated ODSs indicate 

a decline in total reactive halogen (EECl)9 arising from U.S. emissions through 2005, but a 

substantially slower rate of decline since 2003. The slower overall decline in 2004-2005 was 

because of the increases in U.S. emissions of brominated gases during these years (primarily 

CH3Br).

• The direct climate influence (as direct radiative forcing)10 arising from the atmospheric 

abundances of ozone-depleting substances and substitute chemicals attributable to U.S. 

8

 Accessible banks are amounts of ODSs in use in fire extinguishers, refrigeration, and air conditioning sectors (not 

foams). While the accessible bank magnitude given was derived with direct GWP weighting, the ranges of bank 

magnitudes given in parentheses were derived with net GWPs, i.e., where consideration of the indirect, ozone 

depletion influences of ODSs are included (see Box 2.2 for more details).

9

 Equivalent Effective Chlorine, or EECl, is an index to approximately quantify the overall effect of ODSs on 

stratospheric ozone. It is calculated from surface measurements of ozone-depleting substances and accounts 

for the ODSs having different numbers of chlorine and bromine atoms, for the enhanced efficiency by which 

bromine atoms destroy ozone relative to chlorine, and the different rates at which ozone-depleting substances 

decompose in the stratosphere and liberate chemical forms of chlorine and bromine that can participate directly 

in stratospheric ozone-depleting reactions. Equivalent Effective Stratospheric Chlorine (EESC) is a related index, 

except that time lags associated with transporting air from the troposphere to the stratosphere are considered. 

These indices are described additionally in the text and in Box 2.7.

10 Direct radiative forcing is an estimate of the direct climate influence of a chemical and is expressed as energy per 

area (Watts per m2

). It is calculated with knowledge of how a chemical absorbs infrared light in certain wavelength 

regions (its radiative efficiency) and is directly proportional to its atmospheric abundance for the less abundant 

greenhouse gases. Direct forcings do not include indirect radiative effects associated with feedbacks, such as 

those related to ozone depletion. Net forcings discussed in this chapter include the indirect forcings related to 

stratospheric ozone depletion (see Box 2.2 for further details).
The U.S. Climate Change Science Program Chapter 2

32

emissions is estimated as having been between 0.067 and 0.16 W per m2 in 2005. This U.S. 

contribution amounted to between 19 and 49% of the total global direct climate influence 

of these chemicals of 0.34 W per m2. When indirect climate influences of ODSs related to 

stratospheric ozone depletion are considered, the U.S.-attributed forcing is between 0.04 

and 0.18 W per m2.

Direct observations of the atmosphere provide an independent assessment of the Montreal 

Protocol’s success in reducing atmospheric abundances of ODSs and ozone-depleting 

chlorine and bromine. These observations show that:

• The global atmospheric abundances of all ODSs are responding to changes in global production 

and consumption magnitudes. Atmospheric mixing ratios of the most abundant CFCs, the 

most abundant chlorinated solvents, and CH3Br are now decreasing. Increases are still 

observed for halon-1301, HCFCs, and HFCs. Methyl bromide mixing ratios have declined 

each year since global production was first reduced (1999), despite increases in Critical Use 

Exemptions recently, continued use in QPS6 applications, and substantial natural sources 

over which humans do not exert direct control.

• Global emissions magnitudes derived from global atmospheric data exhibit substantial 

declines since the 1980s, and provide independent confirmation of the large changes in global 

production and consumption as shown by UNEP5 in data reported to them. By 2005, global 

emissions had declined 77-82% compared to peak years, considering either the climate or 

ozone-depletion influences of ODSs and substitute chemicals.

• Tropospheric chlorine contained in all regulated ODSs and substitute chemicals has decreased 

since the early 1990s. Furthermore, measures of stratospheric chlorine show changes 

consistent with those observed in the troposphere. Stratospheric measurements also confirm 

that approximately 80% of stratospheric chlorine, which catalyzes ozone destruction, is from 

ODSs regulated by the Montreal Protocol. The remaining 20% is accounted for primarily 

by methyl chloride (CH3Cl), though a small contribution (~2%) is from very short-lived 

chemicals. 

• Tropospheric bromine from ODSs regulated by the Montreal Protocol has declined slowly 

since 1998. This decline has been dominated by tropospheric changes observed for CH3Br. 

Measures of stratospheric bromine show changes consistent with those observed in the 

troposphere, though a decline in stratospheric bromine is not yet discernable. These 

stratospheric measurements indicate that approximately 50% of stratospheric bromine is 

from industrially produced halons and CH3Br. The remainder is from naturally produced 

CH3Br and from very short-lived chemicals produced primarily naturally. 

• Observed changes in global atmospheric levels of ODSs containing chlorine and bromine 

demonstrate a substantial decline in the ozone-depleting halogen content of the atmosphere. 

The decrease since 1994 in the tropospheric halogen burden (EECl)9 accounted for by the 

long-lived ODSs considered here has been 20% of what would be needed to return EECl 

values to those in 1980 (i.e., before substantial ozone depletion was observed). The decline 

in the shorter-lived gases methyl chloroform (CH3CCl3) and CH3Br have contributed most 

to the observed decline. Decreases in stratospheric, ozone-depleting halogen (as Equivalent 

Effective Stratospheric Chlorine [EESC]) have been smaller because of the time delay 

associated with mixing tropospheric air into the stratosphere. 

• The combined direct radiative forcing from ODSs and substitutes including HFCs is still 

increasing, but at a slower rate than in the 1980s. This trend arises primarily from slow declines 

in atmospheric abundances of CFCs and continued increases in abundances of HCFCs and 

HFCs. The total direct contribution of ODSs and substitutes was 0.34 W per m2 in 2005 

(it is 0.18-0.38 W per m2 if the indirect ozone depletion forcing is included), compared to a 

contribution from CO2, methane (CH4), and nitrous oxide (N2O) of 2.3 W per m2

. 
33

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

INTRODUCTION

In an effort to heal the stratospheric ozone 

layer, schedules for the global phase-out of 

manmade ozone-depleting substances (ODSs) 

were set by the 1987 Montreal Protocol on 

Substances that Deplete the Ozone Layer and 

its Amendments and Adjustments. This chapter 

reviews the changes that have resulted from this 

international Protocol by assessing reported 

levels of ODS production and consumption, by 

deriving emissions with techniques independent 

of production and consumption estimates, by 

reporting on how these changes have influenced 

the atmospheric abundance of ODSs and 

chemicals used as substitutes, and by assessing 

To facilitate a rapid phase-out of ODSs, the Montreal Protocol allowed the use of hydrochlorofluorocarbons 

(HCFCs) as interim substitutes for chlorofluorocarbons (CFCs). Temporary use of HCFCs was allowed because, 

even though HCFCs contain chlorine and are ODSs, they are much less efficient at causing stratospheric ozone 

depletion than the ODSs they replaced, and, therefore, have been considered as in-kind replacements to transition 

to a non-CFC world. Elimination of ODSs (including HCFCs) in nearly all applications is anticipated as the phase-out 

schedules run their course. Most uses of ODSs have been replaced with the non-ozone-depleting, non-chlorine-, 

and non-bromine-containing hydrofluorocarbons (HFCs) and other so-called “not-in-kind” alternatives (e.g., nonsolvent-based cleaning processes, and hydrocarbon-based refrigerants). These changes have had a measurable 

influence on the global atmospheric abundance of these gases, with the result that the overall abundance of chlorine 

and bromine reaching the stratosphere has declined in recent years.

Therefore, the key issues, in the form of questions, that are related to ozone-depleting substances in the atmosphere 

and that are covered in this chapter, include:

• What is our best information on global production, consumption, and emissions of ozone-depleting substances, 

primarily CFCs and HCFCs, and HFCs, that are chlorine- and bromine-free, non-ozone-depleting, and longerterm replacements for CFCs and HCFCs? What are the associated uncertainties in these quantities?

• How can the combined influence of chlorinated and brominated ODSs in the stratosphere be quantified, and 

how is it likely to change in the future? 

• What fraction of the produced ODSs is still sequestered and could be potentially released at a later date? (i.e., 

what are the extents of the so-called “banks”?)

• What do the observations of ODS atmospheric abundances show about the levels of total atmospheric chlorine, 

bromine, and equivalent chlorine from these long-lived gases? In other words, are the atmospheric abundances 

actually responding as anticipated to restrictions set forth in the Montreal Protocol?

ODSs and halogenated chemicals used as substitutes have a second important property; they are efficient 

greenhouse gases (GHGs). As a result, they increase atmospheric heating and can influence climate. By requiring 

substantial reductions in global emissions of ODSs, the Montreal Protocol has led to societal benefits related to 

both stratospheric ozone depletion and climate change. The magnitude of this additional climate benefit has been 

diminished slightly, however, by small offsetting influences such as increased HFC emissions, and possibly by resulting 

stratospheric ozone increases, which may have a small warming influence. Therefore, it is important to know:

• What are the contributions of the various ODSs, and their substitutes, to climate forcing, in the past, now, 

and in the future?

Stratospheric ozone depletion is a global environmental issue. Yet, ODS emissions arise from various countries 

and regions. Also, the impact of ozone depletion is felt to different extents by different regions. Therefore, it is 

necessary to ask:

• What are the contributions of the United States to production and emissions of ODSs and substitute chemicals 

in the past? 

This chapter attempts to address many of these issues to the extent possible for those that fall within the purview 

of this document.

BOX 2.1: Key Issues
The U.S. Climate Change Science Program Chapter 2

34

how these atmospheric abundance changes have 

altered the influence of ODSs on stratospheric 

ozone depletion as well as their influence on 

climate. Furthermore, because this is a national 

assessment, this chapter provides estimates of 

these quantities for the United States and how 

they have changed over time. 

This chapter is organized into six sections. 

In the first (Section 2.1), changes in reported 

production and consumption magnitudes of 

ODSs and substitute chemicals are discussed. 

These quantities provide important evidence 

elucidating how the Montreal Protocol has 

influenced human activities. The Protocol 

was written to control production and 

consumption of ODSs. Accordingly, countries 

report these quantities annually to the United 

Nations Environment Programme (UNEP) 

so that compliance with the Protocol can be 

assessed. The data are derived fundamentally 

from industry’s records of production and 

international trade and provide the foundation 

for understanding how emissions of ODSs and 

substitute chemicals could change as a result of 

the Montreal Protocol. Limitations of the UNEP 

data are considered here through comparisons 

to AFEAS compilations (the Alternative 

Fluorocarbons Environmental Acceptability 

Study) (Section 2.1.2) and by considering the 

magnitudes of production and consumption 

for uses not regulated or restricted by the 

Protocol and, therefore, not included in the 

UNEP compilations (Section 2.1.3). Because 

the data compiled by UNEP are published on 

a country-by-country basis, a parallel analysis 

of U.S. consumption and production of ODSs 

and substitute chemicals is presented (Sections 

2.1.4 and 2.1.5).

In the second section (Section 2.2), emissions 

magnitudes and changes are assessed because 

they provide a direct understanding of how 

policy decisions are altering human influences 

on the atmosphere. Global emissions are 

inferred from measured changes in the chemical 

composition of the remote atmosphere (the “topdown” method of estimation). Emissions derived 

in this way provide an important independent 

check on global production and consumption 

data reported to UNEP. Top-down estimates are 

also compared to “bottom up” global emission 

magnitudes estimates, which are derived from 

sales data for different applications and timedependent ODS leak rates from these different 

applications (AFEAS, 2007; UNEP/TEAP, 

2006). As was the case for production and 

consumption, compound-dependent weighting 

factors related to stratospheric ozone depletion 

(Section 2.2.1) and climate (Section 2.2.2) 

are applied to emissions estimates to add 

relevance (see Box 2.2). Banks, i.e., amounts of 

halocarbons that were produced but that have 

not been emitted to the atmosphere, account 

for a large fraction of present-day emissions for 

some halocarbons and are explored in Section 

2.2.3. Banks are a particularly important topic 

because releases from banks account for much 

of the current emission of some ODSs, yet these 

releases are not restricted or addressed in the 

Protocol (Box 2.5). While U.S. regulations 

address recycling and venting of refrigerant, 

this represents a relatively small fraction of 

U.S. ODS banks and most, if not all, of this 

material will ultimately be vented to the 

atmosphere unless collected and destroyed. 

The contribution of emissions from other, nonrestricted influences is discussed subsequently 

(Section 2.2.5).

Annual U.S. emissions of ODSs and substitute 

chemicals are estimated by U.S. EPA (2007) 

using a model analysis of sales and use within 

the United States (Section 2.2.5). U.S. emissions 

estimates are different from “top-down” 

global emissions estimates because they rely 

on the accuracy of industry-related production 

and sales data or assessments of market 

demand for ODSs and substitute chemicals. 

Comprehensive, independent assessments of 

U.S. emissions from atmospheric observations 

are not currently possible, though some useful 

conclusions are drawn from studies conducted 

to date (Section 2.2.6). 

Atmospheric abundances of ODSs and substitute 

chemicals are discussed in Section 2.3. While 

emissions estimates provide a useful metric of 

how changes in human behavior are affecting 

the atmosphere, the influence of ODSs and 

substitute chemicals on stratospheric ozone and 

climate are dependent upon their atmospheric 

abundance, not rates of emission. The sensitivity 

of the atmosphere to emission magnitudes 

is determined by a chemical’s persistence, 

which is quantified as an atmospheric lifetime. 

In an effort to heal 

the stratospheric 

ozone layer, schedules 

for the global phaseout of manmade 

ozone-depleting 

substances (ODSs) 

were set by the 1987 

Montreal Protocol 

on Substances 

that Deplete the 

Ozone Layer and 

its Amendments 

and Adjustments.
35

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

Consistency between observed abundances of 

ODSs and substitute chemicals and calculated 

or expected abundances requires accurate 

estimates of both emissions and lifetimes 

(Section 2.3.1.1) 

Halocarbon abundances in the remote 

atmosphere attributable to U.S. emissions are 

also derived for past years (Section 2.3.2). The 

U.S. contributions to global abundances are 

derived from histories of emissions since 1985 

from the U.S. EPA (2007), and, for earlier years, 

a range of contributions of United States to 

global halocarbon emissions (Box 2.6). 

Subsequently, the overall inf luences that 

the wide ranges of changes observed for 

individual gases are having or will have on 

ozone depletion (Section 2.4) and on climate 

forcing (Section 2.5) are discussed. Quantities 

such as total chlorine, total bromine, and 

Equivalent Effective Chlorine (EECl and 

EESC, see Box 2.7) are calculated to assess 

the changing influences on stratospheric ozone 

(Section 2.4). Radiative efficiencies are applied 

to observed atmospheric changes to assess the 

direct influence these forcings have on climate 

(Section 2.5). Indirect climate influences 

related to stratospheric ozone depletion arising 

from the use of ODSs are also considered. 

Finally, though they are included throughout the 

document, findings related specifically to the 

United States are reviewed in Section 2.6. Results 

related to, for example, atmospheric abundances 

of ODSs calculated from consideration of U.S. 

emissions are summarized, as are the relative 

contributions of U.S. emissions to the measured 

global atmospheric abundances of ODSs and 

substitute chemicals (Box 2.6). Additional 

topics with enhanced relevance to U.S. policy 

are highlighted throughout the text in additional 

boxes. These include a discussion of methyl 

bromide (CH3Br, Box 2.3) and HCFCs (Box 

2.4). 

Throughout this chapter different weighting 

factors are applied to quantities such as 

production, consumption and emission of 

ODSs. These weighting factors are useful 

for considering overall changes because 

different chemicals influence ozone and climate 

to different extents (see Box 2.2 for more 

detail). With regard to stratospheric ozone, 

the weighting factors are Ozone Depletion 

Potentials (ODPs). With regard to climate, the 

weighting factors are 100-year Global Warming 

Potentials (GWPs). Two main influences are 

considered in the GWP calculations, the direct 

effect of a halocarbon on the radiative balance 

of the atmosphere, and the indirect influence 

arising from stratospheric ozone changes caused 

by a halocarbon. Here, the direct influence is 

accounted for by 100-year direct GWPs and 

the indirect influence is included by combining 

direct and indirect GWPs into net GWPs (see 

Box 2.2 for more detail). Different weighting 

factors are applied to atmospheric abundances 

of individual halocarbons to assess their 

influence on ozone or on the direct radiative 

forcing of the atmosphere. See Sections 2.4 and 

2.5 for further discussion of these factors.

2.1 Production and

Consumption of Ozonedepleting Chemicals and

Their Substitutes Derived

from Industry Estimates

2.1.1 Production and 

Consumption: Global Trends 

Historical global data on production and 

consumption or sales of ozone-depleting 

chemicals are available through databases 

compiled from different countries by UNEP and 

from different companies by AFEAS (2007). 

The data provided by these organizations show 

how dramatically dispersive, regulated uses of 

ozone-depleting chemicals have changed over 

the past 20 years in response to the adjusted and 

amended 1987 Montreal Protocol on Substances 

that Deplete the Ozone Layer and to changing 

market conditions. Data are compiled on an 

annual basis by UNEP to assess compliance 

with the Montreal Protocol. The UNEP data 

provide more complete global coverage in 

recent years than AFEAS compilations but are 

not disaggregated by chemical in all instances; 

only production or consumption data aggregated 

by compound class are publicly available for 

CFCs, HCFCs, and halons. Other limitations 

include the UNEP data only being available 

for years since 1986, and not all countries 

have reported production or consumption 

figures to UNEP for all years. Despite these 

limitations, the UNEP compilation provides 

Different chemicals 

influence ozone and 

climate to different 

extents. For this 

reason, weighting 

factors called Ozone 

Depletion Potentials 

(ODPs) and Global 

Warming Potentials 

(GWPs) are useful 

for considering 

overall changes. 
The U.S. Climate Change Science Program Chapter 2

36

Figure 2.1 Panel A: Annual global production and consumption of all regulated ODSs and substitutes (dark and light blue solid 

lines) compared to similar quantities for the United States (dark and light red solid lines), as derived from data reported to 

UNEP (UNEP, 2007). Baseline production and consumption quantities are shown as separate bars with corresponding colors 

in 1986. All of these data are weighted by compound-dependent ODPs. Panel B: Percentage contributions of U.S. consumption 

and production to global totals. Negative consumption indicates exports being larger than the sum of imports plus production 

in a given year (see Section 2.1.4.1).

Figure 2.2 The contribution of different compound classes or compounds to total global, ODP-weighted production of ODS 

and substitute chemicals reported to UNEP for regulated uses for 2005 (left panel), and the global, GWP-weighted production 

of these chemicals in the same year (right panel) (UNEP, 2007; Personal communication, the UN Ozone Secretariat, 2007; 

AFEAS, 2007). The hatched bar heights in the right-hand panel were derived with direct GWPs; the given uncertainties 

represent weighting by net GWPs (see Box 2.2). HFC production includes only the portion of global HFC data reported by 

AFEAS (2007) for HFC-134a, HFC-125, and HFC-143a. For HFC-23, production was inferred from atmospheric data (Clerbaux 

and Cunnold et al., 2007). Relative contributions of less than 1% are not included in these charts; note that the global CCl4

ODP-weighted production was –7% during 2005, though its contribution was not included in this Figure. See Section 2.1.4.1 

for additional discussion regarding negative consumption and production values.
37

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

(CH3CCl3 and CCl4 accounted for less than 1%) 

(Figure 2.2). Despite small declines in total 

production of HCFCs since 2000, the relative 

contribution of HCFCs increased substantially 

over this period so that by 2005 they accounted 

for 33% of total ODP-weighted production. 

Preliminary data suggest that global, ODPweighted consumption of HCFCs equaled ODPweighted consumption of CFCs in 2006.

Global production of ODSs and substitutes 

(unweighted data from UNEP, 2007; IPCC/

TEAP, 2005) can be weighted by direct or 

net GWPs to estimate the potential influence 

that production could have on climate forcing 

(see Box 2.2) (Figure 2.3). When weighting by 

100-year, direct GWPs is considered, the annual 

production of ODSs and substitutes declined 

by 8040 GWP-Mt from 1989 to 2005, which 

corresponds to a decline of 81% (Table 2.1) 

(Figure 2.3). A slightly smaller decline of 61 to 

83% is calculated for global production through 

2005 when net GWPs are used as weighting 

factors to account for the indirect influence of 

stratospheric ozone changes arising from the 

changing mix of ODSs (see Box 2.2). Increases 

in global HFC production have slowed the 

overall decline somewhat; production of HFCs 

critical data for assessing global and national 

changes in production and consumption for 

regulated uses of ODSs including CFCs, halons, 

carbon tetrachloride (CCl4), methyl chloroform 

(CH3CCl3), HCFCs, and CH3Br, particularly in 

recent years (UNEP, 2007). The data through 

2005 indicate that annual global production and 

consumption of ODSs and substitutes for ODSs 

has declined by 1.6-1.7×106 ODP-Tons since the 

Montreal Protocol was ratified (Figure 2.1). 

This corresponds to a 95% decline in both the 

ODP-weighted production and consumption of 

these chemicals across the globe by 2005 (Table 

2.1) (see Box 2.2). The average total global 

production and consumption in 2004-2005 was 

approximately 1.1×105 ODP-Tons per year.

In the data reported to UNEP aggregated by 

compound class, all classes showed declines 

in total global production and consumption 

during 2000-2005, though the relative decline 

was smallest for HCFCs (12 to 16%) and data for 

CCl4 are quite variable year-to-year. Production 

and consumption of CFCs still dominates the 

ODP-weighted global totals. During 2005 ODPweighted annual production (consumption) of 

CFCs accounted for 50% (48%), HCFCs 33% 

(34%), CH3Br 11% (14%), and halons 6% (5%) 

Table 2.1 Declines in reported production or consumption and derived emission of ODSs and 

substitute chemicals (including HFCs) relative to magnitudes in the late 1980s.

Region Production or Consumption 

Decline, 1989-2005 (%)a

Emission Decline 

Through 2005 (%)e

Weighted by Ozone Depletion Potentials

Globe 95b 82

United States 97-98c 81

Weighted by 100-year Global Warming Potentialsd

Globe 81 (61-83)b 77 (77-84)

United States 87 (81-88)c 74 (71-75)

a Considers production and consumption of ODSs for only dispersive uses restricted by the Montreal Protocol as 

shown by UNEP in data reported to them, plus HFC production and consumption or sales data without consideration of use. 

b Derived from the UNEP (2007) compilation of reported ODS production and consumption; AFEAS (2007) production data for HFC-134a, HFC-125, and HFC-143a; and HFC-23 production inferred from atmospheric data 

(Clerbaux and Cunnold et al., 2007). c Derived from consumption of ODSs for restricted uses reported to UNEP (UNEP, 2007) but delineated by compound by the U.S. EPA; and the U.S. EPA (2007) vintaging model estimates for HFCs.

d Declines indicated are calculated with direct GWP weighting but the ranges given in parentheses indicate the decline 

calculated when the indirect influence (and uncertainty) related to stratospheric ozone depletion is included in net 

GWP weighting factors (see Box 2.2). 

e Derived on a global scale from atmospheric data of ODSs and substitute chemicals and so includes all uses, regulated 

and not; derived on a U.S. scale from the U.S. EPA (2007) vintaging model estimates of emissions of ODSs and 

substitute chemicals. HFC global emissions in 2005 were interpolated from 2002 global estimates and the 2015 

“business as usual” scenario in IPCC/TEAP (2005). 

By 2005, annual 

global production 

and consumption 

of ozone-depleting 

substances (ODSs) 

and ODS substitutes 

declined by 

1.6 to 1.7x106

ODP-Tons since the 

Montreal Protocol 

was ratified.
The U.S. Climate Change Science Program Chapter 2

38

in 2005 is estimated here at approximately 630 

GWP-Mt (production of HFC-134a, HFC-125, 

and HFC-143a from AFEAS [2007], and 

inadvertent production magnitudes of HFC-23 

derived from atmospheric measurements 

[IPCC/TEAP, 2005]) (Figure 2.3). 

Based upon these production figures and direct 

GWPs, CFCs accounted for 24%, HCFCs 42%, 

and HFCs 34% of the global, CO2-equivalent 

production of all ODSs and their substitutes 

in 2005. The indirect influence associated 

with stratospheric ozone depletion alters these 

figures somewhat, though HCFCs and HFCs 

still account for the largest fraction of both 

direct and net GWP-weighted global production 

in 2005 (Figure 2.2). The contribution of HFCs 

considered here is an underestimate because 

global production data on HFCs other than 

-134a, -125, -143a, and -23 are not currently 

available, though these four gases alone 

accounted for 95% of the global total, GWPweighted demand for HFCs in 2002 (Campbell 

and Shende et al., 2005).

2.1.2 Production and 

Consumption: Comparing UNEP 

and AFEAS Compilations

AFEAS has compiled production and sales 

data for individual compounds for many years 

(AFEAS, 2007). Though the data compiled 

by AFEAS and UNEP are not independent, 

they do allow for some cross checking and an 

assessment of consistency in the global totals 

reported for CFCs and HCFCs (Figure 2.4). 

The AFEAS compilation only includes data 

for some ODSs (CFCs, HCFCs, and HFCs) and 

only for a subset of companies around the globe 

that are producing ODSs and their substitutes. 

While this compilation accounted for most of 

global production and sales of CFCs and HCFCs 

in the 1980s and early 1990s, it has accounted 

for a smaller fraction since. 

On an ODP-weighted basis, both the AFEAS 

and UNEP compilations show that by 2004 

annual global sales or consumption of CFCs 

and HCFCs (weighted by chemical-specific 

ODPs) declined by approximately 1×106

 ODPTons, or by 93%, since the late 1980s and since 

the ratification of the Montreal Protocol on 

Substances that Deplete the Ozone Layer by 

many countries (Figure 2.4). The totals from 

these two compilations during 1986-2004 

are slightly different; annual AFEAS sales 

figures are 0.1 (±0.03)×106 ODP-Tons lower 

than consumption reported to UNEP, on 

average. This difference may represent errors 

in accounting or reporting of data, but is most 

likely the result of consumption outside the 

companies reporting to AFEAS, such as by 

countries operating under Article 5 of the 

Montreal Protocol (so-called “developing” 

countries). Since 1995, the annual UNEP –

AFEAS difference has been 80 ± 10%, on 

average, of the consumption reported by these 

Article 5 countries.

By 2004, annual 

global sales or 

consumption 

of CFCs and 

HCFCs declined 

by approximately 

93% (1x106

 ODPTons) since the late 

1980s and since the 

Montreal Protocol 

went into force.

Figure 2.3 Global production of ODSs and HFCs compared to U.S. 

consumption estimates. Solid lines represent weighting by 100-year, 

direct GWPs. Dashed lines represent a range of total production or 

demand calculated with weighting by a range of net GWPs that include the 

indirect influence of stratospheric ozone on climate. A range of results, 

rather than a central value, is presented for net GWP weighting because 

of our incomplete understanding of how stratospheric ozone depletion 

has influenced climate (see Box 2.2). Results for subsets of different 

compound classes are also shown. Global ODS production derived from 

UNEP (2007) compilations (blue line) are compared to U.S. consumption 

data for ODSs (light red line). The additional influence of HFCs is shown 

on global (light blue and green lines) and U.S. scales (red line). The light 

blue line is derived by adding AFEAS global production data of HFC-134a, 

HFC-125, and HFC-143a (AFEAS, 2007) to the contribution of all ODSs. 

The green line includes additional inadvertent HFC-23 production derived 

from measured atmospheric trends (IPCC/TEAP, 2005; Clerbaux and 

Cunnold et al., 2007). The contribution of HFCs to U.S. GWP-weighted 

production has been estimated from the U.S. EPA vintaging model (U.S. 

EPA, 2007). The contributions from other HFCs listed in Table 2.2 are not 

included here due to a lack of production information on global scales.
39

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

Accurately assessing the overall effect of changes in production, consumption, and emission of individual gases 

requires consideration of weighting factors that account for compound-dependent influences on ozone and climate 

(Clerbaux and Cunnold et al., 2007; Daniel and Velders et al., 2007). When considering the influence of ODS 

production, consumption, and emission on ozone depletion, Ozone Depletion Potentials (ODPs) are used. Units 

of these quantities are ODP-tons. When considering the direct or net influence of ODS on climate, direct or net 

100-year Global Warming Potentials (GWPs) are used. Units on these quantities are GWP-tons or CO2-equivalent 

tons. Ozone Depletion Potentials, direct and indirect Global Warming Potentials, and other compound-specific 

parameters used in this report are tabulated in Appendix 2A.

ODPs

Ozone Depletion Potentials represent the amount of global ozone destroyed by a particular ODS per unit mass 

compared to the amount destroyed by a reference gas (usually CFC-11) per unit mass. Ozone Depletion Potentials 

provide a simple way to compare ODSs with respect to their ability to deplete stratospheric ozone and have proved 

useful to scientists and policymakers since their initial development (Wuebbles, 1983). Ozone Depletion Potentials 

take into account the number of chlorine and bromine halogen atoms in a chemical, how rapidly these halogen 

atoms become released in the stratosphere, how reactive the halogen atoms are for ozone destruction (Cl vs. Br, 

for example), and how persistent the chemical is throughout the entire atmosphere (its lifetime). Steady-state ODPs 

are most commonly used and are applicable for longer time periods since they represent the steady-state ozone 

responses to ODS perturbations. Time-dependent ODPs also have been proposed (Solomon and Albritton, 1992) 

for use when a particular time horizon or the time-dependence of relative ozone destruction is of interest. Chapter 

2 will use steady-state semi-empirical ODPs in this chapter (Solomon et al., 1992), taken from Chapter 8 of the 2006 

World Meteorological Organization (WMO) ozone assessment (Daniel and Velders et al., 2007). 

GWPs

Global Warming Potentials are analogous indices for comparing the integrated radiative impact of greenhouse gases 

(IPCC, 2007). They represent the cumulative radiative forcing of a unit mass of a gas relative to the same quantity 

for a unit mass of a reference gas (generally CO2) over some time horizon (generally 100 years). Hence, the GWP 

provides an approximate measure of the relative integrated climate forcing of a GHG. While it is acknowledged to 

be an imperfect index, it is generally true that emission of a well-mixed gas that is characterized by a larger GWP 

than another well-mixed gas will lead to a greater climate response.

There are two components to GWPs that we will consider in this chapter. The first is from the direct effect of 

halocarbons. The addition of an ODS to the atmosphere initially leads to a reduction in the outgoing longwave 

radiation at the tropopause, causing a globally averaged warming. This results from the strong infrared absorption 

by the ODS, particularly in the transparent atmospheric window region (8-12 μm). The amount of the net radiative 

imbalance (down flux minus up flux) at the tropopause per unit mixing ratio increase in an ODS is called the 

“radiative efficiency” (e.g., units of W per m2 per parts per billion by mole, ppb). The second component to GWPs 

considered in this chapter is from the destruction of ozone caused by ODSs. This ozone destruction leads to an 

additional radiative forcing that can be considered in discussions concerning the overall climate impact of ODSs. It 

is referred to as an “indirect” forcing because it is caused by the change in ozone (owing to ODSs) and not by the 

change in ODSs directly. 

Our general approach in this chapter is to use direct GWPs as the primary weighting factors when considering 

climate-relevant magnitudes of production, consumption, or emission of ODSs. Uncertainties quoted on these direct 

GWP-weighted quantities, however, represent a range of net GWP weightings that include the rather uncertain 

indirect ODS forcing arising from stratospheric ozone depletion. The indirect and net GWPs are derived here 

from the indirect GWPs in WMO (2007) by considering the recent revision of the radiative forcing attributed to 

stratospheric ozone depletion from −0.15±0.10 (IPCC, 2001) to −0.05 ±0.1 W per m2

 (IPCC, 2007) (see Table 

2A.3). 

BOX 2.2: Weighting Factors
The U.S. Climate Change Science Program Chapter 2

40

Net GWPs

Combining direct and indirect GWPs into net GWPs potentially leads to additional errors and has intentionally 

been avoided in IPCC/TEAP (2005) and WMO (2007). It is known that the direct and indirect processes will cause 

different spatial forcing, and it is possible that the climate response to these forcings will differ both in the global 

mean magnitude and in the spatial pattern (Joshi et al., 2003). Therefore, it is inaccurate to think of direct and 

indirect GWPs or forcing as being additive. For example, if the direct and indirect GWPs were to exactly cancel 

each other, it is still expected that there would be a climate response. Nevertheless, in this chapter we have opted 

to use net GWPs in many situations. Our purpose in doing so is not to present a precise net GWP quantity, but to 

provide an approximate idea of how ozone destruction may affect some of the conclusions drawn from considering 

the direct effect alone. Doing so likely provides a more complete and accurate picture of overall climate forcing 

from ODSs than would be obtained from considering direct GWPs alone. Specifically, throughout the chapter, we 

provide analyses and conclusions based first on the more accurate direct GWPs and forcing. We then also consider 

net GWPs and forcing calculated assuming an indirect forcing for a total ODS-induced ozone depletion of −0.15 W 

per m2 and +0.05 W per m2. These values are chosen to coincide with uncertainties on the −0.05±0.10 W per m2

IPCC (2007) ozone forcing estimate (± one-standard-deviation uncertainty). We neglect direct GWP uncertainties 

and other indirect GWP uncertainties, as we are not aware of a complete error analysis of these processes in the 

current literature.

GWP Uncertainties

The uncertainty in direct GWPs is stated to be about ±35% (± two-standard-deviation uncertainty), due primarily 

to uncertainties in the radiative efficiencies and lifetimes of the halocarbons and to uncertainties in the carbon cycle 

(see, e.g., IPCC, 2001; IPCC/TEAP, 2005; WMO, 2007; IPCC, 2007). Uncertainties in the carbon cycle are thought 

to lead to an uncertainty of about ±15% to the denominator of the GWP, or the CO2 absolute GWP (IPCC, 2007). 

This error contributes to the uncertainty in the absolute GWP value, but affects each GWP in the same way. When 

one is more interested in comparing halocarbon direct GWPs than in the values themselves, an effective uncertainty 

level of something less than 35% can be assumed.

Uncertainties in the indirect GWPs have not been as well quantified. Two of the most important issues likely involve 

the absolute amount of forcing caused by the ozone depletion caused by ODSs, and the relationship between ODSinduced ozone forcing with Equivalent Effective Stratospheric Chlorine (EESC). IPCC (2007) has recently reduced the 

magnitude of the estimated forcing due to stratospheric ozone changes from −0.15±0.10 (IPCC, 2001) to −0.05±0.10 

W per m2. However, this forcing is due to the total ozone change, not only the ozone change due to ODSs. Because 

a better estimate does not currently exist, we will adopt this latest estimate and ignore the potential contribution 

of non-ODS processes to the value. It is not possible for us to estimate the error induced by this assumption at the 

current time. In the procedure for calculating indirect GWPs given in Daniel et al. (1995), there are assumed EESC 

“thresholds” that lead to discontinuities in the forcing/EESC relationship. Because of these forcing/EESC changes, 

indirect GWPs exhibit a dependence on the time of emission; this time dependence has likely been relatively small 

when compared to other uncertainties for emissions between 1970 and 2010 and will not be considered in this 

chapter. Nevertheless, this oversimplification of the forcing/EESC relationship leads to likely additional errors in 

the indirect GWPs that have not been quantified. Because of all the previously discussed uncertainties and errors, 

in this chapter, the indirect and net GWPs are used to provide a very general idea of how ozone depletion might 

affect conclusions obtained from a direct GWP weighting.

Indirect Forcing

The indirect forcing owing to ozone depletion and attributable to individual compounds, compound classes, or 

from the aggregate of U.S. emissions was estimated by scaling the global indirect forcing arising from global ozone 

depletion through 2005 (−0.05±0.10 W per m2) to the fraction of global Equivalent Effective Chlorine (EECl) 

attributable to the compound, a compound class, or the U.S. in 2005 (see Figure 2.15). This procedure suggests 

that the net forcing from ozone depletion attributable to U.S. ODS emissions was between 0.04 and 0.18 W per 

m2

 in 2005 (Figure 2.17).

BOX 2.2: Weighting Factors cont’d
41

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

During the years of highest consumption 

(1986-1990), the AFEAS compilation accounted 

for the majority of global consumption of CFCs 

and HCFCs. During the last decade, however, 

the data reported to UNEP suggest that 59 

(±8)% of global annual consumption was not 

included in the AFEAS compilation.

The accuracy of these data hinges on the 

reliability of sales and import-export magnitudes 

reported to AFEAS and UNEP by individual 

companies and nations. This is difficult to 

assess quantitatively with independent methods, 

though estimates of global emissions inferred 

from atmospheric observations provide an 

independent but qualitative confirmation that 

large decreases in production and consumption 

of ODSs have indeed occurred since the late 

1980s. The smaller declines noted for emissions 

as compared to consumption or production 

(Table 2.1) likely arise in part because emissions 

of ODSs lag production by months to decades 

depending upon the specific application. The 

accuracy of production, consumption, sales, 

and emission data on a national basis is more 

difficult to assess by independent methods, 

though regional estimates of emissions and 

emission changes are an area of active research 

(Section 2.2.6).

The differences between the 

totals compiled by AFEAS 

and UNEP seem reasonable 

given the known differences 

in these databases. Finally, 

the consistency apparent in 

global total production and 

consumption data reported 

to UNEP suggests that the 

accounting of export and import 

activities has been reasonably 

accurate over time on a global 

scale (Figure 2.1).

2.1.3 Production and 

Consumption of ODSs 

and Substitutes Not 

Reported by AFEAS or 

in UNEP Compilations

In UNEP compilations, only 

production and consumption 

of ODSs for dispersive uses 

regulated by the Montreal 

Protocol are included. There 

is substantial additional production of ODSs 

for use as reagents in chemical manufacture 

of other substances (known as feedstock use) 

and for treatments to prevent the introduction 

or spread of pests and diseases during import/

export of goods (known as quarantine and 

pre-shipment (QPS) processes) that are neither 

restricted by the Montreal Protocol nor included 

in the production and consumption data 

compiled by UNEP. Global production for 

feedstock uses was estimated at 3.2×105

 ODPTon in 2002 (UNEP/CTOC, 2007), or about 1.9 

times the total production of ODSs reported 

for dispersive uses in that year (UNEP, 2007). 

Emissions from this production are estimated to 

be 0.5% of amounts produced for feedstock use, 

but this estimate does not include any additional 

emissions that may occur during use. ODSs 

produced substantially as feedstocks include 

CFC-113, CCl4, CH3CCl3, HCFC-22, HCFC142b, CH3Br, and halon-1301. 

In addition to feedstock applications, methyl 

bromide is sold for QPS applications that 

are not restricted or limited by the Montreal 

Protocol. In 2005, global production for 

QPS uses of 0.8×104 ODP-tons was similar 

in magnitude to the non-QPS production 

reported to UNEP of 1.1×104 ODP-tons (UNEP/

Figure 2.4 Comparison of annual AFEAS sales (green line) and 

annual UNEP consumption (blue line) totals for the aggregate of 

CFCs and HCFCs, weighted by Ozone Depletion Potential. Also 

shown is the annual difference (UNEP consumption minus AFEAS 

sales), and the magnitude of global consumption in countries 

operating under Article 5 of the Montreal Protocol, and nonArticle 5 countries. Data were compiled by UNEP (2007) and 

AFEAS (2007).
The U.S. Climate Change Science Program Chapter 2

42

MBTOC, 2007). Based on data for CH3Br 

use in QPS applications during 1999-2005, 

including this non-restricted production would 

increase UNEP-reported, global ODP-weighted 

production for all ODS by 2 to 9%, and it would 

influence the estimate of the total decline in 

ODP-weighted production since the late 1980s 

given in Table 2.1 only minimally (a decline 

of 94.3% when QPS is included, compared to 

94.7%–rounded to 95% in this Table–when not 

included). Global production of CH3Br for QPS 

is expected to increase in 2006-2007 (UNEP/

MBTOC, 2007).

Production magnitudes for three HFCs are 

currently reported by AFEAS. These data are 

thought to account for a large fraction of total 

global HFC production. In 2003, estimates 

of HFC-134a global production capacity 

(Campbell and Shende et al., 2005) exceeded 

AFEAS production data (AFEAS, 2007) for 

this compound by only 10%. Similar data for 

other HFCs are not currently available on a 

global or national basis primarily because of the 

relatively few number of production facilities. 

Most of HFC-23 in the atmosphere today arises 

from overfluorination during the production of 

HCFC-22 rather than direct production. As a 

result, production of HFC-23 can be estimated 

globally based upon emissions inferred from 

atmospheric measurement records (Clerbaux 

and Cunnold et al., 2007) though this would 

be an underestimate if any HFC-23 produced 

during HCFC-22 manufacture were captured 

and destroyed. On national scales, HFC-23 

production has been estimated from HCFC-22 

production magnitudes (U.S. EPA, 2007).

2.1.4 Production and Consumption: 

United States Trends for 

ODSs and Substitutes

2.1.4.1 United states production and

consumption weighted by ozonedepletion potential (ODP)

Production and consumption magnitudes of 

ODSs for regulated, dispersive uses in the 

United States are reported to UNEP as part of 

requirements associated with being a signatory 

to the Montreal Protocol (UNEP, 2007). The 

data indicate large declines in U.S. production 

and consumption of most chemicals as a result 

of the adjusted and amended Montreal Protocol. 

The total decline in annual U.S. production or 

consumption of ozone-depleting substances for 

dispersive uses restricted by the Protocol since 

the late 1980s through 2005 was 0.4-0.5×106

ODP-Tons (Figure 2.1). This represents a 

97-98% decline in both U.S. production and 

consumption of ODSs over this period (Table 

2.1). The total U.S. ODP-weighted consumption 

and production of ODSs reported to UNEP for 

2004-2005 averaged 1.2-1.3×104 ODP-Tons per 

year.

An analysis of data reported to UNEP reveals 

that the contribution of the United States to 

total global ODS production and consumption 

for regulated, dispersive uses decreased 

from a mean of 25 (±2)% in 1986-1994 to 10 

(±2)%, on average, during 2001-2005 (Figure 

2.1). In the interim years (1996-2000) large 

differences between reported U.S. production 

and consumption are apparent owing to negative 

consumption of carbon tetrachloride. Negative 

consumption is reported when exports outweigh 

the sum of production plus imports, or when 

destruction of stockpiles or feedstock use 

outweighs production in any given year. 

Though the mean contribution of the United 

States to global, ODP-weighted production 

and consumption of ODSs for regulated, 

dispersive uses has been 10 (±2)% since 2001, 

the contribution of different compound classes 

to this amount varies. Over this period the 

United States accounted for less than 3% of 

global annual consumption of CFCs, CH3CCl3, 

and halons, between 20 and 39% of HCFC 

annual consumption, and between 17 and 37% 

of CH3Br annual consumption (UNEP, 2007).

The large range observed for some compounds 

and compound classes since 2001 reflects 

changes in U.S. contributions over this period. 

In data reported to UNEP (2007) during 

2003-2005, the United States accounted for 

22±2% of global HCFC consumption during 

these years (19±1% of production). This is 

notably lower than the U.S. contribution during 

the previous decade. During 1992-2002 the 

United States accounted for 38±3% of global 

HCFC consumption (40±4% of production). 

HCFCs accounted for over half of total U.S. 

consumption in 2005 weighted by ODP, the 

remaining consumption was CH3Br (34%) and 

The total decline 

in annual U.S. 

production or 

consumption of 

ozone-depleting 

substances (ODSs) 

for dispersive uses 

restricted by the 

Montreal Protocol 

since the late 1980s 

through 2005 was 

0.4 to 0.5x106

ODP-Tons. This 

represents a 97 to 

98% decline in both 

U.S. production 

and consumption 

of ODSs over 

this period.
43

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

2.3). By 2005, the annual consumption of these 

chemicals had declined by approximately 2600 

Mt CO2-equivalents (87%, Table 2.1) from 

amounts reported and estimated for 1989. The 

ozone depletion arising from use of ODSs may 

have offset some of this warming influence. The 

magnitude of this offset can be approximated by 

considering net GWPs that include this indirect 

effect, though the uncertainties in this indirect 

influence are large (Box 2.2). With this indirect 

effect included, U.S. consumption of ODSs and 

substitutes declined by 81-88% from 1989 to 

2005 (1305-3010 Mt of CO2-equivalents). The 

total U.S. direct GWP-weighted consumption 

of ODSs and substitutes during 2004-2005 

was nearly 400 Mt CO2-equivalent (310-420 

Mt CO2-equivalent if the indirect influence 

associated with ozone depletion is included). 

The decline in CO2-equivalent consumption 

has decreased slightly faster in the United States 

than across the globe; the contribution of the 

United States to total global ODS production 

and consumption for regulated, dispersive 

uses was 30% in 1989 and 21% in 2005 when 

direct GWP weighting is used. If net GWPs 

are considered, the U.S. contribution decreased 

from 24-48% in 1989 to 20-23% in 2005. 

Whereas in the late 1980s more than 90% of 

CO2-equivalent U.S. consumption resulted 

from CFCs, in 2005 more than half of U.S. 

CO2-equivalent consumption was of HFCs and 

nearly all the rest was of HCFCs (Figure 2.5).

CFCs (12%); other compounds contributed less 

than 1% (Figure 2.5).

Consumption of CH3Br in the United States 

for dispersive uses restricted by the Montreal 

Protocol has also varied in recent years. It 

decreased from 1999 to 2002 but then increased 

from 2003-2005 owing in part to Critical Use 

Exemptions (UNEP/MBTOC, 2007). The U.S. 

methyl bromide consumption in 2003-2005 

was 1.3 to 2.8 times higher than consumption 

in 2002. Global consumption has declined 

fairly steadily since 1999 and, as a result, the 

U.S. contribution to global CH3Br reported 

consumption for uses restricted by the Montreal 

Protocol increased from 23 ± 4% during 

2000-2003 to 36 ± 1% during 2004-2005. 

Since 2005, amounts approved for Critical Use 

Exemptions (CUEs) in the United States have 

declined (UNEP/MBTOC, 2007).

2.1.4.2 United states consumption

weighted by global warming

potential (GWP)

United States consumption data for ODSs and 

substitutes (UNEP, 2007) has been combined 

with the U.S. EPA vintaging model estimates 

of HFC demand (U.S. EPA, 2007) to assess 

magnitudes and changes in U.S. consumption 

of halocarbons weighted by climate-relevant 

factors. The data suggest large declines in the 

consumption of ODSs and their substitutes when 

weighted by 100-year, direct GWPs (Figure 

The decline in 

CO2-equivalent 

consumption of 

ozone-depleting 

substances and 

substitute chemicals 

decreased slightly 

faster in the United 

States than across 

the globe from 

1989 through 2005.

Figure 2.5 Contributions of different compound classes or individual compounds to total United States, 

ODP-weighted consumption or sales of ODSs and substitutes reported to UNEP for uses restricted by the 

Montreal Protocol or estimated by the U.S. EPA (left panel), and total United States, GWP-weighted consumption or sales of these chemicals in the same year (right panel) (UNEP, 2007; U.S. EPA, 2007). The hatched bar 

heights in the right-hand panel were derived with direct GWPs; the given uncertainties represent weighting 

by net GWPs (see Box.2.2). Relative contributions of less than 1% are not included in these charts; note that 

the U.S. CCl4 ODP-weighted consumption was −16% during 2005, though this contribution was not included 

in the total or shown in the pie chart.
The U.S. Climate Change Science Program Chapter 2

44

Methyl bromide is unique among ODSs regulated by the Montreal Protocol for several reasons. 

First, natural processes emit substantial amounts, in addition to there being significant releases 

from industrial uses. Emissions arising from human-produced CH3Br accounted for 30 (20-40)% 

of global emissions during the mid-1990s before industrial production was reduced in response to 

Montreal Protocol phase-out schedules (Clerbaux and Cunnold et al., 2007). Since 1998, human 

production for all fumigant-related applications has declined by about 50%. 

Second, a substantial fraction of industrial 

production is for dispersive applications 

not restricted by the Protocol. These 

unrestricted uses, primarily in quarantine 

and pre-shipment (QPS) applications, 

have increased recently and have led to 

a slower decline in total global CH3Br 

production than suggested by UNEP 

values reported to them for assessing 

compliance with the Protocol. For 

example, during 2005 nearly half (43 

(36-49)%) of the global, industrially 

derived emissions of CH3Br were from 

uses not restricted by the Montreal 

Protocol (i.e., QPS applications) and, 

therefore, were not included in the 

production and consumption data shown 

by UNEP as reported to them (Box Figure 

2.3-1) (UNEP/MBTOC, 2007). Such use 

is expected to increase in the future 

(UNEP/MBTOC, 2007). In the United 

States, QPS consumption increased by 

about 13% per year, on average, during 

2001-2006 (U.S.EPA, 2007), leading to 

an annual consumption 30 to 80% higher 

than the annual amounts reported to 

UNEP during these years. 

Third, declines in CH3Br production and consumption have also been slowed by exemptions 

to protocol restrictions for critical uses (Critical Use Exemptions or CUEs) that have allowed 

substantial continued production and consumption past the 2005 phase-out in developed countries. 

Enhanced CUEs in the United States have resulted in higher annual consumption of CH3Br and 

an increased United States/Global consumption ratio during 2004-2005 compared to 2002-2003 

(Box Figure 2.3-1). 

Despite increases in QPS use and enhanced CUEs in recent years and variability in underlying 

natural emissions, global atmospheric mixing ratios of CH3Br have declined continuously since 1998 

(Clerbaux and Cunnold et al., 2007). While the United States contributed much to this atmospheric 

decrease through 2002, this U.S. trend reversed in 2003; the atmospheric abundance of bromine 

attributable to U.S. emissions was higher in 2004-2005 compared to 2002-2003 primarily because 

of enhanced QPS and CUEs consumption of CH3Br in the United States (Figure 2.14).

BOX 2.3: Focus on Methyl Bromide

Box Figure 2.3-1 Annual global production and U.S. consumption magnitudes for restricted uses reported to UNEP 

(UNEP, 2007) (lines with circles), which includes CUE amounts, 

compared to these reported amounts plus use in QPS applications considered (solid lines) (UNEP/MBTOC, 2007; U.S. EPA, 

2007)(feedstock uses not included). 
45

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

HCFCs were attractive substitutes for CFCs because they have similar properties to CFCs in many 

applications, but shorter lifetimes, generally fewer chlorine atoms per molecule, and, therefore, 

lower ODPs and GWPs.

In spite of these attributes, 

HCFCs still lead to stratospheric 

ozone depletion and affect 

climate. Hence, HCFCs are 

considered only temporary 

replacements for the most 

potent ODSs. Production of 

HCFC-22 causes an additional 

climate influence through the 

unintended formation of the 

byproduct HFC-23, itself a longlived, potent greenhouse gas.

The temporary nature of HCFC 

use is reflected in how developedcountry consumption totals have 

changed in recent years (Box 

Figure 2.4-1). Consumption 

has declined substantially in 

developed countries (non-Article 

5) and in the United States in 

response to the HCFC phaseout outlined in the Protocol. 

Production on a global scale has 

remained relatively constant 

over this time, however, as 

production and consumption in 

developing countries (Article 5) 

have increased dramatically. 

U.S. EPA vintaging model estimates suggest that U.S. HCFC annual emissions have increased by 

about 10% since 2002, despite U.S. reported annual consumption during 2003-2005 being about half 

of what it was from 1995-2002 (Box Figure 2.4-1). This apparent discrepancy likely arises from the 

large bank of HCFCs; while HCFC emissions were similar to HCFC consumption in 2005 (~6 ODPKt) the HCFC bank was more than ten times larger (Box 2.5). In the U.S. during 2005, HCFC-22, 

HCFC-142b, and HCFC-141b accounted for 98% of all U.S. HCFC emissions. The remainder was 

contributed by HCFC-225 (1.2%), HCFC-124 (0.6%), and HCFC-123 (0.3%).

An increased awareness of the influence ODSs have on both climate and stratospheric ozone has 

led to recent proposals for more stringent HCFC limits to future use by several Parties to the 

Montreal Protocol, including the United States. The accepted proposal speeds up the production 

and consumption phase-out schedule for non-Article 5 and Article 5 countries and moves the 

Article 5 country consumption baseline year forward to 2009-2010 from 2015. This earlier baseline 

year is expected to reduce Article 5 country consumption beginning in at least 2013, the first year 

consumption limits would be in force. The potential future implications of this accepted proposal 

on the evolution of EESC are summarized in Chapter 5.

BOX 2.4: Focus on Hydrochlorofluorocarbons (HCFCs)

Box Figure 2.4-1 Annual production and consumption totals for 

HCFCs as reported to UNEP for dispersive and regulated uses, 

weighted by ODPs (UNEP, 2007). Global production (red line) is 

compared to U.S. consumption (U.S. Cons.; blue line), consumption in 

all developed countries (non Article 5; developed countries; red dashed 

line), and both consumption and production in developing countries 

(Article 5 country consumption and Article 5 country production; 

green dashed and dotted lines; developing countries).
The U.S. Climate Change Science Program Chapter 2

46

2.1.5 United States Production 

and Consumption of ODSs and 

Substitutes Not Included in 

Published UNEP Compilations

Production and consumption of ODSs for 

chemical feedstock purposes and of CH3Br 

for QPS applications are not included in 

UNEP compilations because these uses are 

not restricted by the Montreal Protocol. While 

losses from feedstock applications are estimated 

to be small (0.5%, see Section 2.2.4), most 

CH3Br used in QPS applications is emitted 

to the atmosphere (UNEP/MBTOC, 2007). 

Furthermore, amounts of CH3Br used in 

QPS applications are substantial compared 

to amounts reported to UNEP for restricted 

uses and they have increased in recent years. 

For example, in the United States, annual 

consumption of CH3Br in QPS applications 

during 2001-2006 was 1.8-2.9 Kt, or 57 (±20)% 

of annual consumption reported by the United 

States to UNEP for restricted uses; this QPS 

use had increased by about 13% per year, on 

average, over this period (U.S. EPA, 2007).

U.S. production data for HFCs are not publicly 

available either through UNEP, AFEAS, or the 

U.S. EPA. Estimates of HFC demand and sales, 

however, are made by the U.S. EPA through 

its vintaging model (U.S. EPA, 2007). These 

estimates show how HFC use in the United 

States has increased by a factor of three over 

the past decade, when use is weighted by 

compound-dependent GWPs. HFC use in the 

United States accounted for about two-thirds of 

the CO2-equivalent consumption of ODSs and 

substitutes in 2005 (Figure 2.5). This vintaging 

model projects a doubling of CO2-equivalent 

HFC use in the United States during 2005-2015 

(U.S. EPA, 2007). 

2.2 Emissions: Ozonedepleting Chemicals

and Their Substitutes

Emissions estimates allow an understanding of 

how human behaviors influence the atmospheric 

abundances of ODSs and their substitutes, and 

how that influence has changed over time as a 

result of international agreements (such as the 

Montreal Protocol) and other factors. Only after 

chemicals become emitted to the atmosphere 

do they contribute to ozone depletion and 

Emissions estimates 

allow an understanding 

of how human 

behaviors influence 

the atmospheric 

abundances of ODSs 

and their substitutes, 

and how that influence 

has changed over 

time as a result 

of international 

agreements (such as 

the Montreal Protocol) 

and other factors.

radiative heating of the atmosphere. Nearly all 

ODSs produced ultimately become released to 

the atmosphere through direct emission (e.g., 

use in aerosol cans) or leakage during use or 

upon disposal. Methyl bromide is an exception, 

because a substantial fraction that is produced 

and applied to soils becomes destroyed through 

hydrolysis and does not reach the atmosphere. 

Global emissions can be estimated from 

production data, knowledge of release rates 

during production, use, and disposal of ODSs 

in different use applications, and information 

on the magnitude of sales for different end uses 

over time (AFEAS, 2007). Uncertainties can 

be significant in this “bottom-up” approach—

but, in general, emissions are delayed after 

production with time lags that are applicationdependent. Because these estimates rely on 

the production data considered in Section 2.1 

of this chapter, they are not independent of 

them. Furthermore, restrictions on reporting 

of production and consumption for ODS and 

substitutes can substantially influence emission 

estimates, particularly when a limited number 

of manufacturers produce a specific chemical.

Independent estimates of global emissions can 

be derived from an analysis of atmospheric 

observations. This “top-down” approach 

provides an important independent check 

on production and consumption magnitudes 

reported to UNEP, and is critical for assessing 

global emissions considering the limitations 

of the “bottom-up” methodology. The 

observationally derived emissions are based on 

the measured change in the global atmospheric 

burden of an ODS relative to the expected 

rate of change in the absence of emissions. 

Accordingly, this calculation incorporates 

the atmospheric lifetime of the ODS, which 

is derived from laboratory measurements of 

destruction rate constants (via photolysis and 

or oxidation by the hydroxyl radical, OH) and 

model-derived parameters such as photolytic 

fluxes, OH abundances, and 3-D distributions 

of ODS atmospheric mixing ratios. This 

method of estimating emissions is susceptible 

to errors in measurement calibration, in 

estimating the global atmospheric burdens of 

trace gases in the entire atmosphere from a 

few measurement locations at Earth’s surface, 

in lifetime, and in the assumption (generally 
47

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

applied) that all observed changes are the result 

of changes in emissions, not changes in loss 

rates. Atmospheric measurement techniques 

have improved over time to the extent that the 

majority of the uncertainty in this approach for 

long-lived ODSs is believed to arise from the 

estimates of lifetime and loss (UNEP/TEAP, 

2006).

Global emissions for ODSs have been derived 

with these different techniques and have been 

compared and reviewed in past WMO Ozone 

Assessment Reports (2003; 2007) and in the 

IPCC/TEAP (2005). Particular discrepancies 

in bottom-up versus top-down emission 

magnitudes were noted in IPCC/TEAP (2005) 

for the years since 1990 and were investigated 

additionally in a special Emissions Discrepancies 

report (UNEP/TEAP, 2006). In this latter report, 

the potential for rapid-release applications and 

time-dependent release functions to influence 

bottom-up emissions estimates was explored 

and a more comprehensive analysis of top-down 

uncertainties was presented. For the compounds 

studied (CFC-11, CFC-12, HCFC-22, HCFC141b, and HCFC-142b), the range (±1 sigma) 

of emissions estimated with top-down and 

bottom-up methods overlapped in nearly all 

years and, therefore, were considered to be 

consistent estimates (Figure 2.6) (UNEP/

TEAP, 2006). The uncertainty ranges are quite 

large in both approaches, however, such that 

the mean CFC-11 emissions estimated from 

these different methods differed generally by 

a factor of between 1.5 to 2. The overall trends 

in emissions estimated for these chemicals 

since 1990 were generally consistent, with the 

exception being HCFC-142b since 2000. While 

the bottom-up analysis suggests a rapid decline 

in emissions of this HCFC over this period, the 

top-down trends indicate only a small decline.

2.2.1 Global Emissions: Estimates 

Derived From Atmospheric 

Observations and Weighted by 

Ozone Depletion Potentials

Estimates of ODS emissions on a global 

scale have been derived for the past from a 

combination of atmospheric observations and 

industrial estimates (WMO Scenario A1, Daniel 

and Velders et al., 2007). This emission history 

indicates substantial declines in total ODPweighted emissions since 1990. By 2005, annual 

emissions had declined nearly 1.1 ODP-Mt from 

peak emissions in 1988. This corresponds to an 

82% decrease in global annual ODP-weighted 

emissions over this period (Figure 2.7; Table 2.1). 

Decreases in emissions of CFCs accounted for 

the majority of this decline (~80%). Decreases 

in emissions of CH3CCl3 and CCl4 accounted 

for 6 and 8% of the decline, respectively; 

emissions decreases in halons and CH3Br each 

accounted for 2-3% of the decline. Increases 

in HCFC ODP-weighted emissions have offset 

some of the overall decline since 1990; annual 

HCFC emissions increased from 1.1×104 in 

1990 to 2.2×104 ODP-Tons in 2005. Total global 

By 2005, annual 

emissions of 

ozone-depleting 

substances declined 

by 82% from peak 

emissions in 1988.

Figure 2.6 Comparison between global emissions derived from 

measured changes in the global atmospheric mixing ratio of CFC-12, 

CFC-11, and HCFC-142b (top down method; emissions bounded 

by red lines) and global emissions derived from an analysis of sales 

for different uses and estimates of releases from those uses over 

time (bottom-up method; emissions bounded by blue lines) (UNEP/

TEAP, 2006).
The U.S. Climate Change Science Program Chapter 2

48

emissions of ODSs and substitutes amounted 

to 2.5×105 ODP-Tons in 2005. Because these 

global emissions estimates are derived from 

atmospheric observations, they include the 

influence of all processes releasing ODSs 

and substitutes to the atmosphere, including 

releases from non-reported, QPS, Critical Use 

Exemptions, and all others.

2.2.2 Global Emissions: Estimates 

Derived From Atmospheric 

Observations and Weighted by 

Global Warming Potentials

When the global emission history compiled as 

the WMO scenario A1 (Daniel and Velders et 

al., 2007) is combined with global emissions 

derived for HFCs (Campbell and Shende et al., 

2005), the results indicate a substantial decline 

in total GWP-weighted emissions since the late 

1980s when the climate influences of ozone 

depletion are not included (Figure 2.8). The 

overall decline in annual emissions amounted 

to 7270 GWP-Mt by 2005, which corresponds 

to a 77% decrease from peak global GWPweighted emissions in 1988 (Table 2.1). With 

weighting by net GWPs to include consideration 

of indirect forcing from ozone changes, the 

global decline in annual emission is estimated 

to be between 3600 and 8500 GWP-Mt, or 

a decline of 77-84%. The decline integrated 

between 1988 and 2005 amounts to a decrease 

of over 90 Gt CO2-equivalents compared to 

constant emissions at 1988 levels. Declines in 

annual CFC emissions accounted for a decrease 

of 7900 (5330-8790) GWP-Mt by 2005, but 

this decline was partially offset by increases in 

HCFC and HFC annual emissions from 1990 

to 2005 of 210 (150-230) and 530 GWP-Mt, 

respectively (numbers in parentheses represent 

quantities with a range of net GWP weightings; 

see Box 2.2). 

In 2005, total global emissions of ODSs and 

substitutes are estimated to have been 2150 

GWP-Mt (direct GWP weighting; a range of 

675-2600 is calculated with the range of net 

GWPs). CFCs accounted for 810 (510-910), 

HCFCs 590 (470-630), and HFCs 625 GWP-Mt 

of this emission (HFC global emissions for 2005 

were interpolated from 2002 estimates and for 

Figure 2.7 Panel A: Aggregated emissions of ODSs derived for the entire globe (blue line; Clerbaux and Cunnold et al., 2007; 

Daniel and Velders et al., 2007) and for the U.S. (red line; U.S. EPA, 2007) over time, weighted by ODP. Lighter lines represent 

projections into the future. Panel B: The percentage of emissions (weighted by ODP) contributed by the United States to the 

global total. Global emissions here are derived from atmospheric observations (Clerbaux and Cunnold et al., 2007); U.S. emissions 

are inferred from a bottom-up analysis of sales data in the U.S. (U.S. EPA, 2007).
49

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

business-as-usual scenario projections for 2015 

[IPCC/TEAP, 2005]).

2.2.3 Global Emissions: The Contribution 

of Banks and Bank Sizes

“Banks” of ODSs exist where there are reserves 

of ODSs that potentially could be released at 

a later date. Though the magnitudes of banks 

are highly uncertain, the release of ODSs from 

these banks has become the most important 

factor in projecting future emissions of many 

ODSs (e.g., CFCs and halons) for two main 

reasons. First, the production of CFCs and 

halons has diminished substantially and is 

expected to continue to decrease in the future 

in response to regulations of the Montreal 

Protocol; and second, the applications for which 

CFCs are used today tend to release ODSs only 

over many years’ time. While the continuing 

production of HCFCs remains important to their 

future evolution in the atmosphere, the HCFCs 

banks are currently large enough so that future 

emissions will also be determined by their size 

and release rates from them.

The estimated sizes of ban ks, annual 

consumption, and annual emissions for the 

CFCs are shown in Figure 2.9 (see also Box 

2.5). The United States’ contribution to these 

values is represented by the lower regions of 

each bar and amounts to nearly a quarter of 

global banks in 2005, when ODP weighting 

Banks of ozonedepleting substances 

(ODSs) exist where 

there are reserves 

of ODSs that could 

potentially be 

released at a later 

date. The release of 

ODSs from these 

banks has become 

the most important 

factor in projecting 

future emissions 

of many ODSs.

Figure 2.8 Panel A: Aggregated, GWP-weighted emissions of ODSs derived for the entire globe (blue lines; Clerbaux and Cunnold 

et al., 2007; Daniel and Velders et al., 2007) and for the U.S. (red lines; U.S. EPA, 2007). Solid lines represent weighting by direct 

GWPs (no indirect influences considered), and the lighter colored lines represent the contribution from emissions of HFCs, which 

were derived on the global scale from Campbell and Shende et al., (2005) and for the U.S. (U.S. EPA, 2007). Dashed blue and red 

lines indicate the range in overall CO2-equivalent emissions when the climate influence of stratospheric ozone depletion is included 

(as indirect GWPs). Panel B: The percentage of emissions (weighted by GWP) contributed by the United States to the global total. 

Global ODS emissions here are derived from atmospheric observations and global HFC emissions are derived from a combination 

of atmospheric observations and an analysis of production data (Campbell and Shende et al., 2005; and AFEAS, 2007); U.S. ODS and 

HFC emissions are inferred from the vintaging model, which is a “bottom-up” analysis of sales and use data in the U.S. (U.S. EPA, 

2007). U.S. HFC emissions include those from replacing ODS use and an additional small (~10-20% since 2001) contribution from 

unintended byproduct emission during HCFC-22 feedstock production.
The U.S. Climate Change Science Program Chapter 2

50

is considered. Both globally and domestically, 

the gradual decline of consumption is evident, 

with the size of the bank remaining as the 

most important driver of future emission. The 

importance of the bank is already apparent 

because the annual emission is substantially 

larger than the reported annual consumption, 

with the difference presumably coming from 

the bank.

The banks of the HCFCs similarly represent an 

important reservoir that will affect future U.S. 

and global emissions (Figure 2.10). However, 

because consumption of HCFCs has not been 

fully phased out in the developed world and is 

not yet limited in countries operating under 

Article 5 of the Montreal Protocol, current 

global consumption plays a larger relative role 

influencing current global emission rates than 

it does for the CFCs (Box 2.5).

Reducing future releases of ODSs from 

banks would necessitate recovering and 

destroying some of them. Technical feasibility 

and the economics of recovery necessarily 

play important roles in determining which 

ODS banks could be feasibly recovered and 

destroyed. Daniel and Velders et al. (2007) have 

evaluated test cases in which the 2007 total 

global banks of CFCs, HCFCs, or halons were 

recovered and destroyed in terms of the ozone 

benefits that could theoretically be achieved 

as a result. This information can be found in 

Chapter 5 of this report.

The direct GWP-weighted annual production 

and emissions are compared to the global 

bank sizes for CFCs, HCFCs, and HFCs in 

Figure 2.11. The only HFC considered in this 

calculation is HFC-134a because there is a 

lack of information regarding banks of other 

HFCs.

The decreases in global production, emissions, 

and bank sizes of the CFCs represent the 

largest changes and have led to overall 

decreases in these quantities for the sum of 

these compounds in a direct GWP-weighted 

sense. Although global banks for HFC-134a 

were still a small fraction of total direct GWPweighted banks in 2005, banks of HCFCs have 

continued to increase over time and accounted 

for almost 25% of the total direct GWPFigure 2.9 Time evolution of CFC banks, annual consumption, and annual 

emission, weighted by compound-dependent ODPs. The total height of the bars 

represents global values, while the lower blue/green portions represent the U.S. 

portions. Global banks are taken from WMO (Daniel and Velders et al., 2007) 

with reliance on IPCC/TEAP (2005) bottom-up estimates used in combination 

with annual production and emission estimates. U.S. values are from the U.S. 

EPA vintaging model analysis of CFCs in the United States.

Figure 2.10 Time evolution of HCFC banks, annual consumption, and annual 

emission, weighted by compound-dependent ODPs. The total height of the 

bars represents global values, while the lower blue/green portions represent 

the U.S. portions. Global banks are taken from WMO (Clerbaux and Cunnold 

et al., 2007; Daniel and Velders et al., 2007) with reliance on IPCC/TEAP 

(2005) bottom-up estimates used in combination with annual production 

and emission estimates. U.S. values are from the U.S. EPA vintaging model 

analysis of HCFCs in the United States.
51

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

The term “bank” refers to an amount of chemical that currently resides in existing equipment or applications 

(including, for example, refrigerators, air conditioners, fire extinguishers, and foams) and stockpiles. Banked 

halocarbons are expected to be released to the atmosphere at some point in the future unless they are 

recovered and destroyed. Before scientists identified the relationship between chlorine- and brominecontaining halocarbons (now referred to as ozone-depleting substances, or ODSs) and stratospheric ozone, 

the majority of ODS usage was in fast-release applications like aerosol sprays and solvents. At that time, 

knowledge of the bank sizes was not critical to an understanding of current or even future projected ODS 

abundances.

Today the situation is far different. Most ODSs are used in slow-release applications, many of which contain a 

significant quantity of a halocarbon compared to its current atmospheric abundance; this quantity is in some 

cases much larger than the amount of halocarbon emitted annually. Accurate knowledge of these bank sizes 

and rates of halocarbon emissions from banks is now important to the future projections of many halocarbon 

abundances and to the amount of ozone that these ODSs will destroy. 

Estimates of current bank sizes are known to be highly uncertain, though various methods have been used to 

make these estimates. In ozone assessments prior to 2007 a “top-down” approach was used in which annual 

changes to bank sizes were determined from the difference between annual production estimates, taken 

from industry databases or reported amounts to UNEP, and annual emissions, estimated from atmospheric 

observations and global lifetimes. This method is particularly susceptible to systematic errors in production 

magnitudes and in the atmospheric lifetime of a chemical; significant errors can arise because the bank size 

is characterized often by small differences between large production and emission numbers and systematic 

errors can accumulate over time. A second method, which was discussed in IPCC/TEAP (2005) and is used 

in the U.S. EPA vintaging model, involves counting the number of application units that use a particular ODS 

and converting this information to a total bank size by knowing the amount of ODS typically residing in a 

single unit. This method is often called the “bottom-up” method and is independent of atmospheric lifetime 

estimates that influence top-down estimates.

Advantages and disadvantages of these methods are discussed in detail elsewhere (IPCC/TEAP, 2005; Daniel 

and Velders et al., 2007). All have significant uncertainties, and in many cases, they do not agree particularly 

well. In the most recent WMO ozone assessment report (Daniel and Velders et al., 2007), the bottom-up 

methodology for estimating the bank, in spite of acknowledged deficiencies, was thought to likely be more 

accurate than the top-down estimate and was used to project future halocarbon abundances.

BOX 2.5: Banks of Ozone-Depleting Substances (ODSs) and Substitute Chemicals

Reducing future 

releases of ozonedepleting substances 

from banks would 

necessitate recovering 

and destroying 

some of them.

Figure 2.11 Comparison of direct GWP-weighted annual global production and emission with total bank sizes. Solid lines represent contributions of CFCs, dotted lines also 

include HCFCs, and dashed lines further add the HFC-134a contribution to the bank.
The U.S. Climate Change Science Program Chapter 2

52

weighted bank in that year. The 2005 global 

banks continue to represent an important 

reservoir in terms of climate forcing, equivalent 

to about 19 Gt of potential CO2-equivalent 

emissions (Figure 2.11). When compared to 

the reductions in emissions that already have 

resulted from compliance with the Montreal 

Protocol (more than 90 Gt integrated through 

2005, direct GWP weighting), they represent 

a non-zero additional contribution to future 

The current bank sizes have important scientific and policy implications. Because banks that are not captured will 

eventually be released into the atmosphere, inaction can result in increased ozone depletion and climate forcing. 

Banks vary in how technically and cost-efficiently they can be recovered so that emission to the atmosphere is 

avoided. The U.S. EPA has identified refrigeration, air conditioning, and fire protection banks as accessible and 

potentially recoverable; other banks may also be recoverable to different extents, albeit with potentially more 

effort and higher costs. 

Sizes and relative contributions of different banks in 2005 weighted by different quantities are displayed in the figure 

(see Box Figure 2.5-1). When compared to annual emissions in 2005 (Figures 2.7 and 2.8), this analysis suggests 

that the 2005 total global and total U.S. banks were about 7 to 16 times larger (weighting by ODP or the range of 

net GWPs). 2005 accessible banks in the U.S. were about four times larger than U.S. annual emissions in that year 

(ODP or direct GWP weighting). For comparison, CO2 global emissions from fossil fuel and cement production 

have been estimated at 27-30 Gt CO2 in 2005 (IPCC, 2001). Furthermore, while CFCs account for the largest 

fraction of both global and U.S. banks in 2005 regardless of the weighting considered, they account for a much 

smaller fraction of the bank classified as accessible in the U.S. (HFCs not included). For 2005, halons accounted 

for two-thirds of the ODP-weighted, accessible U.S. bank and HCFCs a similar fraction of the GWP-weighted, 

accessible U.S. bank. The halon contributions to U.S. banks calculated here should be considered underestimates 

because stockpiles were likely non-zero in 2005. 

BOX 2.5: Banks of Ozone-Depleting Substances (ODSs) and Substitute Chemicals cont’d

Box Figure 2.5-1 The size and contribution of different compound classes to 2005 banks estimated for the globe, 

for the U.S., and for the U.S. but classified as accessible (Daniel and Velders et al., 2007; U.S.EPA, 2007). Top row includes bank estimates weighted by compound-dependent ODPs; banks in the bottom row have been weighted by direct 

(hatched bars) and net (indicated as uncertainties) GWPs (see Box 2.2). Pie charts show relative percentages; units refer 

to weighted total bank sizes, which are given at the bottom of each box in upper panel. Bar charts in lower panels give 

weighted absolute quantities; totals in the lower panels are 19 (5-24) GWP-Gt (or Gt CO2-equivalents) for the global 

bank, 6.2 (2.9-7.3) GWP-Gt for the U.S. bank, and 1.9 (0.9-2.2) GWP-Gt for the U.S. accessible bank. Note that of the 

U.S. accessible bank, HCFCs accounted for approximately 1.0 (0.9-1.1) GWP-Gt. Note that halon stockpiles are not 

included in any of the U.S. bank estimates from the U.S. EPA owing to lack of available data at this time.
53

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

In 2005, global 

emissions from 

non-restricted 

applications accounted 

for nearly half of all 

anthropogenic CH3Br 

emission. More than 

99% of this nonrestricted emission 

is estimated to arise 

from quarantine and 

preshipment use.

climate forcing and ozone depletion (see Section 

2.2.1; Velders et al., 2007). The importance of 

the Montreal Protocol and the quantification 

of the effect of future policy actions regarding 

global bank recovery and destruction have also 

been discussed in WMO (Daniel and Velders 

et al., 2007), and IPCC/TEAP (2005) reports, 

and in Velders et al., (2007). Consideration 

of the indirect forcing influences of ODSs 

associated with stratospheric ozone declines 

(and uncertainty in these influences) changes 

these figures somewhat. When weighted by 

a range of net GWPs, the 2005 global bank of 

ODSs and substitutes is estimated to be 5-24 

GWP-Gt, comprised of 6-16 GWP-Gt of CFCs, 

3-5 GWP Gt of HCFCs, and 0.9 Gt of HCFC134a (See Box Figure 2.5-1; the uncertainty 

range stated in these figures represents the 

range of influences associated with net GWPs; 

see Box 2.2).

2.2.4 Global Emissions: The 

Influence of Non-restricted 

Uses and Other Factors

As indicated in Section 2.1.3 of this chapter, 

production of ODSs for chemical feedstock 

purposes is not restricted under the Montreal 

Protocol. ODP-weighted production of ODSs 

for feedstock purposes was approximately 

1.8 times larger than overall production for 

dispersive uses reported to UNEP during 

2002 (UNEP/CTOC, 2007). Emissions during 

production of the feedstock chemical are 

estimated to be 0.5%, which corresponds to 

1600 ODP-Tons during 2002 based upon ODSs 

produced for chemical feedstock purposes, 

though this estimate does not include emissions 

that might arise after production. At a rate of 

0.5% of production, this emission amounted to 

less than 1% of total ODP-weighted emission 

in 2002.

All production of HCFC-22, including 

feedstock-related production not restricted by 

the Montreal Protocol, has an added influence 

on the atmosphere that arises from HFC-23, 

which is created from over-fluorination as 

HCFC-22 is produced. Byproduct HFC-23 

emissions account for most of the HFC-23 

present in the atmosphere today (Clerbaux and 

Cunnold et al., 2007). Feedstock production of 

HCFC-22 accounted for approximately onethird of total HCFC-22 production in 2005 

(Rand and Yamabe et al., 2005). In a “businessas-usual” scenario regarding HCFC-production 

for restricted and feedstock purposes, HFC-23 

emissions were projected to increase 60% from 

2005 to 2015. 

Production of CH3Br for QPS purposes is also 

not restricted by the Montreal Protocol (Box 

2.3). This global production was similar in 

magnitude to that used for restricted purposes 

in 2005. The emission rate for CH3Br in 

QPS uses is estimated to be 78-90% of the 

amount produced (UNEP/MBTOC, 2007). 

Based upon 2005 production for restricted 

and non-restricted (feedstock and QPS) uses 

of CH3Br, we estimate that global emissions 

from non-restricted applications accounted for 

nearly half (43 [36-49]%) of all anthropogenic 

CH3Br emission during this year. Most of this 

non-restricted emission (more than 99%) is 

estimated to arise from QPS uses.

Unlike other restricted ODSs, a substantial 

amount of CH3Br emission arises from the 

natural environment. These emissions arise 

from the oceans, wetlands, plants, and biomass 

burning; as a result, humans have little direct 

control over them. Emissions of CH3Br arising 

from industrial production are estimated 

to have accounted for 30 (20-40)% of total 

global emissions during the 1990s before 

industrial production was curtailed (Clerbaux 

and Cunnold et al., 2007).

The magnitude of variability in non-industrial 

emissions of CH3Br on annual and decadal 

time scales is not well known, and changes in 

these natural emissions could add to or offset 

the emission declines brought about by the 

Montreal Protocol. Despite these uncertainties, 

atmospheric data (Clerbaux and Cunnold et 

al., 2007) suggest that global emissions of 

CH3Br have declined each year since industrial 

production was first reduced (1999). 

2.2.5 United States Emissions 

and Banks: Estimates Derived by 

U.S. EPA Vintaging Models

While global emissions are fairly straightforward 

to derive from atmospheric measurements 

of the global background abundances of 

ODSs and substitutes, provided loss rates are 

known (Section 2.2), estimating emissions on 
The U.S. Climate Change Science Program Chapter 2

54

national or regional scales is more difficult. 

While regional atmospheric monitoring could 

potentially provide national estimates of 

emissions, to date such estimates have been 

sporadic and are based on very few sampling 

regions (Section 2.2.6). Instead, U.S. emissions 

have been estimated using “vintaging” models 

that incorporate data regarding applicationspecific sales, and leakage rates during and 

after use of ODSs and substitutes (see Box 2.6 

for further description of the vintaging model). 

With this method the U.S. EPA has estimated 

annual, U.S. emissions of ODSs and their 

replacements since 1985 (U.S.EPA, 2007). The 

data compiled by the U.S. EPA covers industrial 

production for uses restricted by the Montreal 

Protocol and for non-restricted uses such as 

feedstock and QPS applications.

When weighted by chemical-specific ODP 

values, the U.S. emissions of ODSs and 

substitute chemicals peaked in 1991 and 

declined thereafter as a result of limits imposed 

upon production and consumption by the fully 

adjusted and amended Montreal Protocol 

(Figure 2.9). By 2005, total annual U.S. 

emissions of ODSs and substitute chemicals 

had declined by 226 ODP-Kt or by 81% (Table 

2.1). Emissions have declined less from their 

peak than consumption or production (Table 

2.1) because much of those ODSs are contained 

currently in in-use foams, fire extinguishers, 

and cooling devices. United States emissions 

of ODSs and substitutes are estimated to 

have been 52.7 ODP-Kt in 2005. Emissions 

for the substitute HCFCs have increased over 

this period; U.S. HCFC emissions in 2005 

were nearly 6 ODP-Kt, which represents an 

increase of 3.5 ODP-Kt since 1990. U.S. HCFC 

emissions in 2005 were predominantly HCFC22 (87%) and HCFC-141b (7%); other chemicals 

contributed lesser amounts (HCFC-142b 4%; 

HCFC-227ca/cb 1%; HCFC-124 0.6%; and 

HCFC-123 0.3%). Over half (55%) of the U.S. 

ODP-weighted emissions in 2005 were from 

CFCs, 25% were from halons, 9% were from 

CH3Br, and 11% were from HCFCs. 

The contribution of the United States to global 

ODP-weighted emissions has varied during 

the period of available data from 18 to 35% 

(1985-2005; Figure 2.7). When weighted by 

chemical-specific 100-year direct GWP values, 

U.S. emissions estimated by the U.S. EPA 

suggest a reduction of 1640 GWP-Mt in annual 

emissions of ODSs and substitutes including 

HFCs by 2005 compared to 1991 when they 

were at their peak (a decrease of 74%; Figure 

2.8; Table 2.1). When indirect influences are 

included, a U.S. emissions decline of 900-1880 

GWP-Mt is estimated with net GWP weighting 

(see Box 2.2), which corresponds to a relative 

decline of 71-75% from peak emissions. The 

largest decline was for CFCs emissions, which 

accounted for 90-95% of the direct or net 

GWP-weighted total annual U.S. emissions in 

1985-1995 and only 40-45% of these emissions 

in 2005. By 2005, direct or net GWP-weighted 

emissions of HCFCs accounted for about onethird and HFCs one-quarter of total annual U.S. 

emissions of ODSs and substitutes. 

Independent assessments of ODS emissions 

in the United States are limited. One analysis 

derived CFC emissions by country during a 

single year, 1986 (McCulloch et al., 1994). This 

investigation was based upon consumption 

data for 1986 compiled by UNEP, AFEAS 

delineations of use by individual CFCs specific 

to different geographic regions, and emissions 

of 86-98% of consumption in each year. These 

results suggest that emissions in 1986 were 

within 50% of those estimated by the U.S. EPA 

for CFC-12, CFC-113, CFC-114, and CFC-115, 

but a factor of three higher for CFC-11. This 

discrepancy likely arises because the U.S. 

EPA analysis suggests that U.S. emissions 

of CFC-11 were not typical of other nations 

during this period. In this report we have used 

the U.S. EPA estimates because they represent 

the most in-depth and comprehensive analysis 

of U.S. emissions available. In this analysis, an 

error range of −25% to +50% is applied to U.S. 

emission estimates of ODSs after 1985.

As was touched upon in the discussion of 

ODS banks on global scales, U.S. emissions 

of many ODSs are dominated currently by 

slow releases from banks. Bank magnitudes 

in the United States are estimated with the 

U.S. EPA’s vintaging model, though these 

estimates likely have large uncertainties. The 

vintaging model estimates banks as part of its 

calculation to estimate emissions through an 

analysis of the number of pieces of equipment 

in use, the charge size of ODS in the equipment, 

The U.S. emissions 

of ozone-depleting 

substances (ODSs) 

and substitute 

chemicals peaked in 

1991 and declined 

thereafter as a result 

of limits imposed by 

the Montreal Protocol. 

By 2005, total U.S. 

annual emissions of 

ODSs and substitute 

chemicals had 

declined by 81%.
55

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

the loss rate of ODS from these applications, 

and estimates of how these variables change 

over time with input from industry. The 2005 

banks amounted to seven to sixteen years 

worth of emissions at rates estimated for 2005, 

depending upon weighting, or 830 ODP-Kt and 

6.2 Gt CO2-equivalents (direct GWP weighting; 

Box 2.5) (halon stockpiles not included). 

The U.S. EPA vintaging model analysis does 

include an estimate of how much of the ODS 

banks are accessible for recovery, where 

“accessible” refers to ODSs in current air 

conditioning, refrigeration, and fire protection 

equipment. While banked halocarbons that are 

not considered accessible by the U.S. EPA could 

be recovered and destroyed with appropriate 

policy measures, market-based incentives, and/

or certain technological advances, halocarbons 

that are not recovered and destroyed will 

eventually escape into the atmosphere. Banks 

in the United States that are classified by 

the U.S. EPA as being accessible amount to 

approximately 25-30% of the total U.S. bank 

(ODP, direct or net GWP weighting). Halons 

account for the majority of the ODP-weighted 

accessible U.S. bank in 2005, even when 

stockpiles are not included. HCFCs make up 

over two-thirds of the accessible GWP-weighted 

bank. Less than 15% of the total U.S. CFC bank 

in 2005 is considered accessible (8-14% when 

net GWP-weighted; ~6% when ODP-weighted) 

compared to over 30% of the U.S. HCFC bank 

(55-68% when net GWP-weighted and 37% 

when ODP-weighted) (Box 2.5). 

Comparing these U.S. EPA vintaging model 

bank size estimates in 2005 with the global 

banks from WMO (Daniel and Velders et al., 

2007) gives that the United States contributed 

21 and 26% to the global ODP- and direct GWPweighted banks, respectively. In this year the 

U.S. accessible banks accounted for about 5 

and 7% of the ODP- and direct GWP-weighted 

global banks. Future projections from these 

models suggest that the total U.S. bank will 

gradually account for less of the global bank 

over the next decade, shrinking to 14% and 17% 

of the global ODP- and direct GWP-weighted 

bank. Similarly, the U.S. accessible bank is 

projected to decrease to 2 and 4% of the ODPand direct GWP-weighted global banks. The 

significance of these banks to integrated EESC 

and to ODS recovery times will be addressed 

in Chapter 5. 

Interpretation of these bank comparisons 

must include consideration of the different 

assumptions and techniques used to generate the 

U.S. and the global bank estimates. Although an 

error analysis has not been performed on either 

set of numbers, the uncertainties are potentially 

large, with this uncertainty representing an 

important gap in our current understanding. 

2.2.6 United States Emissions: 

Derived From Atmospheric 

Data in Non-remote Areas 

Techniques to estimate regional or national 

emissions of ODSs that are independent of 

sales data and vintaging models are currently 

being developed. They rely on high-frequency 

atmospheric observations (multiple samples 

per day) in air downwind of source regions. 

The enhancements observed for ODSs in these 

air masses can be proportional to emission 

rates from the upwind source region provided 

dilution and mixing influences are appropriately 

considered. These estimates are specific to 

the region most directly influencing the air 

reaching a measurement site. Unfortunately, 

U.S. emissions have been derived using this 

method with data from only a small number of 

sites that may not capture regional variations in 

ODS use and emission rates. The extrapolations 

are made to the entire United States, for 

example, based upon population or by reference 

to enhancements observed in co-measured trace 

gases whose national emission rates are thought 

to be better quantified on a national scale (such 

as carbon monoxide or sulfur hexafluoride). 

While the uncertainties associated with this 

general method can be substantial and estimates 

have been made for only a small number 

of years, such an approach offers the only 

independent test of U.S. emission estimates 

derived from production and sales data in 

vintaging models (U.S. EPA, 2007). Estimates 

of U.S. emissions have been made for selected 

CFCs, CH3CCl3, and CCl4 over 1996-2003 

based on individual studies in California and in 

the northeast. The results point to a clear decline 

in U.S. emissions of CFC-12 over this period 

(Clerbaux and Cunnold et al., 2007). Although 

U.S. emissions derived for CFCs from these 

Halocarbons that are 

not recovered from 

banks and destroyed 

will eventually escape 

into the atmosphere.
The U.S. Climate Change Science Program Chapter 2

56

estimates are generally lower than from the U.S. 

EPA, those for CCl4 and CH3CCl3 are generally 

higher. On an ODP- or GWP-weighted basis, 

total U.S. emissions of ODSs derived from 

observations during 2002-2003 were about half 

of those estimated by the U.S. EPA’s vintaging 

model analysis. Because it is not known which 

method is more accurate, no modifications 

were applied to the history of ODS emissions 

compiled by the U.S. EPA (2007).

2.3 Changes in the

Atmosperic Abundance of

Ozone-depleting Chemicals

and Their Substitutes

2.3.1 Global Atmospheric Abundances

The influence an atmospheric trace gas has 

on ozone or climate generally scales with 

its atmospheric abundance. Atmospheric 

abundances reflect the integration of past 

emissions and how persistent a trace gas is in 

the atmospheric environment (i.e., its global 

lifetime). A measure of international efforts 

to minimize the deleterious environmental 

influences of ODSs and substitutes is found 

in how successful they are in reducing the 

atmospheric abundance of these chemicals.

Long-term changes in the global atmospheric 

abundances of ODSs and substitute chemicals 

are estimated with different techniques. The 

atmospheric abundances of the full suite of 

organic ODSs are generally determined at a 

small number (less than ten) of remote locations 

at Earth’s surface by independent national 

and international scientific organizations. 

These determinations are either made in real 

time by direct injection of ambient air into 

on-site instrumentation, or via the analysis 

of flask samples collected at remote sites and 

subsequently shipped to a central laboratory. 

Measurements of the most abundant ODSs 

(e.g., CFC-12, CFC-11, and HCFC-22) are 

also made with ground-based infrared solar 

absorption spectroscopy at selected sites 

across the globe. The absorption spectroscopy 

method provides a measure of the total column 

abundance of these gases above a point on 

Earth’s surface. Measurements are also made 

with absorption spectroscopy instrumentation 

onboard satellites. These instruments provide 

global observations for the most abundant 

ODSs. Long-term spectroscopic measurements 

of these chemicals reveal trends consistent with 

those observed with ground-based, flask or in 

situ sampling techniques and so are not shown 

here (Clerbaux and Cunnold et al., 2007).

Because most ODSs and their substitutes have 

lifetimes of a year or more, they are fairly 

well-mixed in the atmosphere. As a result, 

hemispheric and global atmospheric changes 

can be well captured by measurements at only 

a few remote sites. Evidence for this can be 

found in the good agreement noted between 

global surface mixing ratios derived from 

the different array of sampling locations and 

analytical techniques used by these independent 

organizations. Global surface means derived 

from these independent laboratories typically 

agree within a few percent, and often the 

small discrepancies (typically less than 5%) 

that do exist for the most abundant ODSs 

can be attributed to calibration differences 

(UNEP/TEAP, 2006). Further evidence of this 

can be found in the consistent measures of 

atmospheric composition changes provided by 

spectroscopic total column measurements and 

those provided by ground-based, in situ discreet 

sample analysis or flask sampling (Clerbaux 

and Cunnold et al., 2007).

A summary of measured trends for ODSs 

reveals a wide range of changes in atmospheric 

mixing ratios for these chemicals and their 

substitutes, primarily as a result of changes in 

production and emission brought about by the 

Montreal Protocol on Substances that Deplete 

the Ozone Layer (Figure 2.12). As of 2005, the 

surface mixing ratio and total column burden 

of the most abundant and long-lived CFC, 

CFC-12, had begun to decline slowly after 

reaching a plateau a few years earlier (Clerbaux 

and Cunnold et al., 2007). Mixing ratios of 

other ODSs, including CFC-11, CFC-113, 

CCl4, CH3CCl3, and CH3Br, have declined 

persistently over the past five to ten years at 

rates ranging from −0.5% per year to −18% 

per year. Halons have been slower to respond 

to production restrictions, though most data 

indicate that the atmospheric accumulation 

of these bromine-containing chemicals has 

slowed in recent years. Rates of accumulation 

for halon-1211 and halon-1301 estimated by 

Because most ozonedepleting substances 

and their substitutes 

have lifetimes of a 

year or more, they 

are fairly well-mixed 

in the atmosphere. 

As a result, 

hemispheric and 

global atmospheric 

changes can be 

well captured by 

measurements 

at only a few 

remote sites.
57

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

different laboratories during 2003-2004 range 

from 0 to 3.2% per year.

Global atmospheric mixing ratios of HCFCs 

(the chlorine-containing substitutes for 

CFCs and other ODSs) continue to increase 

(Clerbaux and Cunnold et al., 2007). The most 

abundant HCFC, HCFC-22, was present in 

the background atmosphere at nearly 170 ppt 

(parts per trillion by mole) in 2005 and has 

increased fairly steadily at 4 (±1) % per year 

for over a decade. Other HCFCs are one-tenth 

as abundant (or less) in the global atmosphere 

but increased during 2003-2004 at relative rates 

similar to HCFC-22. For example, though the 

global mixing ratio of the least abundant HCFC, 

HCFC-123, was on the order 0.06 ppt in 2004, 

its mixing ratio increased at about 6% per year 

during 2003-2004.

Global atmospheric mixing ratios of HFCs, 

which are in-kind substitutes that do not 

contribute any Cl or Br to the atmosphere, 

have increased quite substantially over the 

past decade (Clerbaux and Cunnold et al., 

2007). HFC-134a is the most abundant of these 

substitutes; the global mean surface mixing 

ratio in the beginning of 2006 was 36 ppt, and it 

was increasing at a rate of approximately 4.5 ppt 

per year. The long-lived HFC-23 is the second 

most abundant HFC; global measured mixing 

ratios in 2005 were approximately 20 ppt and 

were increasing. A number of other HFCs 

have been measured in the remote atmosphere 

at mixing ratios of a few ppt. After HFC-134a 

and HFC-23, the HFCs currently emitted in 

the most significant quantities are HFC-143a 

and HFC-125.

2.3.1.1 Global atmospheric abundances

of ozone-depleting substances and

substitute chemicals: measured vs. 

expected mixing ratios

As of 2005, the initial success of the Montreal 

Protocol in reducing the threat that ODSs pose 

to the stratospheric ozone layer is made clear 

by noting that production declines have led to 

declining mixing ratios or mixing ratios that 

are increasing more slowly for all regulated 

ODSs not considered to be substitutes. The 

atmospheric response to decreasing production 

varies for different gases owing to differences 

Figure 2.12 Global surface mixing ratios of ODSs and substitute chemicals observed from surface 

sampling networks (open circles), and as estimated for the past and future in WMO scenario A1 (Clerbaux and Cunnold et al., 2007; Daniel and Velders et al., 2007). Past projections are based on histories 

derived from the analysis of archived air samples, the analysis and modeling of firn-air (air trapped in 

uncompacted snow in the polar regions) samples, and historic industrial production data.
The U.S. Climate Change Science Program Chapter 2

58

Tropospheric Chlorine or 

Bromine Contribution 

(ppt) 

EECl

Contribution 

(%)c

Compound Lifetime

(years) 1985 1995 2005 2005 ODS ODSSubstitute Regulationa

CFCs 45%

CFC-11 45 622 808 759 21% X M

CFC-12 100 763 1046 1078 18% X M

CFC-113 85 114 252 237 5% X M

CFC-114 300 25 34 34 0.3% X M

CFC-115 1,700 3 8 9 0.1% X M

HCFCs 2.5%

HCFC-22 12 61 112 165 1.6% X X M

HCFC-141b 9.3 0 5 35 0.7% X X M

HCFC-142b 17.9 0 6 15 0.2% X X M

HCFC-123 1.3 0.06b 0.002% X X M

HCFC-124 5.8 1.7b 0.02% X X M

Other 

Chlorocarbons25%

CH3CCl3 5 324 330 60 1.8% X M

CCl4 26 398 412 376 11% X M

CH3Cl 1 550 550 550 12%c X None

Bromocarbons 28%

CH3Br 0.7 8.6 9.5 7.9 15%c X MA

Halon-1211 16 1.1 3.3 4.2 8% X M

Halon-1301 65 0.7 2.3 2.9 3% X M

Halon-2402 20 0.4 0.8 0.7 1.4% X M

HFCs 0%

HFC-23 270 0 0 0 0 X K

HFC-125 29 0 0 0 0 X K

HFC-134a 14 0 0 0 0 X K

HFC-143a 52 0 0 0 0 X K

HFC-152a 1.4 0 0 0 0 X K

a

 “M” represents regulation by the Montreal Protocol, “K” by the Kyoto Protocol, and “MA” is used to show that only the anthropogenic 

portion of CH3Br production and consumption is regulated. Lifetimes are from WMO (Clerbaux and Cunnold et al., 2007) and halogen 

abundances are from WMO Scenario A1 (Daniel and Velders et al., 2007), which are derived from atmospheric observations. b

 Mixing ratios for these HCFCs are for 2004, not 2005 (Clerbaux and Cunnold et al., 2007). c

 Bold percentages are calculated for the entire compound class. Methyl chloride and methyl bromide fractions are calculated including natural and manmade components. If only manmade emissions of CH3Cl and CH3Br were included (i.e., assuming anthropogenic 

contributions in 2005 of 0 ppt for CH3Cl and 1.25 ppt (7.9 - 9.5×0.7) for CH3Br), the percentages calculated for 2005 abundances 

would be CFCs: 59%; HCFCs: 3.3%; Other Chlorocarbons: 17%; Other Bromocarbons: 20%; CH3Br: 3.1%.

Table 2.2 The most abundant ODSs and substitute chemicals.
59

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

in release rates from the applications in which 

the chemicals were used, and the persistence of 

the chemicals in the atmosphere. For example, 

CH3CCl3 was used in cleaning applications in 

which release to the atmosphere followed sales 

with only a short delay (generally less than one 

year). This, combined with its relatively short 

lifetime of approximately five years, resulted 

in rapid atmospheric decreases once production 

was curtailed (Figure 2.12). The atmospheric 

abundance of CH3CCl3 has declined since 1998 

at near its lifetime-limited exponential rate, 

which is approximately 20% per year.

Conversely, CFC-11 and CFC-12 were used 

largely in foam and refrigeration applications 

in which they only slowly escaped to the 

atmospheric over decades. This, combined with 

atmospheric lifetimes of 45 to 100 years, has 

resulted in only slowly declining atmospheric 

mixing ratios (Figure 2.12; Table 2.2). The 

maximum rate of decline in the atmospheric 

abundance of a chemical with a global lifetime 

of 100 years, such as CFC-12, is 1% per year 

and would be observed only if emissions were 

negligible.

Atmospheric mixing ratios of halons also have 

been slow to respond to production declines 

(Table 2.2). This delay is attributable to large 

banks of chemical in fire-protection installations 

that are released to the atmosphere during use, 

servicing, and from leakage, and in the case of 

halon-1301, its relatively long lifetime.

Carbon tetrachloride is used as a feedstock 

for production of CFC-11 and CFC-12. As 

production of these CFCs decreased, so did 

global emissions and atmospheric mixing 

ratios of CCl4. Atmospheric declines have not 

been as rapid as expected, however, given a 

lifetime of 26 years. The slower than expected 

decline suggests the presence of substantial 

unaccounted emissions (30 to 40 Gg per 

year since the mid-1990s, or greater than 

35% of estimated emissions) or large errors 

in the estimate of the CCl4 global lifetime 

(approximately 26 years; Table 2.2).

The decline in CH3Br mixing ratios was 

somewhat faster than expected in response 

to production declines after 1998, though for 

this chemical the magnitude of the expected 

decline hinges on an accurate understanding 

of the relative importance of industrial 

emissions compared to emissions from the 

natural environment (Clerbaux and Cunnold 

et al., 2007). The measured decline has been 

more variable than observed for other ODSs, 

perhaps because of interannual variability in 

nonindustrial sources of this chemical, such 

as biomass burning. Despite these influences 

over which humans have little direct control, 

the global mean atmospheric mixing ratio of 

this chemical through 2006 has decreased each 

year since 1999, when the gradual phase-out of 

industrial production and consumption began in 

developed countries.

Increases in production of ODS substitutes 

(HCFCs and HFCs) have led to increases in 

atmospheric mixing ratios for these compounds 

over the past two decades (Clerbaux and 

Cunnold et al., 2007).

2.3.1.2 Atmospheric abundances: on the

role of variations in loss rates

Atmospheric abundances of ODSs represent 

a balance between emissions and loss. The 

Montreal Protocol has resulted in declining 

emissions of all regulated ODSs not considered 

to be substitutes. Atmospheric mixing ratios 

begin to decrease as the natural processes 

that decompose trace gases in the atmosphere 

outweigh emissions. For ODSs and their 

substitutes these natural loss pathways 

include photolytic destruction primarily in the 

stratosphere and, for ODSs containing C-H 

chemical bonds, photochemical oxidation by 

the hydroxyl radical.

Both loss processes can vary in strength 

over time because they are influenced by the 

physical and chemical state of the atmosphere. 

Accordingly, long-term or short-term variations 

in rates of photolysis and photo-oxidation 

have the potential to influence atmospheric 

mixing ratios in a way that is independent 

of emission changes brought about by the 

Montreal Protocol. Estimates of the magnitudes 

of these loss changes suggest that they are 

generally small over multi-decadal periods, 

though it has been suggested that decadal 

changes in hydroxyl radical abundance can be 

as large as 15% (Clerbaux and Cunnold et al., 

2007). Furthermore, OH reaction rate constants 

The atmospheric 

response to decreasing 

production varies 

for different gases 

owing to differences 

in release rates from 

the applications in 

which the chemicals 

were used, and 

the persistence of 

the chemicals in 

the atmosphere. 
The U.S. Climate Change Science Program Chapter 2

60

Atmospheric abundances of long-lived ODSs and substitute chemicals can be calculated from an emission history 

and a simple box model (IPCC/TEAP, 2005; WMO, 2007). A box model includes the influence of emissions and 

loss rates (or atmospheric lifetimes) to derive atmospheric abundances over time. This common and widely 

accepted approach is used in this report to estimate atmospheric mixing ratios of ODSs and substitutes arising solely from U.S. emissions of these chemicals. U.S. emission histories are derived over an entire period of 

ODS use based upon assumptions regarding a potential range of United States/Global emission ratios before 

1985 and estimates from a vintaging model analysis thereafter (U.S. EPA, 2007) (Box Table 2.6-1). The U.S. 

EPA vintaging model is a bottom-up modeling approach that considers market size, amount of ODS in each 

unit of equipment or application, and ODS substitution trends in order to estimate time-dependent, annual 

emissions and bank sizes. In the absence of U.S. production or use data for most years before 1985, a number 

of assumptions were made in order to bracket likely U.S. emissions during these early years. Such assumptions 

are essential for estimating U.S. contributions to CFC abundances, for example, because much of the pre-1985 

emissions of these long-lived chemicals are still present in the atmosphere today. In 1974, as the public became 

aware of the threat posed to the ozone layer by ODSs, CFC use in fast-release applications in the U.S. was 

dramatically curtailed. This likely resulted in substantial changes in U.S. emissions of CFCs then. Accordingly, 

a different approach was used to derive ranges of potential U.S. CFC emissions during 1975-1984 compared 

to before 1975 (Box Table 2.6-1).

BOX 2.6: On Deriving Atmospheric Abundances from U.S. Emissions

Box Table 2.6-1 Methodology for Deriving Limits to Compound-Specific, 

Annual U.S. Emissions*:

Upper Range

(emission or emission fraction)

Lower Range

(emission or emission fraction)

Pre 1975

CFCs [GE]×0.67 [GE]×0.33

Non-CFCs GEFUS/Global (1985-1990) ×1.5 GEFUS/Global(1985-1990) ÷1.5

HFC-23 α×GPF(HCFC-22)US/Global(1985-1990) ×1.2 α×GPF(HCFC-22)US/Global(1985-1990) ×0.9

1975-1984

CFCs

interpolate [GE]×0.67

in 1975 to GEFUS/Global(1985-1990) ×2 in 

1984

GEFUS/Global(1985-1990) ÷2

Non-CFCs GEFUS/Global(1985-1990) ×1.5 GEFUS/Global(1985-1990) ÷1.5

HFC-23 α×GPF(HCFC-22)US/Global(1985-1990) ×1.2 α×GPF(HCFC-22)US/Global(1985-1990) ×0.9

1985-2005

All HFCs U.S. EPA (2007) +20% U.S. EPA (2007) −10%

All others U.S. EPA (2007) +50% U.S. EPA (2007) −25%

* where GE = global annual emissions; GEFUS/Global(1985-1990) refers to the mean U.S./global emission fraction over the 

period 1985 to 1990 determined from a ratio of compound-specific emissions from the U.S. EPA (2007) vintaging 

model and WMO scenario A1 global emissions (Daniel and Velders et al., 2007); and GPF(HCFC-22)US/Global(1985-1990)

refers to the mean U.S./global production fraction for HCFC-22 during 1985 to 1990 in data reported by AFEAS 

(2007) and by the U.S. EPA (2007) and the assumption that HFC-23 emissions were proportional to HCFC-22 

production during those years. Alpha (α) indicates the fraction of HFC-23 emitted for a given amount of HCFC-22 

produced (~1% by mass, U.S. EPA[2007]) 

The main uncertainties in calculating atmospheric mixing ratios associated with U.S. emissions are associated 

with U.S. emission magnitudes and, for some gases, global atmospheric lifetimes. Emission errors stem from 

uncertainty in the ability of the U.S. EPA vintaging model to accurately capture the mean annual emissions 

from the many varied applications in which ODSs were used since 1985. Though a chemical-specific uncertainty analysis has not been performed by the U.S. EPA, an uncertainty analysis has been performed on 2005 

emissions derived by the U.S. EPA vintaging model for the high-GWP, ODS substitutes as a group (HFCs and 

PFCs). This analysis, performed with Monte-Carlo techniques, indicated a 95% confidence interval of −9% to 

+20% on 2005 emissions estimated for this class of compounds. 

Compound-specific uncertainties and uncertainties for earlier years are likely to be somewhat larger, though 

the ODS consumption and production tracking system required for compliance with the Montreal Protocol 

has added to data reliability. 
61

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

are temperature sensitive; increases in reaction 

rates of approximately 10% between OH and 

CH4, HCFCs, and HFCs are calculated for a 5°C 

increase in temperature. Because oxidation by 

OH is the primary means by which these gases 

are removed from the atmosphere, compound 

lifetimes would change similarly. Finally, future 

changes in large-scale atmospheric circulation 

patterns have been predicted in some models 

in response to increased radiative forcing from 

elevated greenhouse gas abundances and would 

likely also shorten lifetimes for ODSs where 

stratospheric losses are relatively important 

(Section 4.4.1).

2.3.2 The United States Contribution 

to Global Atmospheric Abundances

Atmospheric mixing ratio histories can be 

derived from a record of U.S. ODS emissions, 

estimates of global loss rates (lifetimes), and 

a simple global box model (Box 2.6). Mixing 

ratios calculated in this way can be compared 

to measured and calculated global mixing ratios 

to estimate the contribution of U.S. emissions 

to the atmospheric abundance of ODSs and 

substitute chemicals in the past and future. 

Though uncertainties in this analysis are 

large, the results suggest that U.S. emissions of 

ODSs and substitutes account for between 10 

BOX 2.6: On Deriving Atmospheric Abundances from U.S. Emissions cont’d

Given these considerations, we have augmented the uncertainties derived for the high-GWP ODS substitutes 

by a factor of approximately 2.5 to derive a lower and upper range of −25% to +50% on compound-specific, 

annual emission estimates derived from the vintaging model during the 1985 to 2005 period. 

These approaches have yielded estimates to expected upper and lower ranges to U.S. emissions of ODSs and 

substitute chemicals (Box Figure 2.6-1). 

Errors on calculated atmospheric mixing ratios that are associated with lifetime uncertainties depend upon the 

use period of an ODS relative to its lifetime. For chemicals that have been emitted for a period that is small 

compared to their lifetimes (CFC-12 and HFC-23, for example), lifetime uncertainties are relatively small. For 

other gases such as CH3Br and CH3CCl3 where their use period is long relative to their atmospheric lifetime, 

lifetime uncertainties are more substantial. 

Errors associated with the simple box-model approach are thought to be substantially smaller than the errors 

already discussed, especially in the case of ODSs whose atmospheric lifetimes are comparable to or much longer 

than atmospheric mixing times (see, for example, UNEP/TEAP, 2006; Clerbaux and Cunnold et al., 2007).

Box Figure 2.6-1 Ranges estimated for U.S. emissions derived from the analysis presented in Box Table 

2.6-1; units are Gg, or 109

 g. These U.S. emissions estimates are compared to North American sales data 

for CH3CCl3 (green points; Midgley and McCulloch, 1999). These data show reasonable consistency within 

the ranges estimated here for U.S. emissions, considering that the U.S. accounted for approximately 95% 

of North American consumption of this chemical in the late 1980s (UNEP, 2007) and that releases of CH3CCl3 generally occurred within a year after sales. Data to allow similar comparisons for other chemicals 

are not available.
The U.S. Climate Change Science Program Chapter 2

62

The smallest fractional contributions of 

U.S. emissions to present-day mixing ratios 

are calculated for CCl4, CFC-11, HCFC142b, HCFC-141b, and halon-1211. Larger 

contributions are estimated for some other 

CFCs, CH3CCl3, HCFC-22, halon-1301, CH3Br, 

and some HFCs (Figure 2.13).

The largest increases in U.S. contributions 

in recent years are calculated for CH3Br, 

HCFC-22, and halon-1301. The increased 

fractional contribution of U.S. CH3Br emissions 

to its atmospheric abundance arises from 

and 50% of the global atmospheric abundance 

measured for most gases in the present-day 

atmosphere (Figure 2.13). These estimates are 

most uncertain for the long-lived CFCs because 

the substantial and poorly constrained emissions 

that occurred before 1985, when U.S. EPA 

estimates begin, still contribute significantly 

to atmospheric abundances today. Uncertainties 

are smaller for gases having shorter lifetimes 

and shorter emission histories (i.e., where 

a higher percentage of total emissions have 

occurred after 1985).

Figure 2.13 Upper and lower ranges to the fractional contributions of U.S. industrial emissions 

to global atmospheric abundances of ODSs and substitute chemicals over time (see Box 2.6 for the 

derivation of U.S. emissions and a discussion of uncertainties). For HFC-23, emission from production of HCFC-22 for ODS substitution and feedstock uses is included. For the HFCs, results are 

only displayed for years when global mixing ratios are greater than 1 ppt. Methyl bromide emissions arising from natural processes are not included, i.e., 146,000 Metric Tons of emission per year 

(Clerbaux and Cunnold et al., 2007; Daniel and Velders et al., 2007).
63

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

increases in U.S. consumption compared to 

global consumption since 2002 (Section 2.1.4 

and Box 2.3). The U.S. fractional contribution 

to atmospheric mixing ratios of halon-1211, 

HCFC-141b, CH3CCl3 and some HFCs has 

decreased in recent years (Figure 2.13).

2.4 The atmospheric

abundance of aggregated

chlorine and bromine from

long-lived ODSs

Many different chemicals contribute to 

atmospheric chlorine and bromine, such as 

sea salt, pool disinfectants, CFCs, and HCFCs. 

Only those with longer lifetimes (greater than 

months) and lower water solubilities escape 

scavenging by aerosols and rain, however, 

and become efficiently transported to the 

stratosphere and contribute to ozone depletion 

(ODSs such as CFCs, HCFCs, halons, CH3Br, 

and others). Hence, tropospheric burdens of 

these long-lived ODSs are closely monitored 

because they provide a useful measure of 

changes and amounts of chlorine and bromine 

being transported to the stratosphere and that 

will ultimately become available for catalyzing 

the destruction of stratospheric ozone. 

Stratospheric abundances of chlorine and 

bromine are more difficult to regularly measure 

with high precision. Such measurements are 

useful, however, to discern whether amounts 

and changes in stratospheric chlorine and 

bromine are well described by the total amount 

of chlorine or bromine measured in longlived ODSs at Earth’s surface. Stratospheric 

measurements take advantage of the fact that 

in the upper stratosphere, nearly all organic 

compounds have become photo-oxidized and 

chlorine and bromine exist primarily in only 

one or two chemical forms (e.g., hydrogen 

chloride, HCl, and chlorine nitrate, ClONO2; 

and bromine oxide, BrO, for example). As a 

result, stratospheric measurements of these 

few inorganic chemicals provide an integrated 

estimate of how ozone-depleting halogen levels 

are changing, and whether or not these changes 

are consistent with ODS observations in the 

lower atmosphere. 

2.4.1 Atmospheric Chlorine

As a result of the restrictions on production 

and consumption of ODSs brought about 

by the Montreal Protocol, the abundance of 

chlorine measured in long-lived gases has been 

decreasing in the lower atmosphere since 1995 

and has continued to decrease through 2004 

(Clerbaux and Cunnold et al., 2007). Global 

tropospheric chlorine in long-lived chemicals 

was 3.44 ppb in 2004, or 0.25 ppb below the 

peak observed in the early 1990s. The rate of 

tropospheric decline in total chlorine from all 

regulated ODSs during 2003-2004 was slightly 

slower than four years earlier, as the influence 

of CH3CCl3 continued to diminish; the mean 

decline during 2003-2004 was 20 ppt per year 

(0.6% per year).

Approximately 80-85% of organic chlorine 

in long-lived trace gases measured in the 

troposphere is accounted for by gases regulated 

by the Montreal Protocol (CFC-11, CFC-12, 

CFC-113, CFC-114, CFC-115, HCFC-22, HCFC142b, HCFC-141b, HCFC-124, HCFC-123, 

CH3CCl3, CCl4, and halon-1211). Most of the 

remaining 15% (or 550 ppt Cl) is accounted 

for by methyl chloride, a chemical having 

predominantly (greater than 95%) non-industrial 

sources. Atmospheric mixing ratios of CH3Cl 

have been relatively constant over the past 

decade, though year-to-year variations on the 

order of a few percent can be observed at Earth’s 

surface during years with enhanced biomass 

burning (Clerbaux and Cunnold et al., 2007).

Small additional amounts of chlorine (~2% of 

the sum from regulated ODSs and CH3Cl, or 

50-80 ppt Cl) may be contributed by short-lived 

gases such as CH2Cl2, CHCl3, C2Cl4, and others 

(Law and Sturges et al., 2007). Tropospheric 

observations suggest that mixing ratios of some 

of these gases have decreased in recent years.

Because HCl and ClONO2 have quite low 

abundances in the lower atmosphere, total 

column measurements of HCl and ClONO2

from surface-based spectroscopic instruments 

provide an independent measure of stratospheric 

chlorine changes (Clerbaux and Cunnold et 

al., 2007). These measurements show that, 

after many years of consistent increases, 

total column inorganic chlorine mixing ratios 

peaked in the mid-to-late 1990s and have since 

Approximately 80 

to 85% of organic 

chlorine in long-lived 

trace gases measured 

in the troposphere 

is accounted for by 

gases regulated by the 

Montreal Protocol. 

Most of the remaining 

15% is accounted for 

by methyl chloride, 

a chemical having 

predominantly nonindustrial sources.
The U.S. Climate Change Science Program Chapter 2

64

declined. The declines have lagged behind 

the decreases observed at Earth’s surface by a 

few years because of time lags associated with 

transporting air in the lower atmosphere to the 

stratosphere.

Satellite-based spectroscopic instruments 

also have the potential to measure changes in 

stratospheric chlorine abundance over time 

(Clerbaux and Cunnold et al., 2007). To date, 

however, long-term trends determined from 

these instruments have added uncertainty 

from numerous complications related to small 

unexplained offsets in HCl measured by 

different instruments, substantial unexplained 

variability in the longest record (Halogen 

Occultation Experiment), and relatively short 

data records for other instruments that offer 

higher precision (Atmospheric Chemistry 

Experiment and Microwave Limb Sounder).

Despite these issues, satellite instruments have 

provided an important independent measure of 

stratospheric chlorine mixing ratios (Clerbaux 

and Cunnold et al., 2007). Results from these 

instruments demonstrate that mixing ratios of 

chlorine observed in the upper stratosphere 

are well explained by measured tropospheric 

abundances of long-lived ODSs regulated by 

the Protocol plus a contribution from CH3Cl of 

approximately 15%. The scatter among results 

from different instruments prevents a precise 

estimate of the contribution of short-lived gases 

to stratospheric chlorine, but they do suggest it 

is on the order of a few percent, consistent with 

the tropospheric observations.

As of 2004, ground-based air sampling results 

show that CFCs still account for most of the 

long-lived Cl in the troposphere, (62% in 

2004; Clerbaux and Cunnold et al., 2007). The 

abundance of the three most abundant CFCs has 

peaked or is decreasing in the troposphere, with 

declines in CFC abundances accounting for 

about half of the decline in total tropospheric Cl 

in 2004 (−9 ppt Cl per year). Methyl chloroform 

still strongly influences total chlorine trends 

despite its dramatically reduced atmospheric 

abundance; in 2004 it accounted for more than 

half of the observed decline in Cl (−13.5 ppt Cl 

per year). Carbon tetrachloride has declined 

fairly steadily at a rate of 1 ppt per year and has 

accounted for an annual change in tropospheric 

chlorine of −4 ppt Cl per year during the past 

As of 2004, the 

abundance of the 

three most abundant 

CFCs had peaked 

or was decreasing 

in the troposphere, 

with declines in 

CFC abundances 

accounting for about 

half of the decline in 

total tropospheric 

chlorine in that year.

Figure 2.14 A) Estimated ranges of the tropospheric abundance of Cl and Br from U.S. emissions of all regulated ODSs. These 

ranges are from upper and lower estimates of U.S. emissions (see Box 2.6 for additional information). B) Estimated ranges in the 

fractional contribution of U.S. emissions to global atmospheric mixing ratios arising from industrial production of ODSs. In both 

panels, only emissions arising from industrial production of ODSs for restricted uses plus QPS uses of CH3Br were considered; 

global and U.S. emissions of CH3Cl and CH3Br from natural processes are not included in these calculations. Chemicals included 

in these estimates of total Cl: CFC-11, CFC-12, CFC-113, CFC-114, CFC-115, HCFC-22, HCFC-142b, HCFC-141b, HCFC-123, 

HCFC-124, CH3CCl3, CCl4, and halon-1211; and of total Br include: CH3Br, halon-1211, halon-1301, and halon-2402.
65

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

decade. In 2004 CH3CCl3 accounted for 2% 

(65 ppt Cl) and CCl4 accounted for 11% (375 

ppt Cl) of all long-lived Cl in the troposphere. 

(These declines total more than 100% owing to 

the offsetting increases observed for chlorine 

from HCFCs.)

HCFCs continue to increase in the atmosphere 

as they are used as substitutes for CFCs 

and other ODSs (Clerbaux and Cunnold et 

al., 2007). They accounted for 6% of total 

tropospheric Cl in 2004, and chlorine from 

HCFCs increased at a rate of nearly 8 ppt Cl per 

year during that year. Though the increase in 

Cl from HCFCs during 2004 was significantly 

slower than observed in 1996-2000, nearterm projections of production and use, and 

continued observations since the publishing of 

the latest WMO Scientific Assessment of Ozone 

Depletion Report (Clerbaux and Cunnold et al., 

2007), show accelerating growth rates since 

2004.

Tropospheric chlorine attributable to U.S. 

emissions of long-lived ODSs also has declined 

since the early 1990s (Figure 2.14). The 

U.S. contribution to global tropospheric 

chlorine from all regulated ODSs (excluding 

consideration of CH3Cl) is estimated as being 

between 17 and 42% in 2005. Despite this 

large uncertainty range, estimates of the rate of 

change have smaller uncertainties and suggest 

that the U.S. relative contribution to global 

atmospheric chlorine from regulated ODSs has 

remained fairly constant over the past decade. 

2.4.2 Atmospheric Bromine

Bromine in the stratosphere catalyzes the 

destruction of ozone with a per-atom efficiency 

that is approximately 60 times that of chlorine 

(WMO, 2007). As a result, small mixing ratios 

of stratospheric bromine play an important part 

in controlling stratospheric ozone abundances. 

Bromine also differs from chlorine because 

emissions from regulated uses account for 

a smaller fraction of the inorganic bromine 

measured in the stratosphere. Whereas chlorine 

emissions from uses regulated by the Montreal 

Protocol accounted for approximately 80-85% 

of stratospheric chlorine at its peak abundance, 

emissions of bromine regulated by the Protocol 

accounted for approximately 50% of the 

bromine measured in the stratosphere at its 

peak abundance of 20-22 ppt (estimated by 

assuming 30% of 9.5 ppt from CH3Br as arising 

from regulated uses, plus 8 ppt Br from halons) 

(Law and Sturges et al., 2007).

Chemicals containing bromine that are regulated 

by the Montreal Protocol include halons and 

methyl bromide, though methyl bromide has 

both natural and anthropogenic emissions. 

When both natural and anthropogenic sources 

of CH3Br are considered together with the 

bromine from halons, these chemicals accounted 

for approximately 80-90% of total bromine 

reaching the stratosphere in 1998. Surface-based 

measurements show that total tropospheric 

bromine from these chemicals peaked in 1998 

and has since declined (Clerbaux and Cunnold et 

al., 2007). By mid-2004, tropospheric bromine 

from these gases was 0.6 to 0.9 ppt below the 

peak amount. The decline was entirely a result 

of declining CH3Br mixing ratios. By 2004 

the tropospheric mean CH3Br mixing ratio 

had declined by 1.3 ppt (14%) from its peak in 

1998. Although the rate of decline of CH3Br was 

variable over this period, global mixing ratios 

of CH3Br have declined each year during this 

period as global production decreased. Global 

atmospheric mixing ratios of the halons were 

still increasing slowly in 2004, albeit at slower 

rates than in earlier years (2004 rates of increase 

for the halons were less than 0.1 ppt per year). 

Continued increases in halon mixing ratios 

arise from continued production allowed in 

developing nations and slow leakage rates from 

large banks of halons in developed countries in 

fire extinguishers that are still in use.

Trends in global mean bromine accounted for 

by short-lived gases are not easily measured 

from ground-based stations, because of the 

high variability observed and the potential 

for local influences to dominate measured 

abundances and changes. In these instances, 

data from firn air (air trapped in uncompacted 

snow in the polar regions) have improved our 

understanding of historical changes in the 

atmospheric abundance of these chemicals. 

Firn air integrates atmospheric abundances 

over decadal periods so short-term variations 

are smoothed, but local inf luences could 

mask broader changes. Atmospheric histories 

of short-lived brominated chemicals such as 

dibromomethane (CH2Br2) and bromoform 

Whereas chlorine 

emissions from uses 

regulated by the 

Montreal Protocol 

accounted for 

approximately 80-85% 

of stratospheric 

chlorine at its 

peak abundance, 

emissions of bromine 

regulated by the 

Protocol accounted 

for approximately 

50% of the bromine 

measured in the 

stratosphere at its 

peak abundance.
The U.S. Climate Change Science Program Chapter 2

66

(CHBr3), derived in this way suggest no large 

long-term changes in atmospheric mixing ratios 

in polar regions during the past two decades 

(Law and Sturges et al., 2007).

As was true for chlorine, the integrated influence 

of changes in the tropospheric abundance 

of brominated gases can be measured in the 

stratosphere from airborne, balloon-borne, 

and satellite instrumentation. These data have 

been important for quantifying the role nonregulated chemicals play in controlling the 

abundance of bromine in the stratosphere. They 

have also demonstrated that the total abundance 

of Br in the stratosphere has increased over 

time in a manner that can be explained by the 

tropospheric mixing ratio changes observed 

for halons and CH3Br, considering lag times 

associated with air transport (Law and Sturges 

et al., 2007). A recent study published since 

WMO (2007) suggests that the accumulation 

rate of bromine in the stratosphere has slowed 

in a manner consistent with the trend observed 

in the troposphere, after considering the time 

it takes to transport air from the troposphere to 

the stratosphere (Dorf et al., 2006).

Tropospheric bromine attributable to U.S. 

emissions of halons and CH3Br also peaked 

in 1998-1999 and declined through 2003, 

as estimated by the U.S. EPA (U.S. EPA, 

2007) (Figure 2.14). Since 2003, however, the 

increased emissions of CH3Br from Critical 

Use Exemptions and QPS uses (Box 2.3) have 

The threat posed to the ozone layer from ODSs is not directly proportional to the summed mixing ratios of these 

chemicals in the troposphere. Instead, it depends upon the number of chlorine and bromine atoms contained in 

the ODSs, how rapidly the ODSs degrade once they reach the stratosphere and liberate ozone-depleting forms of 

chlorine and bromine, and the abundance of bromine relative to chlorine contained in the mix of ODSs reaching 

the stratosphere (given that bromine is 60 times more reactive, on average, than chlorine). To account for these 

influences, indices have been developed to estimate changes in the burden of reactive stratospheric halogen in a 

simple manner based on observed changes in tropospheric abundances of ODSs. 

Equivalent Effective Chlorine (EECl) is one such index used here and elsewhere (WMO, 2007) to quantify overall 

changes in reactive halogen trends based upon the measured mix of ODSs in the troposphere. The timing associated 

with EECl changes correspond to the dates those changes were measured in the troposphere. 

Equivalent Effective Stratospheric Chlorine (EESC) is a closely related index used to estimate the time evolution 

of ozone-depleting halogen in the stratosphere. In most past formulations it has differed from EECl only in that it 

includes a time lag associated with transporting air from the troposphere, where ODS measurements are regularly 

made, to the stratosphere. EESC is often used to estimate when the cumulative effect of all ODSs on ozone 

will return to a level attained at some earlier time, often chosen to be 1980, assuming no changes in dynamical, 

climate, or other non-ODS-related influences (WMO, 2007). Quite different “recovery” times can be calculated 

for midlatitude EESC and Antarctic EESC in springtime when lag times of three years for midlatitudes and six 

years for Antarctica are assumed (see Chapter 5 for additional discussion of recovery times). Furthermore, EESC 

projections for different scenarios of ODS uses have been an important tool for assessing the potential influence 

of various policy choices on ozone. 

Recently, EESC has also been used to improve our understanding of the extent to which changes in ozone 

abundances may be due to policy restrictions under the currently adjusted and amended Montreal Protocol. 

Specifically, attempts have been made to identify both a slowing of the declining ozone trends and even a reversal 

of the decline, and whether these recovery milestones can be attributed to ODS changes.

Despite its usefulness, EESC provides only a rough estimate of the effect of ODSs on stratospheric ozone 

because it incorporates simplified assumptions regarding mixing processes and degradation rates. Recent efforts 

to enhance the formulation perhaps provide a more realistic evolution of stratospheric reactive halogen over 

time and space (Newman et al., 2006, 2007). Other differences in more recent work relating to the calculation 

of EESC suggest some rather large alterations to estimated ODS recovery times and are currently a source of 

uncertainty in this analysis.

BOX 2.7: Equivalent Effective Chlorine (EECl) and 

Equivalent Effective Stratospheric Chlorine (EESC)
67

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

caused the tropospheric abundance of bromine 

attributable to U.S. emissions to increase. 

While the contribution of United States to total 

atmospheric bromine (the industrially derived 

emissions from regulated uses only) declined 

throughout the 1980s and early 1990s, it 

reversed course and actually increased by about 

7% from 2000-2005; the U.S. contribution to 

atmospheric bromine is estimated at between 17 

and 35% in 2005 (only halons and anthropogenic 

CH3Br considered).

2.4.3 Equivalent Effective 

Stratospheric Chlorine and 

Equivalent Effective Chlorine

The combined inf luence of changes in 

chlorinated and brominated ODSs on reactive 

halogen abundances in the stratosphere and 

on stratospheric ozone can be assessed from 

aggregate quantities such as EESC and EECl 

(Box 2.7). These quantities are derived with 

weighting factors applied to tropospheric ODS 

abundances to provide a rough estimate of how 

total reactive halogen abundances are changing 

or will likely change in the stratosphere based 

upon observed trends in tropospheric mixing 

ratios of ODSs. 

EECl from measured global surface mixing 

ratios of regulated ODSs, substitutes, and 

CH3Cl peaked in 1994 and has since declined 

(Figure 2.15). By 2004, EECl had declined by 

277 ppt, or 8-9% from the peak. This decline 

represents about 20% of the decline needed 

for EECl levels to return to their 1980 levels 

(Clerbaux and Cunnold et al., 2007). Most 

of this decline resulted from changes in the 

atmospheric abundance of the shorter-lived 

ODSs: CH3CCl3 and CH3Br.

In 2005 CFCs still contributed the most to 

the atmospheric burden of EECl (45%) from 

all long-lived chlorinated and brominated 

chemicals, including those with large natural 

sources (CH3Cl and CH3Br) (Table 2.2). Other 

chlorinated gases contributed 25%, brominated 

gases contributed 28%, and HCFCs contributed 

2.5% to EECl in 2005 (Table 2.2). These 

percentages include natural contributions to the 

atmospheric abundance of CH3Cl and CH3Br. 

When only anthropogenic contributions to 2005 

EECl are considered, the relative contribution of 

CFCs increases (to 59%), that for chlorocarbons 

Figure 2.15 EECl from measured global mixing ratios of ODSs (blue line) 

and from estimates of ODS mixing ratios attributable to U.S. emissions 

(upper and lower ranges bounded by red lines; see Box 2.6 for discussion of 

U.S. emissions). While all long-lived ODSs, including CH3Cl and the natural 

contribution of CH3Br, are included in global EECl, natural contributions 

are not included in the United States/Global fraction or EECl calculated 

from U.S. emissions alone. Also shown are ranges for the fraction of EECl 

attributable to U.S. emissions (green lines, right-hand axis).

Figure 2.16 Rate of change in global EECl derived from measured global 

atmospheric mixing ratios of ODSs and substitute chemicals (blue line), and 

the rate of change in EECl derived from atmospheric mixing ratios calculated 

from upper and lower bounds on U.S. emissions of ODSs (bounded by red 

lines; see Box 2.6 for discussion of U.S. emissions). High and low U.S. EECl 

estimates express the influence of emission uncertainties on these rates 

(Box 2.6).
The U.S. Climate Change Science Program Chapter 2

68

decreases (to 17%), that for bromocarbons 

decreases (to 20%), and that for CH3Br alone 

becomes 3.1% (see footnote to Table 2.2 for 

grouping definitions).

Declines are also calculated for EECl attributable 

to U.S. emissions of ODSs for regulated 

uses during 1994-2004, though substantial 

uncertainty in atmospheric abundances 

derived from U.S. emissions prevents a precise 

determination of EECl from U.S. emissions 

alone. United States emissions of ODSs for 

regulated uses have accounted for between 15 

and 39% of total EECl from regulated chemicals 

from 1994 to 2004, and between 15 and 36% 

in 2005. 

Despite the added uncertainty U.S. emissions 

before 1985 add to estimates of the U.S. 

contribution to Cl, Br, and EECl in today’s 

atmosphere, these uncertainties have a much 

smaller influence on our understanding of 

changes in these quantities (Figure 2.16). Global 

EECl declined fairly consistently since the 

mid-1990s. EECl from U.S. emissions followed 

global trends until about 2003, when declines 

in U.S. EECl slowed substantially. From 2004 

to 2005, U.S. EECl declines were substantially 

smaller than in earlier years, primarily because 

of the increases in U.S. emissions of brominated 

gases during these years (primarily CH3Br) 

(Figure 2.16).

On uncertainty in the U.S. contribution.

Though estimates of the rate of change in 

EECl attributable to U.S. emissions of ODSs 

and substitutes are much less dependent upon 

pre-1985 emission rates, they do rely on the 

accuracy of emission algorithms of ODSs 

from in-use applications. Such algorithms are 

difficult to verify experimentally on national 

scales (Section 2.2.6).

2.4.3.1 Estimating reactive halogen

tr e n ds f o r the m idl atit u d e

stratosphere

Changes in stratospheric halogen abundance are 

delayed from those observed in the troposphere 

because of the time it takes for air to be 

transported from the troposphere to stratosphere. 

In addition to this time lag, mixing processes 

also influence how tropospheric composition 

changes propagate to the stratosphere. Estimates 

of stratospheric halogen trends in the midlatitude 

stratosphere have been roughly derived with a 

lag of three years on EECl. The EESC calculated 

for the midlatitude stratosphere suggests that by 

2004, the midlatitude stratospheric halogen 

burden had declined by approximately 7% 

from its peak. 

2.4.3.2 Estimating reactive halogen

trends using eesc for the antarctic

stratosphere

Measurements suggest that air found in the 

lower Antarctic stratosphere during the early 

springtime has resided in the stratosphere for 

about six years, or approximately twice as long 

as it takes to transport air from the troposphere 

to the midlatitude stratosphere. This influence, 

combined with the expected slow decline 

in EESC during the twenty-first century 

compared to the relatively fast buildup around 

1980, suggests that it will take 15-20 years 

longer for EESC in Antarctica to fall below the 

1980s levels than it will for midlatitude EESC 

to drop similarly (Daniel and Velders et al., 

2007) (see also Chapter 5). By 2004, EESC 

over Antarctica is estimated to have declined 

from peak levels by only 3%, when estimated 

simply as a six-year lag of EECl. More detailed 

analyses of these projections are currently 

being refined to include mixing effects and a 

better representation of decomposition rates for 

individual ODSs and substitute chemicals (e.g., 

Newman et al., 2006).

2.5 Changes in Radiative

Forcing Arising From

Ozone-Depleting Chemicals

and Substitutes

2.5.1 Changes in Direct 

Radiative Forcing

As previously noted in this chapter, weighting 

emissions by 100-year GWPs allows one to 

compare the integrated radiative forcing from 

various greenhouse gases, including ODSs, 

with the intent to gain insight into the resulting 

climate effects (see Box 2.2). Instantaneous 

radiative forcing is generally calculated by 

multiplying the atmospheric mixing ratios of 

the various GHGs by their radiative efficiencies 

(Section 2.1.1). While the relationship between 

changes in radiative forcing and global average 

temperature vary somewhat among models, 

United States 

emissions of ozonedepleting substances 

have accounted for 

between 15 and 39% 

of total Equivalent 

Effective Chlorine, an 

index used to quantify 

overall changes in 

reactive halogen 

trends, from regulated 

chemicals from 1994 

to 2004, and between 

15 and 36% in 2005.
69

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

radiative forcing remains arguably the best 

simple metric available to compare the direct 

climate effect of greenhouse gas abundances. 

As with “direct” GWPs (Box 2.2), direct 

radiative forcing represents the forcing of GHGs 

due to their own absorption of infrared energy 

and neglects any potential chemical or other 

feedbacks.

The direct, global radiative forcing due to 

ODSs and substitutes reached about 0.34 W 

per m2 in 2005 and was still increasing slowly 

(Figure 2.17). Due to compliance with the 

Montreal Protocol, however, the recent increase 

in radiative forcing was much slower than 

measured in the early 1990s. HFCs contributed 

a noticeable amount to this increase on a global 

scale. Radiative forcing from non-HFC ODSs 

and substitutes changed less than 0.001 W per 

m2 from 2001-2005. For the purpose of putting 

this total direct forcing into perspective, the 

amount of radiative forcing due to CO2, CH4, 

and N2O in 2005 was approximately 1.66 W 

per m2, 0.48 W per m2, and 0.16 W per m2, 

respectively (Forster et al., 2007). While the 

increase in forcing from ODSs has occurred 

relatively rapidly, the decrease will be largely 

limited by the global ODS lifetimes, and will 

occur more slowly. These future projections will 

be further discussed in Chapter 5.

The relative direct forcing contributions of 

classes of ODS chemicals and their replacements 

are shown in Figure 2.18. The CFCs have 

contributed between 79 and 86% of the total 

direct forcing from ODSs since 1980. However, 

over the last decade the fractional contribution 

of the CFCs has declined, as have the fractional 

contributions of CH3CCl3 and CCl4. Increases 

in HCFC abundances, primarily HCFC-22, 

and in HFC abundances are counteracting the 

decline of these other gases. Consideration of 

the indirect forcing due to ODS-induced ozone 

depletion alters this figure somewhat, though 

The direct, global 

radiative forcing due to 

ODSs and substitutes 

reached about 0.34 

W per m2

 in 2005 and 

was still increasing 

slowly. Putting this 

total direct forcing 

into perspective, the 

amount of radiative 

forcing due to CO2, 

CH4, and N2O in 2005 

was approximately 

1.66 W per m2, 0.48 

W per m2, and 0.16 W 

per m2, respectively.

Figure 2.17 Radiative forcing time series arising from changing atmospheric mixing ratios of ODSs and 

their substitutes. Direct radiative forcing is calculated by weighting global atmospheric mixing ratios of 

ODSs or a range of ODS mixing ratios attributable to the U.S. (see Box 2.6) by compound-dependent 

radiative efficiencies for the years 1980-2005. The additional forcing contribution of HFCs to global or U.S. 

radiative forcing is indicated by the lighter-colored lines. The uncertainties indicated for 2005 include the 

range of influences (and uncertainty) of the indirect forcing associated with stratospheric ozone depletion. 

This indirect influence was estimated for the U.S. by considering the range of EECl attributable to the U.S. 

relative to total global EECl in 2005 (see Box 2.2).
The U.S. Climate Change Science Program Chapter 2

70

in 2005, for example, the percent contribution 

of each compound class changes by less than 

4% when a mean radiative ozone influence 

of −0.05 W per m2 is considered (derived by 

scaling ozone radiative forcing of −0.05 W per 

m2 by the fraction of EECl contributed by each 

compound class).

2.5.1.1 Estimating the u.s. contribution

to direc t r a d i at i v e f o rc i n g

changes

Using estimates of atmospheric mixing ratios 

of ODSs arising solely from U.S. emissions 

(Section 2.3), the U.S. contribution to the 

direct global radiative forcing from ODSs and 

substitutes is calculated to have been between 

0.068 and 0.16 W per m2 in 2005 (Figure 2.17). 

The error bars are calculated by summing the 

individual high and low direct forcing estimates 

for each of the ODSs. Since 1990, the U.S. 

contribution has accounted for between 19 and 

49% of the global direct forcing from ODSs and 

substitute chemicals. The contributions of the 

various ODS classes to radiative forcing show 

the same qualitative behavior for U.S. emissions 

as for the global emissions that are apparent in 

Figure 2.18; HCFC contributions are increasing 

while CFC, CH3CCl3, and CCl4 contributions 

Since 1990, the 

U.S. contribution 

has accounted for 

between 19 and 

49% of the global 

direct forcing from 

ozone-depleting 

substances and 

substitute chemicals.

are decreasing. We estimate that the United 

States accounted for about 45% of the increase 

in direct radiative forcing arising from HCFCs 

during 2000 to 2006.

2.5.2 Changes in Net Radiative Forcing 

As stated in Box 2.2, the direct quantities 

represent the effect of the ODSs themselves 

on radiative forcing through their absorption 

of infrared energy. Their destruction of 

stratospheric ozone leads to an additional 

radiative effect, referred to as an “indirect” 

effect of the ODSs. 

In the past, it had been thought that the indirect 

forcing might have offset as much as 50% of 

total direct halocarbon radiative forcing, but 

estimates from different studies varied widely 

(Ramaswamy et al., 2001). More recently, 

now that ozone trends are thought to be better 

quantified, particularly in the radiatively 

important region near the tropopause, the 

magnitude of the ozone forcing has been revised 

to −0.05±0.05 W per m2

 for the changes in 

stratospheric ozone between 1979 and 1998 

and −0.05±0.10 W per m2 for the changes in 

stratospheric ozone from preindustrial times 

to 2005 (Forster et al., 2007). Although revised 

Figure 2.18 Relative contributions of ODS compound classes to global direct radiative forcing. Percentages in 2005 

are CFCs=80%, HCFCs = 12%, other ODSs = 5%, HFCs = 4%.
71

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

to a smaller value, the ozone forcing remains 

potentially significant relative to the direct 

forcing and continues to be highly uncertain, 

with even the sign in doubt (Figure 2.17; see 

Box 2.2). However, it is known that this indirect 

forcing will gradually become negligible as 

ozone recovers from depletion due to ODSs 

(Daniel et al., 1995; see Chapter 5). 

The IPCC (2007) estimate for the magnitude of 

the ozone forcing through 2005 (−0.05±0.10 W 

per m2) is added to the global direct ODS forcing 

in Figure 2.17. As discussed in Box 2.2, there are 

potential problems related to adding these direct 

and indirect forcing quantities; nevertheless, 

such an addition does permit a comparison of 

the sizes of these globally averaged forcings. 

Estimates for the U.S. contribution to the 

indirect forcing are also added to the U.S. direct 

ODS forcing curves. The U.S. indirect forcing 

has been calculated by scaling the global ozone 

forcing range by the U.S. contribution to global 

EECl. For both the global and U.S. cases, the 

indirect forcing remains potentially significant 

but quite uncertain compared to the direct 

radiative forcing estimates. 

2.6 Summary of Findings

Related to the Role

of the United States

in Influencing Past

Changes in Production, 

Consumption, Emissions, 

and Mixing Ratios

of Ozone-Depleting 

Substances and Their

Substitutes

Stratospheric ozone depletion is a global issue 

because the amount of ozone depletion above 

the United States, or any other location, results 

from the global emission of ODSs. A reduction 

in U.S. ODS emissions leads to reduced 

ozone depletion above the United States only 

to the extent to which it reduces global ODS 

emissions.

In response to restrictions put into place under 

the Montreal Protocol, U.S. consumption of 

ODSs for regulated dispersive uses, considered 

in sum, have declined substantially from peak 

levels. By 2005, ODP-weighted consumption 

in the United States had declined by 97-98%, 

or nearly 400 ODP-Kt since the late 1980s 

(UNEP, 2007). These data indicate that the 

United States accounted for 24 (±2)% of total 

global production of ODSs during the years 

of substantial production (1986-1994), and 10 

(±2)%, on average, during 2001-2005 (when 

weighted by ODPs).

United States consumption declines have been 

slightly faster than phase-out schedules for all 

developed countries in the adjusted and amended 

Montreal Protocol for most ODSs. Critical Use 

Exemptions have resulted in ODS consumption 

for CH3Br and CFCs above these scheduled 

allotments in recent years. For example, though 

CH3Br consumption in developed countries 

was to have been zero in 2005 and thereafter, 

U.S. consumption for critical uses was 28% of 

1986 baseline consumption during this year 

(4.4 ODP-Kt) and, in 2006, was 27% of 1986 

baseline consumption (4.1 ODP-Kt) (UNEP/

MBTOC, 2007). Authorized CUEs for CH3Br 

consumption in the United States for 2007 were 

reduced compared to 2005 and 2006 (17% of 

the U.S. 1986 baseline consumption). U.S. 

consumption also has continued for critical uses 

of CFCs despite the 1996 phase-out, though this 

consumption has been comparably small (less 

than 1% of 1986 United States baseline CFC 

consumption in all years since 1996).

The decreases in global and U.S. production and 

consumption have led to substantial reductions 

in emissions of most ODSs to the atmosphere. 

But while global emissions can be derived 

from observed global atmospheric changes and 

knowledge of ODS atmospheric lifetimes, U.S. 

emissions and their changes are more uncertain, 

as they were derived from vintaging model 

analyses of sales, use, and release patterns 

of ODSs (a “bottom-up” analysis; U.S. EPA, 

2007). The results suggest that both global and 

U.S. emissions of ODSs declined overall by 

81-82% since the late 1980s when weighted by 

ODPs (Table 2.1). Furthermore, this analysis 

suggests that the United States accounted for 

a decreasing amount of global ODS emissions, 

from a peak of 35% in 1993 to 20 (±2)% during 

2000-2005 (Figure 2.7).

Similar relative contributions of U.S. emissions 

in 2005 to global atmospheric abundances 

are calculated for the atmospheric abundance 
The U.S. Climate Change Science Program Chapter 2

72

of chlorine (17-42%), bromine (17-35%), and 

EECl (15-36%) during 2000-2005. EECl arising 

from U.S. emissions declined every year from 

1994 through 2004, but did so much more 

slowly during 2004 to 2005, largely due to 

the recent increase in U.S. CH3Br emissions. 

U.S. emissions have also resulted in the U.S. 

accounting for 19-49% of global direct radiative 

forcing from ODSs in the 2000s.

Weighting ODS emissions by 100-year, direct 

GWPs allows the magnitude of these emissions 

to be compared to those of CO2 to approximate 

their direct climate effects. Declines of 77 

(77-84)% and 74 (71-75)% are calculated 

for annual global and U.S. GWP-weighted 

emissions of ODS and substitute chemicals 

through 2004, or a decline in annual emissions 

on a global scale of 7270 (3600-8500) Mt CO2-

equivalents and a decline in annual emissions 

in the United States of 1640 (900-1880) Mt 

CO2-equivalents (quantities in parenthesis are 

calculated with consideration of net GWPs 

that include radiative forcing and uncertainty 

associated with stratospheric ozone changes; 

see Box 2.2). The U.S. decline alone is likely a 

large fraction of the global benefit anticipated 

as a result of adherence the Kyoto Protocol 

(~2000 Mt CO2-equivalent emissions; Velders 

et al., 2007).
73

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

Appendix 2a

Table 2A.1 Lifetimes, relative fractional halogen release factor, and Ozone Depletion Potentials 

for halocarbons. Reproduced from Daniel and Velders et al. (2007).

Halocarbon Lifetime

(years)

Relative Fractional

Release Factora

Semi-Empirical

ODP

ODP in Montreal

Protocol

Annex A-1

CFC-11 45 1 1.0 1.0

CFC-12 100 0.60 1.0 1.0

CFC-113 85 0.75 1.0 0.8

CFC-114 300 0.28±0.02b 1.0 1.0

CFC-115 1700 0.44 † 0.6

Annex A-11

Halon-1301 65 0.62 16 10.0

Halon-1211 16 1.18 7.1c 3.0

Halon-2402 20 1.22 11.5 6.0

Annex B-11

Carbon tetrachloride 26 1.06 0.73 1.1

Annex B-111

Methyl chloroform 5.0 1.08 0.12 0.1

Annex C-1

HCFC-22 12.0 0.35 0.05 0.055

HCFC-123 1.3 1.11 0.02 0.02

HCFC-124 5.8 0.52 0.02 0.022

HCFC-141b 9.3 0.72 0.12 0.11

HCFC-142b 17.9 0.36 0.07 0.065

HCFC-225ca 1.9 1.1 0.02 0.025

HCFC-225cb 5.8 0.5 0.03 0.033

Annex E

Methyl bromide 0.7 1.12 0.51 0.6

Others

Halon-1202 2.9 1.7d

Methyl chloride 1.0 0.80 0.02

†Model-derived values, WMO (2003).

a From WMO (2003), Table 1-4, except for the value for CFC-114. For the EESC calculations in Section 1.8 of WMO 

(2003), slightly different relative fractional release factors were used by mistake for the halons. 

b From Schauffler et al., 2003. c The ODP of halon-1211 should have been reported as 5.3 in the previous Assessment (WMO, 2003), but was incorrectly reported as 6.0 due to a calculation error.

d WMO (2003), with adjustment for updated α value.
The U.S. Climate Change Science Program Chapter 2

74

Industrial Designation(s) or 

Common Name Chemical Formula

Radiative 

Efficiencyb

(W per 

square m per 

ppbv)

Lifetime 

(years)

Global Warming Potential for 

Given Time Horizon

20 years 100 

years

500 

years

Carbon dioxide CO2 1.41 × 10-5 c 1 1 1

Nitrous oxide N2O 3.03 × 10-3 114 d 289 298 153

Chlorofluorocarbons

CFC-11 CCl3F 0.25 45 6,730 4,750 1,620

CFC-12 CCl2F2 0.32 100 10,990 10,890 5,200

CFC-13 CClF3 0.25 640 10,800 14,420 16,430

CFC-113 CCl2FCClF2 0.30 85 6,540 6,130 2,690

CFC-114 CClF2CClF2 0.31 300 8,040 10,040 8,730

CFC-115 CClF2CF3 0.18 1700 5,310 7,370 9,990

Hydrochlorofluorocarbons

HCFC-21 CHCl2F 0.14 1.7 530 151 46

HCFC-22 CHClF2 0.20 12.0 5,160 1,810 549

HCFC-123 CHCl2CF3 0.14 1.3 273 77 24

HCFC-124 CHClFCF3 0.22 5.8 2,070 609 185

HCFC-141b CH3CCl2F 0.14 9.3 2,250 725 220

HCFC-142b CH3CClF2 0.20 17.9 5,490 2,310 705

HCFC-225ca CHCl2CF2CF3 0.20 1.9 429 122 37

HCFC-225cb CHClFCF2CClF2 0.32 5.8 2,030 595 181

Hydrofluorocarbons

HFC-23 CHF3 0.19e 270 11,990 14,760 12,230

HFC-32 CH2F2 0.11e 4.9 2,330 675 205

HFC-41 CH3F 0.02 2.4 323 92 28

HFC-125 CHF2CF3 0.23 29 6,340 3,500 1,100

HFC-134 CHF2CHF2 0.18 9.6 3,400 1,100 335

HFC-134a CH2FCF3 0.16e 14.0 3,830 1,430 435

HFC-143 CH2FCHF2 0.13 3.5 1,240 353 107

HFC-143a CH3CF3 0.13 52 5,890 4,470 1,590

HFC-152 CH2FCH2F 0.09 0.60 187 53 16

HFC-152a CH3CHF2 0.09 1.4 437 124 38

HFC-227ea CF3CHFCF3 0.26e 34.2 5,310 3,220 1,040

HFC-236cb CH2FCF2CF3 0.23 13.6 3,630 1,340 407

HFC-236ea CHF2CHFCF3 0.30 10.7 4,090 1,370 418

HFC-236fa CF3CH2CF3 0.28 240 8,100 9,810 7,660

HFC-245ca CH2FCF2CHF2 0.23 6.2 2,340 693 211

HFC-245fa CHF2CH2CF3 0.28 7.6 3,380 1,030 314

HFC-365mfc CH3CF2CH2CF3 0.21 8.6 2,520 794 241

HFC-43-10mee CF3CHFCHFCF2CF3 0.40 15.9 4,140 1,640 499

Table 2A.2 Direct Global Warming Potentials for selected gases. Reproduced from Daniel and Velders et al. (2007).
75

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

Chlorocarbons

Methyl chloroform CH3CCl3 0.06 5.0 506 146 45

Carbon tetrachloride CCl4 0.13 26 2,700 1,400 435

Methyl chloride CH3Cl 0.01 1.0 45 13 4

Bromocarbons

Methyl bromide CH3Br 0.01 0.7 17 5 1

Halon-1201 CHBrF2 0.14 5.8 1,380 404 123

Halon-1211 CBrClF2 0.30 16 4,750 1,890 574

Halon-1301 CBrF3 0.32 65 8,480 7,140 2,760

Halon-2402 CBrF2CBrF2 0.33 20 3,680 1,640 503

Fully fluorinated species

Sulfur hexafluoride SF6 0.52 3200 16,260 22,810 32,600

Trifluoromethylsulfur-pentafluoroide SF5CF3 0.57 650-950

13,120-

13,180

17,540-

17,960

20,060

-22,360

Perfluoromethane CF4 0.10e 50000 5,210 7,390 11,190

Perfluoroethane C2F6 0.26 10000 8,620 12,200 18,180

Perfluoropropane C3F8 0.26 2600 6,310 8,830 12,450

Perfluorobutane C4F10 0.33 2600 6,330 8,850 12,480

Perfluorocyclobutane c-C4F8 0.32 3200 7,310 10,250 14,660

Perfluoropentane C5F12 0.41 4100 6,510 9,150 13,260

Perfluorohexane C6F14 0.49 3200 6,620 9,290 13,280

Perfluorodecalin C10F18 0.56f 1000 5,500 7,510 9,440

Halogenated alcohols and ethers

HFE-125 CHF2OCF3 0.44 136 13,790 14,910 8,490

HFE-134 CHF2OCHF2 0.45 26 12,190 6,320 1,960

HFE-143a CH3OCF3 0.27 4.3 2,630 756 230

HCFE-235da2 CHF2OCHClCF3 0.38 2.6 1,230 349 106

HFE-245fa2 CHF2OCH2CF3 0.31 4.9 2,280 659 200

HFE-254cb2 CH3OCF2CHF2 0.28 2.6 1,260 359 109

HFE-7100 (HFE-44-9) CH3OC4F9 0.31 5.0 1,390 404 123

HFE-7200 (HFE-56-9) C2H5OC4F9 0.30 0.77 200 57 17

HFE-245cb2 CH3OCF2CF3 0.32 5.1 2,440 708 215

HFE-347mcc3 CH3OCF2CF2CF3 0.34 5.2 1,980 575 175

HFE-356pcc3 CH3OCF2CF2CHF2 0.33 0.93 386 110 33

HFE-374pc2 CH3CH2OCF2CHF2 0.25 5.0 1,930 557 169

CH3OCF(CF3)2 0.31 3.4 1,200 343 104

Industrial Designation(s) or 

Common Name Chemical Formula

Radiative 

Efficiencyb

(W per 

square m per 

ppbv)

Lifetime 

(years)

Global Warming Potential for 

Given Time Horizon

20 years 100 

years

500 

years
The U.S. Climate Change Science Program Chapter 2

76

HFE-43-10pccc124 a

CHF2OCF2OC2F4OCHF2 1.37 6.3 6,320 1,870 569

(CF3)2CHOH 0.28 2.0 764 217 66

HFE-236ca12 CHF2OCF2OCHF2 0.66 12.1 8,040 2,820 859

HFE-338pcc13 

CHF2OCF2CF2OCHF2 0.87 6.2 5,070 1,500 456

Species whose lifetimes have a high

 uncertainty

Nitrogen trifluoride NF3 0.21e 740 13,370 18,000 21,270

Perfluorocyclopropane c-C3F6 0.42 >1000 >12,700 >17,340 >21,800

HFE-227ea CF3CHFOCF3 0.40 11 4,540 1,540 468

HFE-236ea2 CHF2OCHFCF3 0.44 5.8 3,370 989 301

HFE-236fa CF3CH2OCF3 0.34 3.7 1,710 487 148

HFE-245fa1 CHF2CH2OCF3 0.30 2.2 1,010 286 87

HFE-329mcc2 CHF2CF2OCF2CF3 0.49 6.8 3,060 919 279

HFE-338mcf2 CF3CH2OCF2CF3 0.43 4.3 1,920 552 168

HFE-347mcf2 CHF2CH2OCF2CF3 0.41 2.8 1,310 374 114

HFE-356mec3 CH3OCF2CHFCF3 0.30 0.94 355 101 31

HFE-356pcf2 CHF2CH72OCF2CHF2 0.37 2.0 931 265 80

HFE-356pcf3 CHF2OCH2CF2CHF2 0.39 3.6 1,760 502 153

CHF2OCH(CF3)2 0.41 3.1 1,330 379 115

-(CF2)4CH(OH)- 0.30 0.85 254 72 22

Note: Values are calculated for a CO2 mixing ratio of 378 ppm, compared with 370 ppm in IPCC/TEAP (2005) and WMO (2003), which 

tends to increase all GWPs.

______________

a Referred to as H-Galden 1040x in WMO/UNEP ozone assessments prior to WMO (2007).

b All values not otherwise noted from IPCC/TEAP (2005).

c See Section 8.2.3 of Daniel and Velders et al., 2007. d This value is an adjustment time that includes feedbacks of emissions on the lifetime.

e See Table 8-3 of Daniel and Velders et al., 2007. f From Shine et al., 2005.

Industrial Designation(s) or 

Common Name Chemical Formula

Radiative 

Efficiencyb

(W per 

square m per 

ppbv)

Lifetime 

(years)

Global Warming Potential for 

Given Time Horizon

20 years 100 

years

500 

years
77

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

a

 Uncertainties associated with direct GWPs are estimated to be 35% at the 2σ level. b

 Indirect GWPs are affected by numerous uncertainties as discussed in the text. Here, the only contribution 

to the indirect GWP range considered is uncertainty in the ozone forcing through 2005 (–0.05±0.1 W per 

m2

; IPCC, 2007). The ‘Low’ GWP assumes an ozone forcing of −0.15 W per m2, while the ‘High’ is calculated 

for +0.05 W per m2

. Both indirect values are calculated by scaling the indirect GWPs in WMO (2007) by 

the ratio of the appropriate ozone forcing used here to the forcing assumed in WMO (2007), which was 

−0.15 W per m2

.

Gas Direct GWPa Indirect GWPb

Low High

CFC-11 4,750 −3,790 1,263

CFC-12 10,890 −2,160 720

CFC-113 6,130 −2,530 843

CH3CCl3 146 −643 214

CCl4 1,400 −3,630 1,210

HCFC-22 1,810 −286 95

HCFC-123 77 −83 28

HCFC-124 609 −120 40

HCFC-141b 725 −667 222

HCFC-142b 2,310 −362 121

HCFC-225ca 122 −93 31

HCFC-225cb 595 −156 52

CH3Br 5 −2,150 717

Halon 1211 1,890 −40,280 13,430

Halon 1301 7,140 −49,090 16,360

Halon 2402 1,640 −62,000 20,670

Table 2A.3 Direct and indirect GWPs for a 100-year time horizon
The U.S. Climate Change Science Program Chapter 2

78
CHAPTER

3

79

Ozone and UV Observations

Convening Lead Authors: Paul A. Newman, NASA; 

Jay R. Herman, NASA

Lead Authors: Richard Bevilacqua, NRL; Richard Stolarski, NASA; 

Terry Keating, U.S. EPA

KEY ISSUES

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

As atmospheric concentrations of ozone-depleting substances change as a result of implementation of international policies, concentrations of stratospheric ozone and levels of ultraviolet 

radiation reaching the Earth’s surface should also change. However, ozone concentrations and 

ultraviolet (UV) levels are affected by other natural and human processes as well. To understand whether the international policies are working, we must be able to determine changes 

in stratospheric ozone and ground-level UV and separate out the effects of ozone-depleting 

substances (ODS) changes and the effects of other factors.

Stratospheric ozone depletion is a global problem that has its most profound effects in the polar 

regions. However, the processes that drive stratospheric ozone depletion in the polar regions are 

somewhat different than those that drive depletion in the rest of world. Therefore, the impact 

of ODS changes may be different in polar regions than over the midlatitude United States.

In this chapter, we briefly review the observations and current understanding and uncertainties in long-term trends in atmospheric ozone and ground-level UV radiation to address the 

following questions: 

• What is the current state of ozone in the stratosphere in the Earth’s midlatitudes and over 

the polar regions? 

• What do the observations indicate about the abundances and trends of stratospheric ozone 

over the United States and elsewhere?

• How do midlatitude ozone levels and the processes that drive them differ from ozone levels 

and driving processes in the polar regions?

• What is the trend in the occurrence, depth, duration, and extent of the Antarctic ozone 

hole?

• What is the state of stratospheric ozone depletion in the Arctic region?

• How well do we understand the chemical and meteorological processes that determine 

stratospheric ozone concentrations in the polar regions and midlatitudes?

• How have UV radiation levels at the Earth’s surface in the United States and elsewhere 

changed as a result of changes in stratospheric ozone?
The U.S. Climate Change Science Program Chapter 3

80

• Total global ozone has remained relatively constant over the last four years (2002-2006). 

Northern midlatitude ozone reached a minimum in 1993 because of forcings from the Mt. 

Pinatubo eruption and the solar cycle minimum, and has increased somewhat since then. 

Southern midlatitude ozone decreased until the late 1990s, and has been constant since. 

There are no significant ozone trends over the tropics.

• Ozone over the continental United States has followed the behavior of ozone for the entire 

northern midlatitude region; a decrease to a minimum in 1993, and an increase since then 

(see previous bullet). 

• Ozone depletion in the upper stratosphere has closely followed the trends in chlorine. The 

slow down of the negative (or decreasing) trend is attributed to the leveling off of chlorine 

in this region of the stratosphere.

• Over the last decade (1995-2006), the Antarctic ozone hole has not worsened. Most Antarctic ozone hole diagnostics show losses leveling off after the mid-1990s. Saturation of 

ozone loss inside the ozone hole due to complete ozone destruction over a broad vertical 

layer plays the major role in this leveling off. This complete ozone destruction over a deep 

vertical layer is modulated by year-to-year dynamical variations. Antarctic ozone hole diagnostics showed an increase of ozone levels in some recent winter years (e.g., 2002, 2004), 

but these increases resulted from higher levels of dynamical forcing which warmed the 

Antarctic stratosphere, and not decreases in Equivalent Effective Stratospheric Chlorine 

levels. In contrast, the Austral spring of 2006 had below average dynamical forcing resulting in below average Antarctic temperatures, causing the 2006 Antarctic ozone hole to be 

one of the largest on record.

• Arctic spring total ozone values over the last decade were lower than values observed in 

the 1980s. In addition, spring Arctic ozone is highly variable depending on dynamical conditions. For current halogen levels, anthropogenic chemical loss and variability in ozone 

transport are about equally important for year-to-year Arctic ozone variability. Colderthan-average vortex conditions result in larger halogen-driven chemical ozone losses. 

Warmer-than-average vortex conditions result in smaller halogen-driven chemical ozone 

losses. Variability of temperatures and ozone transport are correlated because they are 

both driven by dynamic variability.

• Erythemal irradiance (which is a weighted combination of UVA and UVB based on skin 

sensitivity) over the United States increased roughly by 7% when the ozone minimum was 

reached in 1993 and is now about 4% higher than in 1979. 

• Ground-based measurements of UV irradiance can detect UV trends related to ozone change 

when data from only days with clear sky are used by correcting for aerosol scattering and 

absorption using measured aerosol data.

• UV irradiance estimated from satellite data are usually 10% to 30% too high because satellite 

algorithms neglect the effects of absorbing aerosols. 

• UVB irradiance trends can be estimated directly from satellite-measured ozone changes 

since regional cloud cover and aerosol loadings have not undergone large changes since 

1979 except for a short period after the June 1991 Mt. Pinatubo eruption.

• Increased adverse human health effects associated with excessive UV exposure have been 

observed in Australia, where there are lower ozone amounts and less cloud cover, compared 

with similar latitudes in the United States. 

KEY FINDINGS
81

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

Ozone is destroyed when it reacts with oxides 

of nitrogen, hydrogen, chlorine, bromine, or 

oxygen atoms in catalytic reactions to reform 

molecular oxygen. 

3.1 INTRODUCTION 

Ozone is a trace constituent of the atmosphere, 

with maximum volume mixing ratios of about 

10 to 12 molecules per million air molecules 

(i.e., 10 to 12 ppm). Figure 3.1 (top) shows the 

annually averaged, longitude-averaged ozone 

distribution. 

The total amount of ozone (i.e., the vertical 

integral of ozone density from the surface to 

space) is highest in the mid-to-high latitudes. 

The bottom panel of Figure 3.1 shows the 

total ozone integrated from the top panel. In 

midlatitudes, ozone density is highest in the 

lower stratosphere between 12 and 25 km 

(Figure 3.1 middle panel). While the maximum 

of the ozone mixing ratio (Figure 3.1 top panel) 

is highest in the tropics at 32 km, the total 

column ozone is highest in the midlatitudes, not 

in the tropics (illustrated in the bottom panel). 

The distribution of ozone mixing ratios (Figure 

3.1 top), density (Figure 3.1 middle), and 

total ozone (Figure 3.1 bottom) is controlled 

by the photochemical production, catalytic 

destruction, and transport. The basic circulation 

(shown as the yellow streamlines in the upper 

two panels of Figure 3.1) is known as the 

Brewer-Dobson circulation (Shepherd, 2007). 

This Brewer-Dobson circulation carries air 

into the stratosphere in the tropics near 16 km, 

leading to very low ozone in the tropical lower 

stratosphere as the low ozone air is carried 

upward from the troposphere. The poleward 

and downward flow of ozone from the tropics 

produces the midlatitude maximum in both 

hemispheres.

As air rises in the tropical stratosphere, ozone 

is produced when molecular oxygen (O2) is 

split by solar ultraviolet radiation (hν) to form 

oxygen atoms (O) that combine with O2 in the 

presence of a third air molecule (M) to form 

O3.

O2 + hν → O + O (1a)

 2 (O + O2 + M → O3 + M) (1b)

 Net: 3 O2 + hν → 2 O3

This solar production of ozone leads to very high 

ozone concentrations in the mid-stratosphere in 

the tropics (near 32 km). 

Figure 3.1 Annual longitudinal averaged ozone mixing ratios (top), 

ozone density (middle), and annual longitudinal averaged total ozone 

(bottom). Top panel units are parts per million (ppm); middle panel 

units are molecules per cm3

; bottom panel units are Dobson Units 

(DU). One DU is equal to a column amount of 2.69x1016 molecules 

per cm2

 or about 0.01 mm of pure ozone at standard temperature and 

pressure. Equivalently, an ozone number density of 1012 molecules per 

cm3

 is equivalent to 3.7 DU per kilometer (km). The bottom panel is 

the vertical integral of the middle panel. The annual average flow field 

stream lines are shown in the top and middle panels. The rising motion 

in the tropical stratosphere and sinking motion in the polar region is 

known as the Brewer-Dobson circulation. Adapted from McPeters et 

al. (2007).
The U.S. Climate Change Science Program Chapter 3

82

bromine oxides to reservoir compounds such 

as hydrochloric acid (HCl), chlorine nitrate, 

(ClONO2), and bromine nitrate (BrONO2). 

These chlorine and bromine species are 

eventually returned to the troposphere, where 

they are removed in wet processes.

These ozone catalytic cycles involve oxygen 

atoms (O), and thus operate most rapidly in 

the mid-stratosphere of the tropics and the 

midlatitudes, where the concentration of oxygen 

atoms increases with increasing altitude. 

Oxygen atom concentrations increase with 

altitude because their loss slows as the density 

of O2 and M (O + O2 + M → O3 + M) decreases 

with altitude. Maximum halogen-catalyzed 

ozone loss at midlatitudes occurs around an 

altitude of about 40 km (just above the peak 

ozone concentrations), where these oxygen 

atoms are more abundant. While fractional 

ozone loss peaks near 40 km for a stratosphere 

unperturbed by cold temperatures (about 8 

to 10% of the naturally-occurring ozone at 

that altitude), the contribution of ozone loss at 

40km to the fractional loss in the total column 

is small, since ozone density falls off rapidly 

above the 20-25 km layer (Figure 3.1, middle 

panel).

Ozone depletion in the polar lower stratosphere 

involves different chemistry than described 

above. During winter, the lower stratosphere 

over the poles is characterized by air that the 

Brewer-Dobson circulation has carried poleward 

and downward from the upper stratosphere and 

mesosphere (Figure 3.1, top panel), extremely 

low temperatures (less than 200 K (−73°C)), 

and a circumpolar jet stream that isolates the 

air over the polar regions from midlatitude 

influence (the polar vortex). These extremely 

cold and isolated conditions enable polar 

stratospheric clouds (PSCs) to form (Crutzen 

and Arnold, 1986; Toon et al., 1986). The ozone 

loss occurs in two steps. First, heterogeneous 

chemical reactions occur on the surfaces of the 

PSC particles, liberating chlorine from the two 

reservoir species (HCl + ClONO2 [on PSCs]→

Cl2 + HNO3) (Solomon et al., 1986). Second, 

two principal chlorine and bromine catalytic 

reactions that do not involve oxygen reactions 

(Equation 1a) produce rapid depletion:

O3 + X → O2 + XO (2a)

 XO + O → O2 + X (2b)

 O3 + hν → O2 + O (2c)

 Net: 2 O3 + hν → 3 O2

Here, X represents the catalysts chlorine 

atoms (Cl), bromine atoms (Br), and the 

oxides of nitrogen (nitric oxide, NO) and 

hydrogen (hydroxyl, OH), while hν represents 

the absorption of solar ultraviolet light to 

photochemically break a chemical bond of 

ozone. The net effect of the catalytic cycle is to 

destroy two ozone molecules while regenerating 

the catalytic agent. All of these catalysts are 

highly reactive free radicals, meaning they have 

an unpaired electron, which tends to attach to 

other molecules in order to form a chemical 

bond. Since these reactions have an initial 

energy barrier to reaction, warmer temperatures 

will speed up this catalytic cycle, and cooler 

temperatures (as predicted to occur by recent 

climate models) will slow down this ozone loss 

cycle. In Figure 3.1, ozone decreases above 32 

km as this ozone destruction begins to dominate 

over the ozone production. 

The source gases for the ozone-destroying 

catalysts are compounds such as chlorofluorocarbons, CFCs (chlorine), halons and methyl 

bromide (bromine), nitrous oxide (nitrogen), 

and methane (hydrogen) (see Chapter 2 for a 

complete discussion of these source gases). The 

relative contributions of the oxides of hydrogen 

(HOx), oxides of chlorine (ClOx), and oxides 

of nitrogen (NOx) reactions can be found in 

Figure 1.11 of IPCC/TEAP (2005). As the air 

rises in the stratosphere, the catalytic agents are 

liberated from the source gases by both the UV 

radiation and chemical reactions. 

The catalytic reactions that cause stratospheric 

ozone decreases are principally those involving 

chlorine and bromine. These chlorine and 

bromine compounds are from halogen species 

such as CFCs and halons. These species are 

inert in the troposphere, but are carried into 

the stratosphere by the slow rising circulation 

(Figure 3.1, top panel). As they ascend in the 

stratosphere, the halogen species are broken 

down by UV radiation or oxidation, releasing 

chlorine and bromine to catalytically destroy 

ozone. The rate of catalytic destruction of ozone 

is limited by the conversion of the chorine and 

The catalytic 

reactions that 

cause stratospheric 

ozone decreases 

are principally those 

involving chlorine 

and bromine.
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

at an earlier time in spring than the Antarctic, 

shutting off the ozone loss. 

In the midlatitudes, ozone destruction can 

take place locally or ozone-depleted air may 

be transported from polar regions. During 

periods following major volcanic eruptions, the 

sulfur injected into the stratosphere can lead 

to enhanced aerosols in the lower stratosphere. 

The surfaces of these aerosols promote the 

conversion of reservoir compounds of chlorine 

and bromine back to catalytically active oxides 

that increase ozone destruction.

The solar UV radiation that reaches the Earth’s 

surface is strongly screened by ozone. The UV 

radiation important for biological processes 

is described by two bands, UVA (315 to 400 

nanometers, nm) and UVB (280 to 315 nm). In a 

cloud-free atmosphere, both UVA and UVB are 

scattered by both molecules (Rayleigh scattering) 

and aerosols, while UVB is also significantly 

absorbed by ozone. Ozone absorption increases 

rapidly with decreasing wavelength, which is 

why there is little detectable radiation below 

280 nm at the Earth’s surface. For a given sun 

angle, the relationship of percent UV increase 

to percent ozone decrease is proportional to 

the ozone absorption. Human exposure to UV 

radiation has both negative (e.g., skin cancer 

and eye cataracts) and positive (e.g., Vitamin 

D production) effects. The negative effects 

of UV overexposure are the major reason for 

concern over stratospheric ozone decreases. 

In addition to changes in ozone, long-term 

changes in the amount of aerosols and cloud 

cover affect exposure at the surface to all UV 

wavelengths. 

The following sections of this chapter briefly 

review the observed trends in ozone and ground 

ultraviolet radiation levels and discuss our 

current understanding of the processes that 

determine these levels. For each of these issues, 

the polar regions will be discussed separately 

from the low and midlatitudes because of the 

fundamentally different issues associated with 

those regions.

 ClO + ClO + M → ClOOCl + M (3a)

 ClOOCl + hν → 2 Cl + O2 (3b)

 2 [ Cl + O3 → ClO + O2 ] (3c)

 Net: 2 O3 + hν → 3 O2

BrO + ClO + hν →Br + Cl + O2 (4a)

 Br + O3 → BrO + O2 (4b)

 Cl + O3 → ClO + O2 see (3c)

 Net: 2 O3 + hν → 3 O2

Equation (4a) represents a sequence of reactions 

that together lead to the products shown. Again, 

hν represents the absorption of solar light to 

photochemically break the chemical bonds, 

and M represents any air molecule, typically 

nitrogen (N2) or oxygen (O2), which carries away 

the excess energy of the reaction. In contrast to 

the intense UV necessary to photolyze oxygen 

molecules in (1a), the reactions (3b and 4a) 

require only visible light. These two catalytic 

cycles account for all but a few percent of the 

polar ozone loss, which occurs in the lowermost 

stratosphere (12 to 24 km altitude). This effect is 

strongest in the Antarctic stratosphere where the 

stable polar vortex allows the nearly complete 

destruction of ozone between about 12 and 22 

km altitude each spring, forming the Antarctic 

ozone hole (see the low ozone amounts in Figure 

3.1, bottom panel). The principal ingredients for 

large ozone losses in the polar regions are: (1) 

cold temperatures (< 195 K or −78°C) for the 

formation of PSCs; (2) high concentrations of 

chlorine and bromine; and (3) visible light for 

photolyzing both molecular chlorine (Cl2) and 

the ClO dimer (ClOOCl).

The dramatic seasonal ozone losses occur over 

Antarctica during the Austral spring AugustOctober period (with more than 50% of the 

total column ozone depleted) and to a smaller 

extent over the Arctic during the Boreal spring 

February-March period. The difference in 

hemispheres has to do with the contrast between 

the presence of polar stratospheric clouds and 

the timing of the break up of the polar vortex 

in the two polar regions. First, PSC extent is 

much greater in the Antarctic due to colder 

stratospheric temperatures than in the Arctic. 

Thus, molecules to participate in the two 

catalytic cycles involving chlorine and bromine 

atoms are much more abundant in the Antarctic. 

Second, the Arctic vortex breaks up and warms 

The dramatic seasonal 

ozone losses occur 

over Antarctica 

during August to 

October and to 

a smaller extent 

over the Arctic in 

February and March.
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our current understanding and recent findings 

related to the chemical and meteorological or 

dynamical processes that affect ozone (Section 

3.2.3).

3.2.1 Total Ozone Observations

3.2.1.1 Global ozone (excluding polar

regions)

After nearly two decades of decrease, the 

column amount of ozone at midlatitudes of the 

Northern and Southern Hemispheres has been 

relatively stable over the last decade. Polar 

ozone is considered in more detail in Section 

3.2.1.3 below. We can integrate over the globe 

to get a simple measure of the recent changes in 

the ozone layer (Figure 3.2). The global mean 

total column ozone values for 2002-2005 were 

approximately 3% (about 10 Dobson Units or 

DU) below 1964-1980 average values. The 

2002-2005 values are similar to the 1998-2001 

values and this indicates that, overall, ozone is 

no longer decreasing. Several global datasets 

confirm this conclusion, although differences 

of up to 1% between annual averages exist 

between some individual sets (WMO, 2007).

Total column ozone over the tropics (25°S to 

25°N) remains essentially unchanged. Total 

ozone trends in this region for the period 

1980-2004 are not statistically significant, 

consistent with earlier assessments (Figure 3.4, 

WMO, 2007). 

The behavior of ozone at midlatitudes in the 

Northern Hemisphere during the 1990s was 

different from that in the Southern Hemisphere 

during the same period. The Northern 

Hemisphere shows a minimum around 1993 

resulting from forcings from the Mt. Pinatubo 

eruption and the solar cycle minimum, followed 

by an increase. The Southern Hemisphere shows 

an ongoing decrease through the late 1990s, 

followed by relatively constant levels (Figure 

3.3). The average for the period 2002-2005 of 

total ozone at midlatitudes in each hemisphere 

is similar to the average for the previous 

four years, 1998-2001. Ozone in the southern 

midlatitudes remains about 5.5% below its 

1964-1980 average, while ozone in the northern 

midlatitudes remains about 3% below (Figure 

3.3).

3.2 OZONE

In this chapter we briefly review the most 

recent observed trends in observations of 

total ozone (Section 3.2.1) and ozone vertical 

distributions (Section 3.2.2). We then discuss 

Figure 3.2 Top panel: Ozone observations for 60°S to 60°N 

estimated from ground-based data and individual components 

that comprise ozone variations (Dobson Units or DU). Bottom 

panel: Ozone deviations after removing annual cycle (blue line), 

solar cycle (red line), quasi-biennial oscillation or QBO (magenta 

line), and volcanic effects (green line) from original time series. 

Seasonal variations in the Equivalent Effective Stratospheric 

Chlorine (EESC) related component (up and down variations 

in orange line) are also removed. The thick orange line in the 

bottom panel represents the annual average EESC component 

derived from the regression model. See Box 3.1 for additional 

details.
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

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BOX 3.1: Estimating Ozone Trends

Isolating the ozone response to manmade ozone-depleting substances from natural variations in the 

ozone, such as seasonal changes or volcanic perturbations, is accomplished using a statistical time series 

regression analysis. The top panel of Figure 3.2 (black line) shows total ozone time series from ground-based 

measurements taken over the period 1964 to 2006 and averaged seasonally and over the 60 ̊N-60 ̊S area 

(87% of the Earth’s area). The observations are statistically modeled as a linear combination of the known 

individual processes that cause ozone to vary. In this analysis (following Fioletov et al., 2002), the regression 

model used is:

O3(t) = μ + seasonal cycle + α•EESC + β•QBO + γ•Solar + δ•Volcano + noise

Here, μ, α, β, γ, and δ are constants estimated such that the model (terms on the right hand side) best 

matches the observed ozone time series. The mean (μ) and seasonal cycle are calculated directly from the 

ozone data from 1979 to 1987 (blue line in Figure 3.2). Equivalent Effective Stratospheric Chlorine (EESC, 

see Chapter 5) is used to represent anthropogenic trace gases that react with ozone (orange line in Figure 

3.2). The magenta line shows the quasi-biennial oscillation (QBO). The QBO is a variation in stratospheric 

winds with a period of about 26 months that is represented using equatorial radiosonde wind observations 

(Reed et al., 1961). The solar term is represented using the 10.7 cm radio flux measured at Ottawa, Canada 

(red line). The volcanic term is derived from stratospheric aerosol observations (dark green). The noise term 

includes all variations required to make the model exactly equal the observed ozone (grey). The coefficients 

(μ, α, β, γ, and δ) are estimated by a mathematical regression that minimizes the noise term.

The bottom panel highlights the ozone changes due to chlorine and bromine (i.e., EESC) with the natural 

forcings (seasonal cycle, QBO, solar, and volcano) removed. This line is the original observations with only 

the annually-averaged EESC-related time series (smoothed orange line) and the residual noise term remaining 

(grey line).

Figure 3.3 Top: deseasonalized, annually averaged, area-weighted total ozone deviations from satellite (red) and ground stations 

(black) for the latitude bands 35°N to 60°N (left) and 35°S to 60°S (right). Anomalies were calculated with respect to the time 

average for the period 1964-1980. Updated from Fioletov et al. (2002) and WMO (2003). Bottom: Average total ozone over the 

United States from the TOMS/SBUV series of satellite instruments (red), and seven ground stations in the United States. Both 

time series are plotted relative to the 1964-1980 mean of the ground-station data. Updated from Stolarski and Frith (2006). 
The U.S. Climate Change Science Program Chapter 3

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3.2.1.2 Polar ozone

Significant ozone depletion has occurred in 

the polar regions over the last few decades as 

a result of anthropogenic halogen-containing 

compounds. The ozone loss chemistry, as 

described in Chapter 1 (also WMO, 2007 

and references therein), begins with very 

cold temperatures that lead to the formation 

of PSCs. Chlorine is rapidly converted from 

inactive to reactive forms on the cold aerosol 

surfaces. The Antarctic ozone hole is the most 

extreme manifestation of this phenomenon. 

Reactive chlorine is released within the 

stratospheric polar vortex beginning in the 

winter darkness. In August through September, 

when sunlight has returned to the Antarctic, 

halogen photochemistry rapidly destroys ozone. 

Some ozone loss is also observed in the JuneAugust period at the edge of the polar vortex 

(Roscoe et al., 1997). Ozone loss maximizes by 

the late September to early October period, after 

which temperatures warm, ozone loss ceases, 

the polar vortex breaks up, and high ozone air 

from midlatitudes mixes in, rapidly filling in 

the ozone hole (typically in the NovemberDecember period).

Total ozone over the United States tends 

to parallel the entire Northern Hemisphere 

because these levels are driven by the response 

to the worldwide chlorine and bromine releases 

and by hemispheric-scale transport processes 

(Figure 3.3, bottom). Releases of ozonedepleting substances in the United States 

affect global ozone levels, and releases across 

the globe affect the United States, because of 

the long lifetimes of CFCs and their mixing, or 

spread, around the world. Total ozone over the 

United States is shown in the bottom panel of 

Figure 3.3. The total ozone changes are similar 

to ozone over the entire northern midlatitudes 

(compare to top right panel). The minimum 

value was reached shortly after the eruption of 

Mount Pinatubo and was about 5 to 8 percent 

below the amounts present prior to 1980 (as 

a result of the volcanic aerosol effect). The 

ozone increase since 1993 has diminished the 

ozone deficit to about 2 to 5 percent below the 

pre-1980 amounts. The average for the last four 

years (2002-2005) is essentially the same as the 

previous four years.

Releases of ozonedepleting substances 

in the United States 

affect global ozone 

levels, and releases 

across the globe 

affect the United 

States, because of 

the long lifetimes 

of CFCs and their 

mixing, or spread, 

around the world.

Figure 3.4 March monthly averaged total ozone. The 1971 and 1972 images are from the Nimbus-4 BUV instrument, 

the 1979 is from the Nimbus-7 TOMS satellite instrument, the 1997 and 2004 images are from the Earth Probe 

TOMS, and the 2005, 2006, and 2007 images are from the Aura Ozone Monitoring Instrument (OMI). This figure 

is updated from Figure 4-6 of WMO (2007).
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

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showed a less pronounced impact on the March 

polar average total ozone. Although Northern 

Hemisphere polar column ozone averages are 

a general indicator of Arctic ozone depletion 

and trends (WMO, 2007), the chemical loss 

can oftentimes be masked by the 63-90°N 

polar averaging. For example, the 2005 March 

average had a strong influence of dynamics. 

Vortex fragments moved outside the 63°-90°N 

and the total ozone showed a distinct minimum 

near 60°N (Figure 3.4). This created a higher 

value relative to other recent cold winters 

even though chemical ozone loss in the lower 

stratospheric vortex in mid winter of 2005 was 

as high as or higher than ozone loss in other 

recent cold winters. 

3.2.1.2.2 Antarctic total ozone

In the Southern Hemispheric (SH) polar 

region, very large ozone depletions in the 

Austral spring have led to extremely low 

ozone values over Antarctica during October, 

the “ozone hole” (Figure 3.5, bottom line). 

Figure 3.6 displays a series of Antarctic total 

ozone images (values shown in Figure 3.5 are 

averaged from these images). A comparison 

of the moderate values of total ozone over 

In this section, we illustrate trends in total 

ozone for both the Arctic and Antarctic. The 

ozone content in the polar lower stratosphere is 

dependent on background chemical conditions, 

temperatures, transport, and dynamics. The 

Arctic polar stratosphere shows large interannual 

variability, while the Antarctic is more stable 

because the Antarctic polar vortex is more 

stable. This section discusses the behavior of 

polar ozone over the last few decades. Section 

3.2.1.2.1 focuses on the Arctic, while 3.2.1.2.2 

shows the Antarctic.

3.2.1.2.1 Arctic total ozone

Arctic total ozone has had a substantial 

downward trend since the 1970s with slightly 

higher values over the last ten years than in the 

previous six years. Figure 3.4 displays a series 

of March polar averages for selected years 

from 1971 to 2007 (updated from Figure 4-6 in 

WMO, 2007). The 60°N latitude circle generally 

encloses the region of ozone depletion, but in 

some years (e.g., 2005) the vortex and low ozone 

region are displaced from the pole, extending 

somewhat southward of 60°N. Nevertheless, 

Arctic ozone for recent March averages is low 

compared to the observations prior to 1980 

(shown in the upper row of Figure 

3.4).

The springtime average total ozone 

values in the Arctic poleward of 

63°N latitude (upper line) are shown 

in Figure 3.5, in comparison with 

the average total ozone for the years 

1970-1982 (gray horizontal line). 

The difference between the observed 

values and the 1970-1982 average 

indicates the combined changes in 

ozone due to chemistry and dynamics. 

In the last ten years Arctic column 

ozone is higher than the low values of 

the mid-1990s, except in the cold and 

chemically active winter of 1999/2000, 

when a large decrease of 63°-90° 

Northern Hemispheric total ozone 

was observed (Rex et al., 2002).

The record-cold winter of 2004/2005 

led to very large ozone losses (Manney 

et al., 2006; Rex et al., 2006; Singleton 

et al., 2007; Goutail et al., 2005; Feng 

et al., 2007a). However, this large loss 

Arctic total ozone 

has had a substantial 

downward trend since 

the 1970s, with slightly 

higher values over the 

last ten years than in 

the previous six years.

Figure 3.5 Total ozone average of 63 ̊-90 ̊ latitude in March (Northern Hemisphere) 

and October (Southern Hemisphere). 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 for the Northern Hemisphere and Southern Hemisphere. The 

grey shading shows the contribution of chemical ozone destruction and natural variations. 

Updated from Figure 4-7, WMO (2007).
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Antarctica in the early years (1970s, top row) 

to the reduced values over Antarctica in the 

last two decades (bottom row) illustrates the 

Antarctic ozone hole. In Figure 3.5, the years 

from 2000 to 2005 showed an increase in polar 

column ozone averages compared to 1998 

and 1999. The interannual variations in ozone 

depletion observed from 2001 to 2005 primarily 

result from variations in the dynamics (i.e., 

stratospheric weather variations), and have not 

been caused by changes in Equivalent Effective 

Stratospheric Chlorine (EESC). See Box 2.7 

of Chapter 2 for a definition of EESC and see 

Chapter 5 for more discussion on its usage. 

Since the early 1990s, total loss of ozone occurs 

in the lowermost stratosphere inside the polar 

vortex in September and October (Solomon et 

al., 2005). Estimates of EESC inside the vortex 

reached a value of about 3.2 parts per billion by 

mole (ppb) in 1990 and peaked in early 2001 at 

about 4.0 ppb (Newman et al., 2007). Hence, the 

EESC concentrations since the early 1990s have 

exceeded those necessary to cause total loss. 

The Antarctic ozone hole, therefore, has had 

low sensitivity to moderate decreases in EESC 

and the unusually small ozone holes in some 

recent years (e.g., 2002 and 2004) are strongly 

Figure 3.6 Satellite observations of October monthly averaged total ozone. The 1970, 1971, and 1972 images 

is from the Nimbus-4 BUV instrument, the 1979 image is from the Nimbus-7 TOMS instrument, and the 2004, 

2005, 2006 and 2007 images are from the Aura OMI instrument.

Figure 3.7 The minimum ozone values over Antarctica are averaged for the 

period from 21 September to 16 October (black dots). The vertical grey bars 

indicate the range of ozone values used in the average. The blue line shows 

the fit to these ozone values as was shown in Newman et al. (2004), and now 

using EESC, as derived in Newman et al. (2006) (also Box 2.7 in Chapter 2). The 

EESC has a mean age of 5.5 years, an age spectrum width of 2.75 years, and 

a bromine-scaling factor of 60. The fit is quadratic in EESC. The background 

lighter grey shading shows the expected variation of minimum ozone values 

between warm (upper side = +10°C) and cold years (lower side = −10°C). 

This figure was generated using TOMS and OMI total ozone. Updated from 

Figure 4-8 of WMO (2007).
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

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remote sensing, Light Detection and Ranging 

(lidar) remote sensing, and microwave ozone 

measurements confirm these findings.

The bulk of column ozone is found in the 

lower part of the stratosphere (Figure 3.1). 

The evidence shows that lower stratospheric 

ozone declined over the period 1979-1995, but 

has been relatively constant with significant 

variability over the last decade. Figure 3.9 

shows the vertical profile of ozone trends in 

midlatitudes of the Northern (left panel) and 

Southern (right panel) Hemispheres. The trends 

are actually fits to EESC (∆O3 = α• ∆EESC, see 

Box 3.1) that is converted to a percentage per 

attributable to a dynamically driven warmer 

Antarctic stratosphere.

Various metrics that capture different aspects of 

the Antarctic ozone hole are used to describe the 

severity of ozone depletion, such as Antarctic 

ozone hole area, ozone minimum, ozone mass 

deficit, and profile shape (Section 3.2.2.2). The 

polar average from 63-90°S tends to exaggerate 

dynamical fluctuations (Figure 3.5). Figure 3.7 

displays the Antarctic ozone hole minimum 

values averaged for the period 21 September 

to 16 October. Because the Antarctic ozone 

hole chemical losses peak in late September, 

the average minimum ozone columns in this 

period provide a very useful metric for the 

depletion severity. Again, this figure shows a 

clear decrease from 1979 to the mid-1990s, with 

particularly low values in the mid- to late 1990s. 

Following Newman et al. (2006), we have added 

a statistical fit of these metrics (blue line) to 

a quadratic function of Antarctic EESC. The 

fit shows how ozone levels have responded to 

chlorine. In addition to the fit to chlorine, the 

figure also includes a background grey shading 

that shows the expected natural variation of the 

ozone minimum values for warmer than average 

years (+2σ = 10 K, upper part) and colder years 

(−2σ = −10 K, lower part). The 2002 minimum 

value stands out because it was the warmest 

year on record. The minimum ozone values in 

2002 and 2004 were higher than the expected 

values (the blue line) because of the warmer 

temperatures.

3.2.2 Vertical Distribution of Ozone

3.2.2.1 Global

In addition to the polar regions, the upper 

stratosphere also shows clear evidence for 

ozone destruction due to increasing chlorine 

compounds. Measurements from both the 

Stratospheric Aerosol and Gas Experiment 

(SAGE I+II) and Solar Backscatter Ultraviolet 

(SBUV[/2]) satellite instr uments show 

significant declines in upper stratospheric 

ozone from 1979 through 2004 (Figure 3.8). The 

net ozone decrease over the 1979 to 1995 period 

was approximately 10-15% over midlatitudes, 

with smaller but significant changes over the 

tropics (Figure 3-7 in WMO, 2007). During 

the last decade, upper stratospheric ozone 

has remained relatively constant. Available 

independent Umkehr ground-based optical 

Figure 3.8 Time series of upper stratospheric ozone anomalies measured 

by ground-based lidar and microwave radiometers at five stations and 

corresponding zonal means from satellite (SAGE, HALOE, and SBUV) 

measurements (updated from Steinbrecht et al., 2006 and WMO, 2007).
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decade by scaling the α coefficient with the 

linear 1 ppb change of EESC observed during 

the 1980s. Measurements by SAGE I+II and 

SBUV(/2) showed declines of 7 to 9% (or 10 to 

15% cumulative by 1995) between 40 and 45 

km altitude (Figure 3.9).

These midlatitude ozone decreases are 

not linear, and did not continue in the 

last decade. This non-linear trend has 

been accounted for by using the ozone 

regression against the EESC time series 

and then converting to % per decade 

using the variation of EESC with time in 

the 1980s. At lower altitudes, between 12 

and 15 km, in the Northern Hemisphere, 

a strong decrease in ozone was observed 

from ozonesonde data between 1979 and 

1995, followed by an overall increase 

from 1996 to 2004, leading to no net longterm decrease at this level. These changes 

in the lowermost stratosphere have a 

substantial influence on the column 

because most of the ozone resides in the 

lowermost stratosphere.

3.2.2.2 Polar

The Antarctic ozone hole first began to 

develop in the early 1980s and reached 

its current full extent by the mid-1990s 

(Hofmann et al., 1997; Solomon et al., 2005). 

The most complete record of the morphology 

of the Antarctic ozone hole vertical structure 

is found from the balloon-borne ozonesonde 

The Antarctic 

ozone hole first 

began to develop 

in the early 1980s 

and reached its 

current full extent 

by the mid-1990s.

Figure 3.9 Vertical profile of ozone trends over northern and southern midlatitudes estimated from 

ozonesondes, Umkehr, SAGE I+II, and SBUV(/2) for the period 1979-2004. The trends were estimated 

using regression to an EESC curve and converted to % per decade using the variation of EESC with time 

in the 1980s. The trends were calculated in geometric altitude coordinates for SAGE and in pressure 

coordinates for SBUV(/2), sondes, and Umkehr data, and then converted to altitude coordinates using 

the standard atmosphere. The two standard deviation error bars are shown.

Figure 3.10 Observations of the October average ozone 

profiles measured at the South Pole in different time 

periods; prior to the Antarctic ozone hole (1966-1971), 

after the Mt. Pinatubo eruption when aerosol abundances 

were enhanced in (1992-1995), and current conditions 

(1996-2004). Adapted from Solomon et al. (2005).
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

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also tend to persist for shorter lengths of time in 

the Arctic and these cold regions are generally 

not concentric with the Arctic polar vortex, but 

are frequently centered roughly in the region 

between Greenland and Norway. Thus, ozone 

levels in the Arctic lower stratosphere exhibit 

a large amount of variability, which is well 

correlated with temperature. This is primarily 

the result of the fact that in the Arctic lower 

stratosphere the average temperature is very 

near the PSC formation threshold temperature. 

Therefore, in cold winters, PSCs tend to be very 

abundant and large halogen-catalyzed ozone 

depletion occurs, whereas in warm winters 

PSCs are very infrequent and little chemical 

ozone depletion occurs (Rex et al., 2004). This 

is illustrated later in Section 3.2.3.1.1 in Figure 

3.12, which shows a very good correlation 

between the volume of air with temperatures 

cold enough to be capable of forming PSCs 

and the chemical loss of ozone in the lower 

stratosphere. 

A particular problem with regard to assessing 

trends in polar ozone loss is that the distribution 

and variation of stratospheric ozone are 

controlled by both transport processes and 

photochemical processes. Ozone trends 

resulting from changes in atmospheric halogen 

loading must be separated from trends resulting 

measurements at the South Pole, which extend 

back to the mid-1960s. Figure 3.10, from 

Solomon et al. (2005), uses the South Pole 

ozonesonde data to delineate the Antarctic 

ozone hole region relative to the pre-ozone hole 

conditions of the 1970s. The altitude range of 

the Antarctic ozone hole has been very stable 

in the 1990s. In the vicinity of the lower edge of 

the Antarctic ozone hole (10 to 14 km), Figure 

3.10 shows that ozone abundances were lowest 

in the 1992-95 time period. This is presumably 

the result of increased ozone loss resulting 

from the enhanced aerosol loading after the 

Mt. Pinatubo eruption (Hofmann et al., 1997; 

Solomon et al., 2005). 

Also of interest is the ozone variability near 

the top edge of the Antarctic ozone hole. 

Ozone abundances in this layer between 18 

and 22 km may provide an early indication 

of Antarctic ozone hole recovery (Hofmann 

et al., 1997). However, as discussed further 

below, the higher abundances in the 2001-2004 

period have been attributed to meteorological 

variations rather than to ozone recovery (e.g., 

Hoppel et al., 2005). During 2002-2004, the 

temperature in the 20-22 km region tended 

to be warmer than average from mid-August 

through September, resulting in fewer PSCs 

which inhibited ozone loss (Hoppel et al., 

2005). The most extreme manifestation of 

this inhibited ozone loss occurred in 2002. As 

described in Section 3.2.3.1.1, in September of 

that year the first documented Antarctic major 

warming event took place (Roscoe et al., 2005). 

Major warmings are defined as reversals of both 

the vortex flow and the temperature gradient 

in the middle stratosphere; these events are 

relatively common in the Arctic, but had not 

been previously observed in the Antarctic. 

In 2002, anomalously high ozone levels and 

temperatures extended down to 15 km.

The Antarctic ozone hole generally behaves in 

a regular fashion, since the Antarctic winter 

stratosphere is consistently cold, with a stable, 

isolated vortex and an abundance of PSCs each 

winter. As discussed in Section 3.2.3.1.1, the 

Arctic winter stratosphere exhibits much more 

variability. Compared to the Antarctic, the 

Arctic is generally warmer with fewer PSCs 

(Fromm et al., 2003 Figure 3-13). Periods of cold 

temperatures with elevated reactive chlorine 

Compared to the 

Antarctic, the Arctic 

is generally warmer 

with fewer polar 

stratospheric clouds 

(PSCs), which play a 

role in the depletion 

of the ozone layer.

Figure 3.11 Ozone loss estimates from the Polar Ozone and Aerosol 

Measurement (POAM II & III). Southern Hemisphere October 5 estimates 

(solid lines) and Northern Hemisphere March 10 estimates (dashed lines) 

are based on the “vortex average technique” described in WMO (2007) 

and Hoppel et al. (2002, 2003). Estimates are shown only for Northern 

Hemisphere winters which had a relatively persistent, isolated vortex from 

January 1 - March 10.
The U.S. Climate Change Science Program Chapter 3

92

This section is divided into two subsections. 

The first subsection discusses the influence 

of dynamics on polar ozone, while the second 

subsection addresses the influence of dynamics 

on midlatitude ozone. 

3.2.3.1.1 Polar

Variability in the dynamical conditions in the 

troposphere/stratosphere system results in 

variability of ozone transport and temperatures 

in the polar stratosphere. Previous World 

Meteorological Organization (WMO)/

United Nations Environment Programme 

(UNEP) assessments have shown that, on 

short timescales, interannual variability in 

polar ozone chemistry is mainly driven by 

temperature variability, which in turn is the 

result of variable dynamical conditions. The 

combined effect of dynamically-induced 

variability in both chemistry and transport is 

the main driver of interannual variability of the 

abundance of ozone in the polar stratosphere. 

As described in the Introduction (Section 

3.1), the air in the polar lower stratosphere 

is transported downward from the upper 

stratosphere and mesosphere over the course 

of the winter period by the Brewer-Dobson 

circulation. The Brewer-Dobson circulation 

is driven by large-scale atmospheric waves 

that propagate upward from the troposphere. 

Figure 3.1 shows this poleward and downward 

circulation in the annual average. This upper 

stratospheric air has on average been in the 

stratosphere for five to six years since entering 

the stratosphere at the tropical tropopause. In 

the absence of polar ozone destruction, this 

air would be characterized by relatively high 

ozone concentrations. Furthermore, because 

the air has been in the upper stratosphere 

and exposed to intense solar UV, the organic 

chlorine and bromine compounds have been 

almost completely converted to inorganic forms 

that can participate in ozone loss processes.

The chemical ozone loss processes precipitated 

by the presence of halogens are initiated by the 

formation of PSCs in the extremely cold polar 

lower stratosphere (Crutzen and Arnold, 1986; 

Toon et al., 1986). PSCs provide a surface upon 

which heterogeneous (not gas-phase, but at 

the surface between a solid/liquid and a gas) 

reactions take place that convert comparatively 

from transport variations. Instruments measure 

ozone abundances and their variations, but do 

not directly measure ozone photochemical loss. 

Isolating the photochemical ozone change in the 

Arctic is more complicated than in the Antarctic 

because of the much larger degree of dynamical 

variability. Several different methods have 

been developed for isolating photochemically 

driven ozone change from transport-driven 

change. For cold Arctic winters (in which there 

is measurable loss), ozone loss derived from 

each of these methods now agree fairly well 

(e.g., WMO, 2007, Figure 4-11). Therefore, 

we now have a fairly reliable record of ozone 

chemical loss for all Antarctic winters, and 

for cold Arctic winters, dating back to the 

mid-1990s. As an example, Figure 3.11 shows 

vertical profiles of photochemical loss derived 

from Polar Ozone Aerosol and Measurement 

(POAM II & III) measurements, for both the 

Arctic and Antarctic, during the 1994-2005 

time period (Hoppel et al., 2003). Ozone loss 

in the Antarctic ozone hole was fairly stable 

in the 1990s, with nearly complete loss in the 

14 to 19 km altitude range. Figure 3.11 shows 

that the anomalously high ozone levels in the 

upper region of the Antarctic ozone hole in 2001 

through 2004 were the result of reduced ozone 

chemical loss. In contrast to the Antarctic, the 

ozone loss profiles for the Arctic are highly 

variable with peak losses of almost 50% (losses 

up to approximately 60% have been reported by 

other analyses [Rex et al., 2004; WMO, 2003; 

WMO, 2007]). 

3.2.3 Processes That Affect Ozone

3.2.3.1 Transport and dynamics

Stratospheric ozone levels are strongly influenced 

by both transport and the temperatures of the 

stratosphere. In this section, we will summarize 

the influence of dynamical processes on ozone 

levels. First, there is the direct influence of 

winds that carry ozone-enriched air from the 

photochemical production region into other 

regions, thereby increasing ozone. Second, the 

opposite process can occur where winds carry 

ozone-depleted air into other regions, thereby 

decreasing ozone (e.g., from the Antarctic ozone 

hole into the midlatitude stratosphere). Third, 

the radiatively and dynamically driven local 

temperature can influence ozone by affecting 

catalytic loss reaction rates. 

The Antarctic 

ozone hole is more 

severe in colder 

than average years, 

while less severe 

in warmer than 

average years.
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

correlation of the north-south wind and the 

temperature) at an altitude of 16 km in the 

45-75°S zone (see Andrews et al. (1987) for a 

more complete description of the wave driving 

of the stratosphere by the troposphere). In 

September 2002, a major warming had a 

dramatic impact on total ozone, splitting the 

Antarctic ozone hole into two pieces (Stolarski 

et al., 2005). Meteorological conditions in 

2002 showed that the early winter was already 

unusually disturbed (Hio and Yoden, 2005; 

Newman and Nash, 2005; Allen et al., 2003). 

There were several significant wave events 

from May to October that each warmed the 

stratosphere by a few degrees until the major 

warming in late September. Several models 

reproduced the chemistry and dynamics of this 

2002 warming, revealing the direct impact of 

tropospheric waves on Antarctic ozone levels 

(Manney et al., 2005; Ricaud et al., 2005; 

Konopka et al., 2005; Grooß et al., 2005; 

Sinnhuber et al., 2003; Feng et al., 2005).

3.2.3.1.2 Midlatitude dynamic and transport 

effects on ozone

The influence of transport and dynamics on 

the midlatitude lower stratosphere (16 to 30 

km) and lowestmost stratosphere (8 to 16 km) 

principally occurs through the Brewer-Dobson 

circulation and through mixing processes. 

While photochemistry plays an important role 

unreactive chlorine reservoirs into ones that 

are exceedingly reactive in sunlight. While the 

chlorine and bromine levels in the stratosphere 

directly cause ozone loss, year-to-year variation 

of the chemically driven polar ozone loss is 

directly tied to the temperature by a modulation 

of polar stratospheric clouds and transport. 

A number of studies have shown that the 

Antarctic ozone hole is more severe in colder 

than average years, while less severe in 

warmer than average years (Newman and 

Randel, 1988; WMO, 1989). In the Arctic, 

Rex et al. (2004) quantitatively related the 

volume of polar stratospheric clouds (VPSC) 

to the chemical ozone loss estimated from 

ozonesondes (extended to the Antarctic by 

Tilmes et al., 2006). Figure 3.12 shows ozone 

loss plotted against VPSC for the years from 

1992 to the present (the latest data are available 

through 2007). This 1992-2007 period has 

high chlorine and bromine levels (Chapter 2). 

For the coldest Arctic winters, the volume of 

air with temperatures low enough to support 

polar stratospheric clouds (VPSC) increased 

significantly since the late 1960s (Rex et al., 

2006). The cooling of the lower stratosphere 

is much larger than expected from the direct 

radiative effect of increasing greenhouse gas 

concentrations. The reason for the change 

is not clear and it could be due to long-term 

natural variability or an unknown dynamical 

mechanism.

The year-to-year variation of spring temperatures 

in the polar stratosphere is primarily driven by 

year-to-year variability of planetary waves 

that propagate upward from the troposphere to 

the stratosphere. The relationship of waves to 

stratospheric ozone was recognized by a number 

of early investigators who saw large increases 

of total ozone following major stratospheric 

warmings (London, 1963). 

The large variability of polar total ozone shown 

in Figures 3.5 and 3.7 is directly tied to the 

variations in the levels of the planetary waves 

(Randel et al., 2002). The Southern Hemisphere 

winter of 2002 provides an excellent example of 

a year with extremely high levels of planetary 

waves propagating into the stratosphere. The 

planetary wave forcing of the stratosphere is 

estimated from the eddy heat flux (a cross 

The lifetime of ozone 

in the midlatitude 

lower stratosphere 

(16 to 30 km) is 

more than 100 

days. Therefore, 

transport through 

the atmosphere plays 

a very important 

role in determining 

midlatitude 

ozone levels.

Figure 3.12 Scatter plot of vortex-average chemical loss of column ozone 

(∆O3, calculated over the range 380 to 550 K) versus VPSC inferred from 

ozonesonde observations for the 1991/1992 to 2006/2007 Arctic winters 

(update from Rex et al., 2004). Colored squares and the red fit line show results 

based on ozonesonde analyses; colored circles and the green fit line show results 

from tracer correlation studies based on HALOE data (update from Tilmes et 

al., 2006). Adapted from Rex et al. (2006, 2004), and Tilmes et al. (2006).
The U.S. Climate Change Science Program Chapter 3

94

Empirical studies using inferred circulation 

changes indicate that a substantial fraction 

of the observed Southern Hemisphere and 

Northern Hemisphere variability results from 

variations of the wave driving and by inference 

the Brewer-Dobson circulation (Salby and 

Callaghan, 2002; Salby and Callaghan, 2004a; 

Salby and Callaghan, 2004b; Randel et al., 

2002; Hood and Soukharev, 2005). Randel et 

al. (2002) and Hood and Soukharev (2005) 

estimated that 18-30% of the spring column 

ozone trends from 1979 to the mid-1990s might 

be attributable to long-term changes in lower 

stratospheric circulation. Such circulation 

changes may also have been responsible, at 

least in part, for the increase that has been 

observed at these latitudes since the mid1990s. It is important to recognize that these 

empirical studies are correlative and statistical 

in scope, and so are only proxies for actual 

ozone transport.

Estimates of the dynamically induced 

contributions to ozone interannual variability 

and trends can be derived by using chemical 

transport models (CTM) driven by observed 

temperature and wind fields (Hadjinicolaou 

et al., 1997; 2002; 2005). Using the SLIMCAT 

three-dimensional (3-D) CTM, Hadjinicolaou 

et al. (2005) found that about one-third of the 

observed ozone trend from 1979 to the mid1990s could be explained by transport-related 

changes. In addition, Hadjinicolaou et al.

(2005) also found that all of the midlatitude 

“increase” (see the period from the mid-1990s 

to 2004 in top right panel of Figure 3.3) could 

be explained by transport alone, and not by 

halogen decreases. However, the interannual 

variation discrepancies between CTMs and 

observations are large, making it difficult to 

place much weight on CTM results to attribute 

long-term transport changes.

The midlatitude ozone is influenced by polar 

loss via air mass mixing after the polar vortex 

breakup in early spring. Using regression 

analysis, Dhomse et al. (2006) concluded that 

this mechanism is one of the main factors 

responsible for the recent increase in Northern 

Hemisphere total ozone.

for ozone in the midlatitudes, the lifetime of 

ozone in the lower stratosphere is long (greater 

than 100 days), and hence, transport plays 

a very important role in determining ozone 

levels. In the upper stratosphere, dynamically or 

radiatively forced temperature changes can have 

a large effect on ozone loss rates by modifying 

the catalytic loss processes. Dynamically forced 

ozone changes in the lower stratosphere occur 

because of:

• interannual and long-term changes in 

the strength of the stratospheric BrewerDobson circulation (Figure 3.1), which is 

responsible for the winter-spring buildup of 

extratropical ozone (e.g., Fusco and Salby, 

1999; Randel et al., 2002; Weber et al., 

2003; Salby and Callaghan, 2004a; Hood 

and Soukharev, 2005) and; 

• changes in tropospheric circulation, 

particularly changes in the frequency of 

local nonlinear synoptic wave forcing 

events, which lead to the formation of 

extreme ozone minima (“mini-holes”) and 

associated large increases in tropopause 

height and horizontal mixing (Steinbrecht 

et al., 1998; Hood et al., 1997, 1999, 2001; 

Reid et al., 2000; Orsolini and Limpasuvan, 

2001; Brönnimann and Hood, 2003; Hood 

and Soukharev, 2005; Koch et al., 2005).

The effects of dynamics on ozone trends and 

variability are extremely difficult to quantify. 

This difficulty is caused by the relationship 

between the strength of the Brewer-Dobson 

circulation, the wave mixing processes, and 

the position and strength of the polar vortex. 

As is well recognized, the propagation of 

planetary scale waves from the troposphere 

into the stratosphere drives the BrewerDobson circulation, while at the same time the 

breaking of these waves irreversibly mixes air 

latitudinally. The estimation of ozone advection 

is further confused by the need to multiply the 

transport “variables” by the ozone horizontal 

and vertical gradients. This effect of the ozone 

gradient is mainly evident in two regions: the 

mixing of lower stratospheric ozone depleted 

air from the polar latitudes to the midlatitudes 

during the spring period, and the mixing of 

air from the tropical upper troposphere (with 

very low ozone amounts) into the midlatitude 

lowermost stratosphere. 

Heterogeneous 

reactions on polar 

stratospheric clouds 

(PSCs) convert 

the comparitively 

unreactive chlorine 

reservoirs into 

chlorine that is very 

reactive in sunlight.
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

photolyzes faster than assumed in WMO (2003). 

However, one recent laboratory study of the 

absorption cross-section of ClOOCl does not 

support this. It indicates that ClOOCl may 

photolyze (Equation 3b) slower than previously 

understood (Pope et al., 2007). However, 

this slower photolysis rate results in severe 

underestimates by photochemical models of 

observed O3 depletion rates and observed ClO 

levels, and hence poor representations of the 

severity of polar ozone losses (von Hobe et 

al., 2007). Current models (without Pope et 

al., 2007) reproduce the basic features of the 

Antarctic ozone hole and Arctic ozone losses 

using previous laboratory recommendations 

for photochemical parameters (e.g., WMO, 

2003; WMO, 2007). Clearly, more work will be 

required to understand this discrepancy.

Recent measurements show that bromine exists 

in the stratosphere at higher concentrations than 

is found in most 3-D models (WMO, 2007 and 

references therein). Profiles of BrO measured 

in the Arctic vortex suggest that inorganic 

bromine levels may be 3 to 8 parts per trillion 

(ppt) by volume larger than the amount of 

bromine carried to the stratosphere by methyl 

bromide (CH3Br) and halons alone (Canty et 

al., 2005; Frieler et al., 2006). Although still 

uncertain, the additional 3-8 ppt of bromine is 

probably derived from very short lived (VSL) 

species containing bromine that enter the 

stratosphere at the tropical tropopause (WMO, 

2007). Considering that the BrO + ClO cycle is 

now estimated to contribute up to half of total 

chemical loss of polar ozone, using the more 

efficient ozone loss by the ClO dimer cycle, this 

observation indicates the BrO + ClO catalytic 

cycle is likely to be a more efficient ozone 

loss process than considered in WMO (2003). 

Hence, bromine may play a more important 

role in polar ozone depletion than previously 

thought.

Polar stratospheric clouds are critically 

important in ozone photochemistry primarily 

through two processes: chlorine activation and 

denitrification. The chlorine heterogeneous 

reactions and sunlight lead to chlorine 

activation, while removal of nitric acid (HNO3) 

occurs as PSCs fall out of the lower stratosphere 

and remove nitrogen, or denitrify, that air. 

Satellite observations of aerosols and clouds 

3.2.3.2 Chemistry

3.2.3.2.1 Polar chemistry

Heterogeneous reactions on PSCs convert the 

comparatively unreactive chlorine reservoirs 

hydrochloric acid (HCl) and chlorine nitrate 

(ClONO2) first to chlorine gas (Cl2) in the 

long, dark polar night. As soon as the Sun first 

appears over the horizon in the Antarctic spring 

in August each year, the Cl2 photolyzes (breaks 

apart into chlorine atoms in the presence 

of sunlight, Cl2 + hν → 2 Cl) and Cl atoms 

react with ozone to make chlorine monoxide 

(ClO) (see reaction 3c in Section 3.1). These 

reactions are often called “chlorine activation,” 

since the chlorine compounds are converted 

from comparatively unreactive forms to much 

more photochemically reactive forms. At high 

concentrations, ClO reacts both with itself 

(reaction 3a forms the ClO dimer, dichlorine 

peroxide [ClOOCl], a reaction that actually 

proceeds faster at lower temperatures) and with 

the analogous bromine monoxide, BrO (see 

reaction 4a). Almost all of the rapid ozone loss 

in the Antarctic spring is attributed to catalytic 

cycles formed from the reaction of ClO with 

itself (reactions 3) and with BrO (reactions 4) 

(Frieler et al., 2006).

Thus, stratospheric chlorine levels provide 

the fundamental driver for polar ozone loss, 

since chlorine is involved in the principal 

catalytic cycles responsible for polar ozone loss. 

Beyond this basic understanding, however, the 

calculated chemical loss rates of polar ozone 

are still quantitatively uncertain. Questions 

remain to be resolved on the photolysis rate 

of the ClOOCl (Equation 3b) and the balance 

between ClO and ClOOCl in the Antarctic 

stratosphere and the atmospheric abundance 

of bromine. Higher levels of bromine would 

improve the comparison between theory and 

observation for Arctic and Antarctic loss rates, 

but the exact sources of the extra bromine are 

somewhat uncertain.

From in situ aircraft measurements, Stimpfle 

et al. (2004) suggested that the ClO dimer 

cycle (reactions 3) may be a more efficient 

process for polar ozone loss than previously 

thought (Frieler et al., 2006), and good overall 

consistency between in situ observations of 

ClO and the ClOOCl and model calculations 

can be achieved if it is assumed that ClOOCl 

Bromine may play 

a more important 

role in polar 

ozone depletion 

than previously 

thought.
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variability and the average temperature is close 

to the PSC formation threshold temperature. 

In warm Arctic winters, little or no PSC 

activity is evident (for example, in the winter of 

1984/1985). However, even in the coldest Arctic 

winters, PSCs only reach a 25% frequency.

3.2.3.2.2 Global and midlatitude chemical 

processes

As in the polar regions, halogen increases (chlorine and bromine) have been the principal driver 

of ozone depletion over the past few decades in 

the midlatitudes. There is good overall agreement between observed long-term changes in 

ozone outside of the polar regions and model 

simulations that include the effects of increasing halogens. The models generally reproduce 

the observed ozone changes as a function of 

altitude, latitude, and season, confirming our 

understanding that halogen changes are the 

main drivers of global ozone changes (WMO, 

2007). These models predict that the decline in 

ozone should have ceased and that the next few 

decades should show the beginning of a recovery from the maximum depletion. This is supported by the statistical fit of globally averaged 

ozone observations with Equivalent Effective 

Stratospheric Chlorine (EESC), a quantity that 

peaked in the late 1990s (Figure 3.2).

The explosive eruption of Mt. Pinatubo in 

1991 injected large quantities of sulfur into 

the stratosphere (Trepte et al., 1993). The 

sulfur-enhanced stratospheric sulfate aerosols 

provided significantly more surfaces that could 

support heterogeneous chemical reactions, thus 

converting a higher fraction of stratospheric 

chlorine to catalytically active forms. The 

impact of aerosols on midlatitude ozone was 

greatest in the early 1990s after the eruption 

of Mt. Pinatubo in 1991 (Figure 3.3). The 

observed decrease in Northern Hemisphere 

column ozone in 1993 agrees with chemical 

dynamical models that include these effects 

(WMO, 2003, 2007). The same models predict 

that the aerosols from Mt. Pinatubo should 

have produced a significant decrease in ozone 

over midlatitudes of the Southern Hemisphere, 

but no effect has been seen in either satellite 

measurements or ground measurements at 

stations such as Lauder, New Zealand.

in the polar stratosphere began with NASA’s 

Stratospheric Aerosol Monitor (SAM) II in 

1978, and continued nearly uninterrupted 

to 2005. These measurements used solar 

occultation in the visible and shortwave infrared 

portion of the electromagnetic spectrum. In 

addition to SAM II, other instruments included 

the NASA series of Stratospheric Aerosol and 

Gas Experiment (SAGE) I-III and the Naval 

Research Laboratory’s Polar Ozone and Aerosol 

Measurement (POAM) II-III. 

As noted in Section 3.2.3.1.1, VPSC is a key 

parameter for estimating ozone loss. It is 

important to recognize that VPSC actually 

represents the volume of temperatures cold 

enough to form PSCs, not the actual PSC 

volume over the polar region. Nevertheless, 

temperatures are directly related to PSC 

occurrence frequency (Steele et al., 1983). 

The long-term PSC statistics are presented 

in Figure 3.13 (Fromm et al., 2003). Here the 

PSC frequency (the number of profiles with a 

PSC divided by the number of profiles inside 

the polar vortex) for entire winter seasons is 

shown. In the Antarctic, PSCs are more frequent 

than the Arctic (Fromm et al., 2003). There 

are large interannual variations in Antarctic 

PSC frequency but no obvious long-term 

trend. In the Arctic, as described in Section 

3.2.2.2, stratospheric temperatures exhibit large 

Models predict 

that the decline 

in ozone outside 

the polar regions 

should have ceased 

and that the next 

few decades should 

show the beginning 

of a recovery from 

the maximum 

depletion.

Figure 3.13 PSC frequency for the entire winter season. The frequency is 

calculated as the number of profiles with a PSC divided by the number of profiles 

inside the polar vortex. See Fromm et al. (2003) for details.
97

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

3.3 ULTRAVIOLET RADIATION 

AT THE EARTH’S SURFACE

3.3.1 Background (Factors 

Controlling UV Surface Irradiance)

The amount of UV radiation reaching the 

Earth’s surface is controlled by several key 

factors including cloud cover, aerosols, and 

amount of atmospheric ozone (with most of 

the ozone being in the stratosphere). Ozone and 

cloud cover are the most important atmospheric 

components limiting the amount of UVB (280 

to 315 nm) radiation able to reach the ground. 

Clouds and scattering aerosols reduce UV 

radiation at all wavelengths by reflecting a 

fraction of UV energy back to space, whereas 

ozone absorbs a fraction of the UV radiation 

only in the 280 to 340 nm range, with more 

absorption at shorter wavelengths than at longer 

wavelengths. Under special conditions, clouds 

can locally increase UV from 1% to 10% by 

cloud edge reflections. Extremely heavy cloud 

cover (black thunderstorm) can decrease UV 

almost 100%. Radiation with wavelengths 

shorter than 280 nm does not reach the surface 

in significant amounts because of absorption 

by the atmosphere (O3 and O2). Air pollution 

is an additional factor that can affect UV 

reaching the surface through the absorption 

and scattering by aerosols and absorbing trace 

gases such as tropospheric O3 and nitrogen 

dioxide (NO2). UV radiation at the surface is 

generally highest near the equator following the 

seasonally changing sub-solar point (latitude 

between ±23°), where stratospheric ozone is a 

minimum and the solar zenith angle (SZA) is 

the smallest. Larger amounts of UV radiation 

are seen at high altitude sites, especially those 

with predominantly dry and clear weather 

and large surface reflectivity (e.g., from snow 

or ice cover). Understanding, modeling, and 

measuring the factors affecting the amount 

of UV radiation reaching the Earth’s surface 

is important, since increases in UV radiation 

affect human health adversely through skin 

cancer (Diffey, 1991), eye cataracts (Taylor, 

1990), and suppression of the immune system 

(Vermeer et al., 1991), and positively through 

increased Vitamin D production (Grant, 2002; 

Holick, 2004). Changes in UV radiation also 

have important effects on ecosystem biology 

(Smith et al., 1992; Ghetti et al., 2006). 

The inclusion of additional inorganic bromine 

(Bry) from very short-lived substances (VSLS) 

in models leads to larger ozone destruction at 

midlatitudes, compared with studies including 

only long-lived bromine source gases (e.g., 

Salawitch et al., 2005; Feng et al., 2007b). 

The enhanced ozone loss occurs in the lower 

stratosphere via interactions of this bromine 

with anthropogenic chlorine. Midlatitude ozone 

loss is primarily enhanced during periods of 

high aerosol loading. The impact on longterm midlatitude ozone trends (1980-2004), 

assuming constant VSLS Bry, is calculated to 

be small because aerosol loading was low at the 

start and end of this time period.

The profile of upper stratospheric ozone trends 

from 1980-2004 is generally consistent with our 

understanding of gas-phase chlorine chemistry 

as the cause of declining ozone, modulated 

by changes in temperature and other gases 

such as methane (WMO, 2007). However, 

global dynamical-chemical models have not 

demonstrated that they can simultaneously 

reproduce realistic trends in all relevant 

parameters, although observations over the full 

time period are limited (Eyring et al., 2006). 

Chemical models without interactive radiation 

obtain ozone changes that peak at about 14% for 

1980-2004 (in altitude coordinates), consistent 

with SAGE observations. 

Our ability to reproduce observed past changes 

in the Northern Hemisphere is better than that 

for the Southern Hemisphere. Two-dimensional 

(2-D) models show large model-model 

differences in the Southern Hemisphere due 

to different treatments of the Antarctic ozone 

loss and how it is spread to the midlatitudes. 

Three-dimensional CTMs are inherently better 

at simulating the polar regions and this leads 

to smaller model-to-model differences. These 

CTMs, however, still do better at reproducing 

long-term changes in the Northern Hemisphere 

than in the Southern Hemisphere (WMO, 

2007). This ongoing disagreement between 

model-observation comparisons in the Northern 

versus the Southern Hemisphere indicates that 

we do not yet have a full understanding of the 

combined chemical and transport processes 

controlling ozone changes at midlatitudes.

Reductions in ozone 

lead to increases 

in ultraviolet 

(UVB) radiation 

at the Earth’s 

surface. Higher 

UVB radiation 

can increase 

skin irritation.
The U.S. Climate Change Science Program Chapter 3

98

Diffey, 1987). The UV erythemal irradiance 

data shown in Figure 3.14 was obtained under 

clear-sky conditions at Mauna Loa, Hawaii, and 

shows the measured inverse relationship between 

ozone change and UVB radiation, which is 

the dominant portion of erythemal radiation. 

The relation to UV index and the units for 

Both theory and observations (Figures 3.14 

and 3.15) show that reductions in ozone lead to 

increases in UV erythemal radiation and UVB 

at the Earth’s surface. Erythemal radiation is 

a weighted average of UVA (315 to 400 nm) 

and UVB used as a measure of skin irritation 

caused by exposure to sunlight (McKinlay and 

Figure 3.14 Measured erythemal irradiances (lower curve) from an ultraviolet 

spectroradiometer at SZA 45° compared with total ozone (upper curve) for 132 clear mornings 

during July 1995 to July 1996 at Mauna Loa Observatory (19.5°N, 155.6°W, 3.4 km), showing the 

inverse relationship between erythemal UV and ozone amount (adapted from WMO, 1999).

(UV index 10 = 25 μW per cm2

 = 250 mW per m2)

Figure 3.15 Validation of Equation 1 using the measured (dark circles) changes in 

ozone and UV irradiance from Mauna Loa, Hawaii, shown in Figure 3.14 (adapted from 

WMO, 1999).
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

approximately gives the fractional change in 

irradiance for a change in ozone amount.

Fioletov et al. (1997) reported an extensive 

analysis of UVB irradiance and its dependence 

on total ozone. The analysis provides an 

empirical wavelength-by-wavelength measure 

of the increase of UVB irradiance for a 1% 

decrease of total ozone. These values were 

found to be essentially the same for clear and 

cloudy conditions (except for very heavy clouds) 

and are in good agreement with model results 

for longer wavelengths and moderate SZA.

UV radiation reaching the Earth’s surface varies 

on all time scales, from seconds to seasons. 

Hourly to daily changes, i.e., the short-term 

variations, are mostly due to cloud cover 

changes and aerosols, and to ozone in the UVB 

range. The extent of cloud cover also causes 

changes on daily and monthly time scales as 

the weather changes. In today’s atmosphere, the 

multi-year variations are controlled principally 

by changes in stratospheric ozone, changes in 

the extent of cloud cover, and other longer-term 

changes such as in the amount of aerosol and 

irradiance and dose are discussed in Appendix 

3B.

Increases in 280 nm to 340 nm UV radiation 

caused by decreases in ozone are easily 

estimated using radiative transfer calculations. 

For clear-sky conditions, the changes can 

also be accurately estimated using a simple 

relation between ozone and irradiance given 

in Equation 1:

dF/F= −dΩ/Ω αΩ sec(θ) = −dΩ/Ω (RAF) (1)

where the quantity αΩ sec(θ) is known as the 

Radiation Amplification Factor (RAF). 

The relationship is derived from the standard 

Beer’s Law of irradiance F attenuation in an 

absorbing atmosphere, F = Fo exp(−αΩ sec(θ)), 

where Ω isthe ozone column amount in Dobson 

Units (DU, equal to milli-atm-cm), α is the 

ozone absorption coefficient (in cm−1), θ is the 

solar zenith angle, and Fo is the irradiance at 

the top of the atmosphere (Madronich, 1993). 

An example to show the magnitude of the RAF 

as a function of wavelength is shown in Figure 

3.16 for θ = 45° and Ω = 330 DU = 0.33 

atm cm. The RAF method accurately 

estimates UV irradiance change 

compared to clear-sky radiative 

transfer (Herman et al., 1999b). For 

example, radiative transfer shows 

that a 1% decrease in O3 produces a 

2.115% increase in 305 nm irradiance, 

while the RAF method estimates a 

2.064% increase (Ω = 375 DU, θ = 

30°). Changes in measured erythemal 

irradiance are approximated very 

accurately using Equation 1 with an 

RAF = 1.38 when the ozone amount 

changed by 1% (Figure 3.14, Ω = 275 

DU, θ = 45°). For most conditions, 

erythemal irradiance change with 

ozone change behaves roughly the 

same as 308 nm irradiance.

The RAF approximation is useful for 

mid-day during the spring, summer, 

and autumn at most latitudes. During 

summer solstice, Equation 1 applies 

up to 83° latitude. In the presence of 

constant attenuation by cloud cover 

or scattering aerosols, Equation 1 still 

Figure 3.16 Ozone absorption coefficient α (cm-1) and the Radiation Amplification Factor 

RAF45 for a solar zenith angle SZA = 45° and ozone amount of 330 DU (Ω = 0.330 atm 

cm). Note that at 310 nm the RAF45 is approximately 1, so that a 1% increase in O3 would 

produce a 1% decrease in 310 nm irradiance.
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100

clear-sky conditions, the 305 nm irradiance 

typically changes by ±25% during June just 

from to day-to-day ozone changes. As will be 

discussed later, the day-to-day variability of 

clear-sky 40°N UV June irradiance is about 

three times larger than the change caused by the 

long-term June decrease in ozone from 1980 to 

2007 (dΩ/Ω ∼ −0.04).

Identification of long-term (decadal) changes 

from ground-based measured surface UV 

radiation due to stratospheric ozone depletion 

can be accomplished if the data are filtered to 

remove the effects of clouds. Trend detection 

from ground-based measurements under all 

sky conditions, though appealing and relevant, 

has many difficulties. This is primarily because 

the surface UV is highly variable, as noted 

above, due to factors such as cloud cover and 

aerosols, and because the stratospheric ozone 

depletion has been rather small (less than 10%) 

over the past decades, with the exception of high 

latitudes (greater than 60°). 

pollution. Day-to-day ozone-caused changes 

can be significant, but are usually smaller 

than changes due to cloud cover, because the 

stratospheric abundance of ozone only changes 

as a small percentage of its seasonally changing 

mean value. On longer time scales (decadal) 

most regional changes in cloud cover have been 

small (Herman et al., 2008), so that global and 

zonal average changes in UVB due to long-term 

ozone depletion are dominant. In some regions 

(e.g., northern Europe), decadal-term cloud 

changes are also important.

Ozone data from Nimbus-7/TOMS, obtained 

during June for the entire 5° longitudinal zone 

centered at 40°N, shows that the ozone amount 

can vary by 50 DU about the mean value of 

350 DU, or dΩ/Ω = ±0.14. The day-to-day June 

ozone variation is obtained from figures similar 

to those shown in Herman et al. (1995). Using 

an average noon SZA for June of about 23° and 

an ozone absorption coefficient for 305 nm α 

= 4.75 cm−1 yields a typical 305 nm irradiance 

change dF/F = −dΩ/Ω αΩ sec(θ) = ±0.14 * 

4.75 * 0.35 * 1.09 = ±0.25. In other words, for 

Figure 3.17 Frequency of occurrence of reflectivity values from 1979 to 1987 (3285 days) for six different 

locations. The mean and standard deviation (SD) are in RU (1 RU = 1%). Based on Herman et al. (2001a).
101

Trends in Emissions of Ozone-Depleting Substances, Ozone 

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Australia, where the average reflectivity (due 

to cloud cover) is very low and cumulative UV 

exposure is high compared to the same latitude 

in the United States.

Satellite observations of reflected UV indicate 

that ref lectivities for typical midlatitude 

cloud-covered scenes have a wide range of 

values, which can reach 90 RU over high 

altitude cloud tops that occur most frequently 

in the tropics. Under snow-free conditions, 

the surface reflectivity RG is usually between 

2 RU and 4 RU, reaching about 10 RU in 

the Libyan Desert and similar small areas 

(e.g., Andes Mountain high deserts). Areaaveraged clear-sky UV surface irradiance is 

then approximately reduced as a linear function 

of the cloud plus aerosol reflectivity, which can 

be written in terms of effective transmission 

(Krotkov et al., 2001), T ≈ (1 – R)/(1 – RG), with 

local values occasionally exceeding clear-sky 

irradiances by about 10% because of reflections 

from the sides of clouds. Midlatitude UV 

irradiance reductions caused by clouds range 

up to 50%, which is larger than the day-to-day 

305 nm UV variability caused by ozone (25%), 

and comparable to the change at 300 nm. 

Snow-covered scenes cannot be distinguished 

from cloud cover by observations in the UV 

wavelengths. Because of this, the use of 

reflectivity to estimate the amount of UV 

radiation at the surface in the presence of snow 

is likely to be in error. For example, the very 

high reflectivity values observed in Greenland 

(Figure 3.17) are almost independent of the 

cloud cover.

Long-term changes in regional cloud and 

aerosol reflectivity must be considered when 

estimating long-term changes in UV irradiance. 

However, for most populated regions of 

the Earth, long-term (decadal) cloud and 

aerosol scattering changes have been shown 

to be small even where they are statistically 

significant (Herman et al., 2001b, 2008). Local 

values of aerosol amounts and absorption 

are currently estimated from the widely 

distributed AERONET network of groundbased sunphotometers (Holben et al., 2001). 

Other factors, such as Rayleigh scattering 

and surface reflectivity, affect the magnitude 

of measured or theoretically estimated UV 

irradiance. However, these factors do not 

significantly affect the short- or long-term 

changes in irradiance, since their changes are 

small. Hourly or daily changes in Rayleigh 

scattering follow the small changes in 

atmospheric pressure, which usually are less 

than 2%. There have been no long-term changes 

in mean atmospheric pressure. The UV surface 

reflectivity, RG, is small (3 Reflectivity Units, 

RU, to 10 RU, where 1 RU = 1%) and almost 

constant with time except in regions seasonally 

or permanently covered with snow or ice. 

Based on radiative transfer studies, clear-sky 

atmospheric backscattering to the surface 

contributes less than 0.2 RG to the measured 

UV irradiance, which is quite small for most 

ice/snow-free scenes.

3.3.1.1 Reduction of uv by clouds

A measured daily cycle of UV reaching 

the surface will show large UV irradiance 

reductions from clear-sky conditions as 

clouds pass over a site. These reductions 

are frequently in excess of those caused by 

measured ozone changes from climatological 

values for wavelengths longer than 305 nm. In 

general, the effect of clouds is to reduce the UV 

amount at all wavelengths reaching the Earth’s 

surface. The average amount of UV radiation 

reduction by clouds can be estimated from the 

Lambert Equivalent cloud reflectivity R, which 

varies significantly between locations (Figure 

3.17). The operational definition of R is given 

in Appendix 3A.

Satellite data (Figure 3.17) show that the most 

commonly occurring values of R are about 3 

to 5 RU greater than the surface reflectivity 

representing haze or very sparse cloud cover. 

Central Europe, represented by Germany, is 

quite different from North American sites in 

that the most frequent values are around 10 RU 

(127 days out of 3285 days) and around 50 RU 

(128 days), with almost the same number of days 

(80 to 128 days) having 10 to 70 RU. Greenland 

is another extreme, where the reflectivity is 

always high because of the ice cover. Nevada 

and Virginia are similar, except that Nevada has 

a lower average reflectivity representing less 

cloud cover. An extreme case is represented by 
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3.3.1.3 Estimating uv trends: groundbased

Instr umental requirements for making 

long-term UV irradiance measurements are 

well understood in terms of calibration and 

stability for both spectrometers and broadband 

radiometers. While useful work can still be 

done with broadband instruments, much more 

information can be derived from high spectral 

resolution spectrometers (e.g., the global 

network of Brewer spectrometers represented 

in the United States by the NOAA-EPA network 

of single-grating Brewers <http://www.esrl.

noaa.gov/gmd/neubrew>, the National Science 

Foundation/Biospherical network, and at NASA 

by a modified double-grating Brewer (Cede 

et al., 2006). Long-term surface UV spectral 

irradiance measurements must be carefully 

made and analyzed to preclude variations due 

to clouds that could be mixed into UV trend 

estimates, or whose variability can mask the 

detection of small changes. If ground-based 

data are filtered for cloud-free observations, 

then UVB changes caused by changes in 

ozone amount are easily observed in multiyear data records. Aerosols and other forms of 

pollution can also produce apparent changes 

in UV irradiance that mask the effect of ozone 

changes. These can be taken into account if 

measurements are made simultaneously in 

the UVB range (e.g., 305 nm) and outside of 

the ozone absorbing range (e.g., 324 nm). The 

lack of ability to separate aerosol and pollution 

effects from ozone-induced changes limits 

the usefulness of broadband instruments (300 

to 400 nm) for understanding the observed 

irradiance changes. 

Radiometric and wavelength calibration of 

spectrometers used for trend estimates must 

be carefully maintained to detect the relatively 

small changes caused by ozone and aerosols. 

Making accurate spectral measurements is 

quite difficult, since the natural UV spectrum 

at the ground changes by several orders of 

magnitude from 300 to 400 nm. A slight 

wavelength misalignment can cause significant 

errors in the measured UVB irradiance amount. 

Wavelength misalignment is less important 

for integrated quantities such as the erythemal 

irradiance. 

3.3.1.2 UV absorption

The amount of UV reaching the surface can 

also be affected by air pollution, i.e., absorption 

by aerosols (black carbon, dust, and smoke), 

tropospheric O3, NO2, and other gases. These 

can cause reductions in UV of up to 10% in 

polluted sites, but with much higher reductions 

occurring in certain highly polluted cities, e.g., 

occasionally in Los Angeles and frequently 

in Beijing. Nitrogen dioxide causes small 

reductions mainly to UVA since its absorption 

cross-section peaks near 410 nm, but is still 

significant at 330 nm. Aerosols of most types 

have much weaker wavelength dependence 

and affect UV and visible radiation at all 

wavelengths. Pollution abatement, especially 

in highly polluted regions, can decrease the 

atmospheric reflectivity and absorption, which 

has the effect of increasing the amount of UV 

and visible light reaching the ground.

Figure 3.18 Combined effects of ozone, aerosols, and other 

absorbing components on UV radiation. Long-term variability in 

monthly mean solar spectral irradiances at 324 nm (upper panel) 

and at 305 nm (middle panel) measured at Thessaloniki, Greece, 

under clear skies at 63° solar zenith angle, shown as departures from 

the long-term (1990-2006) averages. The lower panel shows the 

corresponding departures in the ozone column of 375 DU (adapted 

from WMO, 2007).
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

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approximately the same amount. 

Combining the changes seen for 

324 nm with those observed for 

305 nm (8.1% per decade) implies 

that the effect of increasing ozone 

amount (0.9% per decade) on 305 

nm irradiance is a statistically 

significant decrease of 3.2% 

(11.3% minus 8.1%) per decade. 

This is also shown in Figure 3.19, 

where the time series for 324 nm, 

D324, was subtracted from the 

time series for 305 nm D305. The 

difference, D305 − D324, was fit 

with a linear regression having a 

slope of –3.4% per decade.

An easy way to check this 

conclusion is t h roug h t he 

radiation amplification factor 

defined as part of Equation 1. The 

radiation amplification factor, 

RAF=−αΩ sec(θ) = −4 for Ω=375 

DU and θ=63°, the average measured values 

for Thessaloniki, Greece. Based on the RAF 

and the observed ozone change of 0.9% per 

decade, the change in 305 nm UV irradiance 

dF/F = RAF dΩ/Ω should be approximately 

–4(0.9) = –3.6% per decade, consistent with the 

measurements of –3.2% and –3.4% per decade 

discussed above. In addition to the smaller 

ozone effects, Figure 3.18 shows that a decline 

in air pollutants can cause increases in surface 

UV irradiance of 11.3% per decade in a local 

industrial site such as Thessaloniki.

When data from cloudy and clear days are 

present in the UV time series, the measured 

trends in UV radiation at individual stations can 

have sufficient variation (typically 0 to 50%, 

and occasionally larger caused by clouds) to 

make estimated long-term trends lose statistical 

significance. As shown in the WMO (2007) 

report, trend estimates for the period from 

1998 through 2005 for Toronto was 1.5±5% 

per decade (1 standard deviation, 1σ) (WMO, 

2007) during a period in which the total ozone 

amount was relatively constant. Even using 

Toronto UV radiation data going back to 1990, 

no statistically significant trend is observable 

in the extended Toronto UV data despite ozone 

decreases that took place during the 1990s, 

because of variability introduced by clouds. To 

A climatology of UV erythemal irradiance for 

the United States and Canada has been derived 

from Brewer and pyranometer data for the 

United States and Canada (Fioletov et al., 2004). 

The ground-based climatology is lower by 10 to 

30% than satellite estimates because of aerosol 

and pollution absorption that are neglected in 

the satellite estimates.

An excellent example of UV trend detection is 

from measured solar irradiances at 305 nm and 

324 nm at Thessaloniki, Greece. The irradiances 

shown in Figure 3.18 are for cloud-free skies at a 

constant solar zenith angle of 63° (WMO, 2007, 

which are an extension of Bais and Lubin et al., 

2007). These data are obtained from a carefully 

maintained Brewer spectrometer located in an 

industrial area of Thessaloniki that is subjected 

to moderate amounts of pollution generated 

both locally and reaching Greece from other 

countries in Europe. There are also occasional 

dust episodes originating in northern Africa.

The radiation at 324 nm should not be 

significantly affected by ozone so that the 

cause of the upward trend at 324 nm (11.3% 

per decade) is almost certainly due to aerosol 

and pollution decreases. Decreasing amounts 

of aerosol and pollution that cause the upward 

trend at 324 nm will also affect 305 nm by 

Figure 3.19 The difference between the 305 nm and 324 nm 

departures from the long-term (1990-2006) averages shown in 

Figure 3.18 showing the measured 3.4% per decade decrease in 

305 nm irradiance caused by an ozone amount increase of 0.9% 

per decade.
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104

surface reflectivity RG, T ≈ (1 – R)/(1 – RG), 

where 0 < T < 1. The irradiance at the surface 

is then approximately 

 FSURFACE = T FCLEAR (2)

The two methods agree quite closely (Krotkov et 

al., 2001), except when there is enough multiple 

scattering within a cloud to give enhanced 

ozone absorption at wavelengths less than about 

310 nm, where CT is the better estimate. Both 

the CT and the simplified method are frequently 

10% higher than measured irradiance values 

on the ground, and sometimes 20% higher. 

The differences are usually caused by an 

underestimate in the satellite calculation of 

aerosol amount and aerosol absorption (Herman 

et al., 1999a; Krotkov et al., 1998; 2001; 

Kalliskota, 2000). The differences become 

much less when the aerosol amount is small or 

is known from ground-based measurements. 

Other sources of difference between groundbased measurements and satellite estimates of 

UV irradiance arise from the large satellite field 

of view (50 x 50 km2 for TOMS and 12 x 24 

km2 for OMI) compared to the smaller groundbased field of view, and also from terrain height 

differences within a satellite field of view. 

A recent comparison of measured UV erythemal 

irradiance from ground-based measurements 

and OMI satellite estimates has been made 

(Tanskanen et al., 2007). The comparison shows 

that for flat, snow-free regions with modest 

loadings of absorbing aerosols or trace gases, 

the OMI-derived daily erythemal doses have a 

median overestimation of 0 to 10%, and 60 to 

80% of the erythemal doses are within ±20% 

compared to ground-based measurements.

Similar errors occur when interpolating between 

widely separated ground-based stations, where 

the aerosol, ozone, and cloud amount varies 

between the stations. Given the need for 

global coverage of UVEST and the sparsely 

located ground-based stations, calculations of 

UVEST from satellite-observed column ozone 

abundances and cloud reflectivities, which are 

validated by ground-based measurements, are 

a useful method for estimating regional, zonal 

average, and global UV irradiance trends.

relate the estimated trends to ozone changes 

requires knowledge of changes in aerosol and 

cloud amounts, which can be obtained from a 

wavelength not affected by ozone.

3.3.1.4 Estimating uv trends: 

satellites

The data for estimating long-term changes of 

surface UV irradiance can come from individual 

local ground-based measurements or from 

global estimations using satellite ozone, aerosol, 

and cloud data. Global estimates of surface UV 

irradiance, UVEST, as a function of latitude and 

longitude have been calculated from satellite 

measurements of atmospheric backscattered 

UV and the small amount reflected from the 

surface. UVEST data are obtained from vector 

radiative transfer calculations that include 

polarization effects, ozone absorption, cloud 

reflectivity and transmission, aerosol scattering 

and absorption, and the measured surface 

reflectivity climatology (Herman and Celarier, 

1997). The long-term precision and stability of 

a satellite instrument’s in-flight calibration, 

especially the single channel radiances used to 

estimate cloud transmission and reflectivity, 

make it very useful for estimating trends in 

UVEST. In the absence of a widely distributed 

closely spaced network of well-calibrated UV 

spectrometers, satellite UV irradiance estimates 

are extremely useful, especially over ocean 

areas where there are no other measurements. 

Since ozone amount, aerosol amount, and 

cloud reflectivity are the measured quantities, 

it is straightforward to separate their respective 

effects on estimated UV irradiance from 

satellite data.

There are two ways of estimating the UV 

irradiance reaching the ground from satellite 

ozone, aerosol, and reflectivity data. First, one 

can enter these quantities in a detailed plane 

parallel radiative transfer model to compute 

cloud transmission, CT, using Mie theory to 

approximate the cloud and aerosol properties 

in addition to Rayleigh scattering and ozone 

absorption (Krotkov et al., 1998; 2001). The 

second, and easier method, is to estimate the 

irradiance reaching the ground for a Rayleigh 

scattering and ozone absorbing atmosphere 

FCLEAR, and then add the cloud and aerosol 

transmission as a correction factor based on 

the measured fractional scene R (0 < R <1) and 
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

Satellite-observed long-term changes in ozone 

amount averaged over the United States suggest 

that there were significant UV changes for both 

erythemal irradiance and for UVB. Compared 

to the annual mean levels in 1980, the change 

in UV averaged over the United States was 

approximately 20% (erythemal irradiance) and 

approximately 40% (305 nm irradiance) early in 

1993. Fortunately, these large percent changes 

were during the winter months when the solar 

zenith angles are large, so that the absolute 

irradiances are comparatively small. The 

calculated annual average irradiance increase 

during 1993 was about 7% and about 14% for 

erythemal and 305 nm irradiances, respectively. 

By 2007, the irradiance increase moderated to 

4% and 8%, respectively, in response to a partial 

recovery of stratospheric ozone, which model 

calculations show is a direct consequence of the 

Note that year-to-year shifts in cyclic weather 

patterns (e.g., clouds, ozone transport, etc.) 

by even a tenth of a degree in latitude and 

longitude (about 10 km) have a minimal effect 

on area-averaged satellite ozone and reflectivity 

measurements (and the UV estimates derived 

from them), but strongly affect ground-based 

UV measurements and their estimates of UV 

irradiance trends. Therefore, the surface UV 

changes deduced from ozone amounts and 

reflectivity measured by satellites, UVEST, are 

expected to be equivalent to those from cloudfiltered ground-based observations of UV 

irradiance, and superior for estimating regional 

and global changes. Satellite measurements 

provide both local and global long-term 

coverage, which can be used to construct zonal 

and regional averages and long-term trends 

that have much less geophysical variance 

from clouds than corresponding groundbased measurements. The use of satellite 

estimates, however, presupposes ground-based 

measurements for validation and as a bridge 

between successive satellite instruments. 

Satellite measures of UVEST have used data 

from Nimbus7-TOMS (N7, 1979 to 1992), 

global weekly averages from multiple SBUV-2 

instruments (1988 to present), global coverage 

from Earth-Probe TOMS (EP, 1997 to 2002), 

and the Aura satellite’s Ozone Monitoring 

Instrument (OMI, 2005 to present). Other data 

are available from European satellites (e.g., 

GOME). 

It has been shown that cloud plus aerosol 

reflectivity over the United States has only 

changed by a small amount for the periods 1980 

to 1992 (Herman et al., 2001b) and for 1997 to 

2007 (Herman et al., 2008), where there are 

well-calibrated satellite reflectivity data records. 

Because of this, the change in UV irradiance 

over the United States can be estimated from 

just the change in satellite measured ozone 

amounts as shown in Figure 3.20. Fioletov et 

al. (2001) has made ground-based estimates of 

erythemal irradiance changes from two Brewer 

spectrometer stations (Montreal and Edmonton), 

and found that the UVB trends were similar to 

those expected from just changes in ozone, but 

with much larger uncertainty because of clouds 

and aerosols.

Barrow, Alaska, has 

experienced UVB 

increases related to 

springtime ozone 

depletion, but these 

increases are a factor 

of ten smaller than 

those observed 

at the southern 

high latitudes.

Figure 3.20 The calculated percent change in UV irradiance caused by percent 

changes in ozone over the continental United States. The ozone change is 

estimated from satellite measurements over the United States.
The U.S. Climate Change Science Program Chapter 3

106

cycles, which become very efficient during 

the 24-hour insolation that occurs in the Arctic 

summer. During summer 2000, two low-ozone 

episodes brought about erythemal UV increases 

on the order of 10 to 15%, each lasting more 

than five days (WMO, 2007 and references 

therein).

Because of the extreme Antarctic springtime 

ozone depletion (ozone hole) compared to all 

other regions, it is useful to compare (Figure 

3.21) the measured amounts of UV irradiance 

at Palmer Station, Antarctica (64°S) with 

San Diego, California (32°N) and Barrow, 

Alaska (71°N). For seasons other than spring 

in Antarctica, there is a decrease in UVB 

irradiance caused by the increased path through 

the atmosphere resulting in less UVB than in 

San Diego. The Antarctic ozone depletion that 

occurs each spring causes the UVB portion 

of the erythemally weighted irradiance to 

increase dramatically to where it exceeds even 

the summertime values observed in San Diego 

at 32°N. Similar wide-area springtime low 

ozone amounts do not occur in the Arctic region 

because of the degree of meteorological wave 

activity in the north that leads to a weaker polar 

vortex and higher ozone amounts.

3.3.3 Human Exposure to UV

From the viewpoint of human exposure to UV, 

the maximum clear-sky UV irradiance and 

exposure occurs in the equatorial zone, 23.3°S 

to 23.3°N, following the seasonal sub-solar 

point, and at high mountain altitudes. In general, 

UV erythemal, UVA, and UVB irradiance 

decreases with increasing latitude outside 

of the equatorial zone, since the maximum 

daily noon solar elevation angle decreases. 

An exception occurs for UVB wavelengths at 

southern mid- to high latitudes when reduced 

ozone amounts from the Antarctic ozone hole 

remain late into the spring and are pushed 

away from Antarctica towards lower latitudes, 

which includes some populated areas. For 

example, UV measurements indicate equatorial 

irradiance levels can occur in the southern part 

of South America for several days.

Global images of daily-integrated UV erythemal 

exposure (kilojoules, kJ per m2

) averaged 

during the months of January (Southern 

Hemisphere summer) and July (Northern 

implementation of the Montreal Protocol and its 

subsequent Amendments.

3.3.2 UV in the Polar Regions

The expansion of the Antarctic polar vortex 

during the 1990s, both in spatial extent and 

temporal into early summer, has increased the 

frequency of elevated UVB episodes over subAntarctic populated areas. These episodes are 

no longer just small pockets of ozone-depleted 

stratospheric air coming from the breakup of the 

polar vortex, but include occasional excursions of 

the polar vortex edge over Ushuaia, Argentina 

and Punta Arenas, Chile. This occurred 44 times 

in the years 1997, 1998, and 2000 combined, with 

some episodes lasting three to four days. Surface 

measurements show average erythemal UV 

increases of about 70% over Ushuaia since 1997, 

and episodic total UVB increases of up to 80% 

over Punta Arenas (WMO, 2007 and references 

therein).

Diaz et al. (2003) show that Barrow, Alaska, 

has experienced UVB increases related to 

springtime ozone depletion in March and April, 

but these increases are a factor of ten smaller 

than those observed at the southern high latitudes. 

Summertime low-ozone episodes in the Arctic 

also affect surface UVB irradiances. These 

summertime events result from gas-phase 

chemistry involving nitrogen and hydrogen 

Figure 3.21 A comparison of measured erythemally weighted UV irradiance in 

Antarctica, the Arctic, and a midlatitude site in relative units. (Fahey, 2007)
107

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

Mountains (over 100 peaks exceeding 7000 

meters in height) during July as shown in 

Figure 3.22. Excluding high altitude locations, 

the largest monthly UV exposures occur in 

Australia and South Africa during summer 

(January) because of their very low amount 

of day-to-day cloud cover from late spring to 

early autumn. Other midlatitude low altitude 

areas also receive high doses, e.g., summertime 

(July) in the southwest United States and the 

Mediterranean countries. 

Other factors contribute to the high Southern 

Hemisphere UV doses. There is a five million 

km decrease in Earth-Sun distance for the 

Southern Hemisphere summer solstices, as 

compared to the Northern Hemisphere, causing 

a 6.5% increase in summer solstice irradiance 

Hemisphere summer), and the two equinox 

months September and March, are shown in 

Figure 3.22 (based on WMO, 1999). Because 

of cloud cover, the high equatorial clear-sky 

irradiances do not translate into the highest 

monthly cumulative exposures. The maximum 

erythemal doses at the equator occur when the 

sun is directly overhead during March, which 

has lower cloud cover than during September. 

The difference is related to the annual cycle of 

the cloud cover associated with the Intertropical 

Convergence Zone (ITCZ), which is usually 

over the equator in September, but is south of 

the equator in March. Two extreme examples 

of very high UV exposures occur in the South 

American Andes (e.g., the sparsely populated 

Atacama Desert in Chile at 4400 to 5600 meters 

altitude) during January and in the Himalayan 

The southwest 

United States 

receives high UV 

exposures during 

summertime.

Figure 3.22 Erythemal exposure kJ per m2 for the months of January, March, July, and September 1988 (from WMO, 

1999) based on Nimbus-7/TOMS ozone and reflectivity data. In terms of the UV index, the numbers would be divided 

by 25. High UV levels are observed over Antarctica in the Southern Hemisphere late Spring and Summer (Figure 3.21). 

These extreme levels are not seen in the September 1988 panel because the sun is just beginning to rise over Antarctica 

and the 1988 ozone depletion was not extreme (Figures 3.5 and 3.7).
The U.S. Climate Change Science Program Chapter 3

108

Through the 

Montreal Protocol 

and subsequent 

agreements, CFCs 

were almost 

completely phased 

out by 1995, 

preventing major 

chlorine-driven 

ozone decreases 

and UVB increases.

the annual average clear-sky UV erythemal 

irradiance averaged over the continental United 

States increased from 1979 to the mid-1990s by 

about 7%. Since the mid-1990s the erythemal 

irradiance has decreased, so that the current 

level is about 4% higher than it was at the start 

of the record in 1979. Year-to-year and seasonal 

variations ranged from only a few percent to 

about 20% with the largest changes occurring 

during the winter months when UV irradiance 

is at an annual minimum. In the absence of 

the Montreal Protocol, summer maximum 

and annually integrated UVB doses over the 

United States would have been much larger 

with adverse consequences for public health 

and ecosystems (U.S. EPA, 1999). 

Ground-based measurements of surface UV 

trends present a challenge that can be overcome 

with proper analysis of the data for cloud-free 

conditions along with simultaneous aerosol 

measurements. UV estimates from satellite 

measurements of ozone, aerosols, and cloud 

reflectivity are averages over large areas on 

the order of 25 km to 100 km, which minimizes 

many problems with local variability of cloud 

and aerosol amounts. Both ground and satellite 

UV estimates are critically dependent on 

establishing and maintaining an accurate 

calibration over the lifetime of an instrument 

and between successive instruments. Groundbased measurements are essential to provide 

validation of satellite calibration and as a bridge 

between successive satellite instruments.

While the UV irradiance maximum in 1993 

was associated with the massive equatorial 

Mt. Pinatubo eruption in 1991, a portion of 

the total increase occurred before 1991 and 

was associated with ozone destruction from 

chlorine loading in the atmosphere before 

being limited by the Montreal Protocol. Major 

chlorine-driven ozone decreases and UVB 

increases were prevented by this and subsequent 

agreements that were effective for limiting 

releases of chloroflourocarbons (CFCs) and 

other chlorine-bearing compounds, with CFCs 

being almost completely phased out by 1995.

in the Southern Hemisphere. Average summer 

ozone in the Southern Hemisphere (270 DU) 

is lower than the Northern Hemisphere (320 

DU) by about 13%, which would lead to a 13% 

increase in 310 nm and a 26% increase in 305 

nm irradiance. The exact percent increase is a 

function of latitude. In general, the Southern 

Hemisphere has less pollution aerosols, which 

can cause another few percent increase in UV 

irradiance relative to Northern Hemisphere.

In Australia and South Africa, the combination 

of high UV exposure and residents of European 

descent have lead to a major skin cancer health 

problem. Based on National Institutes of Health 

data, the same problems are present in the United 

States, with more skin cancer occurring at lower 

latitudes where the UV exposure is higher. 

The seriousness of the very high UV exposure 

problem is observed in Australia, where skin 

cancer rates have increased dramatically (20% 

for basal cell, to 788 per 100,000; over 90% for 

squamous cell, to 321 per 100,000 carcinomas), 

based on household surveys conducted in 1985, 

1990, and 1995 (Staples et al., 1998). This 

compares to the U.S. National Cancer Institute 

estimate of 14.5 per 100,000 for the United 

States. Lucas et al. (2006) give a comprehensive 

review of health problems and benefits related 

to UV exposure.

3.3.4 UV Summary

Measurements from ground-based instruments 

at different midlatitude sites around the globe 

show a mixture of UVB increases and decreases 

that depend on changes in local cloud cover, 

ozone, and aerosol amounts. Trends in UV 

in the polar regions, especially Antarctica, 

are dominated by changes in springtime 

stratospheric ozone. In the latitude range 60°S 

to 60°N, all three main factors governing UVB 

must be taken into account (for UVA, clouds 

and aerosols are the dominant factors). Groundbased stations located in or near urban sites 

have observed increases in cloud-free sky UV 

radiation from pollution abatement comparable 

to those from observed total column ozone 

changes. 

Measurements of ozone and cloud plus aerosol 

reflectivity from satellites have been used to 

estimate the changes in UVB over the last 

28 years. Based on the satellite ozone record, 
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Layer Recovery, and Implications for Ultraviolet Radiation Exposure

APPENDIX 3A: Lambert Equivalent Reflectivity (LER)

The Lambert Equivalent Reflectivity, R, is calculated by requiring that the measured radiance, 

ISM, match the calculated radiance, IS, at the observing position of the satellite (Equation A1) by 

solving for the single free parameter, R, in the formal solution of the radiative transfer equation

I R P

RI P f P

RS P

I P I S O

d O O

b O

dO O SM ( , , , ) ( , , ) ( , , )

( , ) Ω Θ ( , , )

Ω Θ Ω Θ

Ω = Ω Θ − + = 1 (A1)

 

where Ω = ozone amount from shorter wavelengths (e.g., 317 nm) 

Θ = viewing geometry (solar zenith angle, satellite look angle, azimuth angle)

R = LER at PO 0 < R < 1 

PO = pressure of the reflecting surface (e.g., ground or cloud)

Sb = fraction scattered back to PO from the atmosphere

Id = sum of direct and diffuse irradiance reaching PO

f = fraction of radiation reflected from PO reaching the satellite 

IdO = radiance scattered back from the atmosphere for R=0 and P=PO

The quantities Sb, Id, f, and IdO are calculated from a radiative transfer solution and stored in 

tables. From Equation A1,

R

I I

I f I I S

SM dO

d SM dO b

= −

+ − ( ) (A2)

APPENDIX 3B: UV Index and Units

Erythemal irradiance is frequently expressed in terms of the UV index = 25 milliWatts (mW) 

per m2 = 2.5 microWatts (μW) per cm2 (the units of Figure 3.14). The index is an arbitrary unit 

such that very high values reported by weather services have a UV index of 10. In Figure 3.14, 

the highest value is about 22 μW per cm2, which is a UV index of 8.8. High altitude locations 

with extreme UV amounts can exceed 10 on clear days. Erythemal exposure or dose is a timeintegrated quantity normally expressed in kJ per m2

.
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How Do Climate Change and 

Stratospheric Ozone Loss Interact?

CHAPTER

4

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Layer Recovery, and Implications for Ultraviolet Radiation Exposure

Convening Lead Author: David W. Fahey, NOAA

Lead Authors: Anne R. Douglass, NASA; V. Ramaswamy, NOAA; 

Anne-Marie Schmoltner, NSF

KEY ISSUES

Stratospheric ozone abundances are dependent on a balance of chemical processes that both 

produce and destroy ozone and dynamical processes that transport ozone throughout the 

stratosphere. The chemical processes depend on atmospheric temperatures and the abundances 

of ozone-depleting substances and other trace gases, such as water vapor and nitrogen oxides. 

Transport depends on heating in the atmosphere, which also depends on the distribution and 

abundance of trace gases, such as carbon dioxide, ozone-depleting substances, and ozone. 

Atmospheric temperature, transport, and trace gas amounts, for example, are all aspects of 

Earth’s climate. As these and other climate parameters change as a result of human activities 

and natural variability, ozone abundances will decrease or increase in a manner that depends 

on a variety of factors. 

This complex coupling of ozone and climate parameters is not fully defined at present and has 

significant uncertainties associated with known key aspects. Chemistry climate models (CCMs) 

of the atmosphere are in development and use by researchers aiming to reduce the uncertainty 

in the ozone-climate interaction and to explore other aspects of the interrelationship. Key 

questions related to the coupling of ozone and climate are:

• How do ozone-depleting substances and ozone depletion contribute to the radiative forcing 

(RF) of climate?

• How do long-term changes in greenhouse gases affect stratospheric ozone?

• How have stratospheric temperatures changed in recent decades and what are the causes 

of these changes?

• Is stratospheric water vapor changing in a way that influences ozone abundances?

• How do ozone changes influence the climate of the stratosphere and troposphere?

• Will volcanic eruptions influence future ozone amounts?
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Linking of ozone and climate change

• Ozone and climate change are linked because both ozone and ozone-depleting substances 

(ODSs) are greenhouse gases, which contribute to radiative forcing. The ODS contribution to global radiative forcing between 1750 and 2005 is approximately 20 percent (+0.34 

W per m2) of that from carbon dioxide, the largest human contribution. The ODS contribution is expected to decline in coming decades as ODS emissions and their atmospheric 

abundances continue to decline in the atmosphere. 

• Each ODS contributes to ozone depletion and climate warming with a different level of 

effectiveness as represented, for example, by the Ozone Depletion Potentials (ODPs) and 

Global Warming Potentials (GWPs), respectively. For the principal ODSs, these values 

vary over orders of magnitude for equal mass of emissions. 

• The abundance of stratospheric ozone is dependent on a balance of production and loss 

processes. These processes are dependent on several features of the atmosphere, namely, 

its chemical composition, air motions, radiation, and temperatures. Climate change will 

lead to changes in these features, which in turn will affect ozone abundances. Climate 

change has the potential to increase or decrease ozone abundances depending on the 

region and extent of climate change.

• Chemical ozone depletion also contributes to climate change by modifying atmospheric 

radiative properties. Ozone losses can also alter atmospheric temperatures and atmospheric transport. Ozone depletion can affect the climate of both the troposphere and 

stratosphere. The ozone depletion due to ODSs is expected to offset the direct climate 

forcing by ODSs. The extent of the offset is uncertain, with an estimated value of –0.05 W 

per m2

 with an uncertainty range of –0.15 to + 0.05 W per m2.

Impact of climate change on ozone

• The complexity of the interactions of ozone changes with climate parameters requires development and evaluation of coupled models of Earth’s atmospheric chemistry and climate 

processes (called chemistry climate models, CCMs). These CCMs are used to predict 

future ozone amounts. In addition, CCMs are needed to evaluate the sensitivity of ozone 

to climate parameters and the response of climate to ozone changes.

• Global average stratospheric temperatures have decreased in the observational records 

that begin in the 1960s. The decrease is attributed mainly to ozone depletion, increased 

carbon dioxide (CO2), and changes in water vapor. 

• Stratospheric temperatures influence ozone amounts through chemical and transport 

processes. Future increases in CO2 will continue to contribute to global stratospheric 

cooling. The photochemical loss of ozone is slowed in some regions when temperatures 

decrease with the result that ozone recovery may be accelerated.

• Human activities are expected to increase the future abundances of greenhouse gases 

that influence stratospheric ozone amounts, principally, CO2, methane (CH4) and nitrous 

oxide (N2O).

• Stratospheric water vapor has increased in recent decades but since 2001 has been decreasing in the lower stratosphere. The oxidation of methane emissions is an important 

contributor to increasing water vapor trends. Tropical tropopause temperatures modulate dehydration of air entering the stratosphere, and recent decreases in water vapor are 

well correlated with negative tropical tropopause anomalies. Future water vapor trends 

are uncertain because of uncertainties in projecting methane emissions and the temperatures of the tropical tropopause. Stratospheric water vapor influences stratospheric ozone 

through reactive hydrogen chemistry and polar stratospheric cloud formation.

KEY FINDINGS
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

• Chemistry climate model simulations predict that the atmospheric circulation between 

the stratosphere and troposphere will increase in a changing climate in the coming decades. Increased circulation will change stratospheric ozone amounts and increase the 

stratospheric flux of ozone to the troposphere.

Impact of changes in stratospheric ozone on climate change 

• Depletion of stratospheric ozone since about 1980 has caused a negative radiative forcing 

of climate change (approximately –0.05 W per m2, with an uncertainty that encompasses 

a range from −0.15 to +0.05 W per m2). The forcing is a balance between a shortwave 

cooling of the lower stratosphere and a longwave cooling below the region of ozone depletion. In comparison, increases in ozone from pollution chemistry in the troposphere have 

caused a positive radiative forcing (approximately +0.35 W per m2).

• Ozone depletion causes changes to the temperature and circulation of the stratosphere 

and troposphere. Observational analyses indicate that stratospheric ozone depletion over 

Antarctica has strengthened circumpolar flow throughout the troposphere over Antarctica and caused surface temperature changes. This effect has been well simulated using 

many different general circulation models (GCMs).

Importance of volcanic eruptions 

• If volcanic eruptions that inject material into the stratosphere, referred to here as explosive eruptions, occur again in the coming decades, they will decrease stratospheric 

ozone levels for several years as a result of the heterogeneous reactions occurring on 

volcanic sulfate aerosols. These reactions increase halogen loss processes by reducing the 

abundance of key reactive nitrogen compounds. For a given eruption size, the resulting 

enhanced ozone destruction from halogens will decrease as halogen amounts in the atmosphere decrease in the coming decades.

• Explosive volcanic eruptions are expected to cause major temperature and circulation 

changes in the stratosphere as have occurred after past eruptions. These changes are in 

response to the large increases in sulfate aerosol amounts in the stratosphere following 

such eruptions. The increases result in only a short-term shift in stratospheric climate 

because natural processes remove most of the additional sulfate aerosols within two to 

three years after the eruption. 
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Chapter 4

In the natural 

atmosphere, ozone’s 

warming of the 

planet makes it the 

third most important 

longwave greenhouse 

gas after water vapor 

and carbon dioxide.

4.1 Introduction

Ozone occurs naturally in the atmosphere as 

a result of photochemical processes. In the 

stratosphere, ozone is beneficial to life on Earth 

because it absorbs ultraviolet radiation (UV) 

from the sun. Ultraviolet light absorption heats 

the stratosphere. Ozone is also a greenhouse 

gas that helps trap terrestrial infrared radiation, 

which leads to heating of the troposphere 

and stratosphere. In the natural atmosphere, 

ozone’s warming of the planet makes it the 

third most important longwave greenhouse gas 

after water vapor and carbon dioxide (Kiehl 

and Trenberth, 1997). As a consequence, 

changes in ozone amounts have the potential to 

change climate parameters in the stratosphere 

and troposphere. Anthropogenic pollution 

has led to increased ozone production and 

abundances in the troposphere, particularly 

near Earth’s surface. In contrast, emissions of 

ozone-depleting substances (ODSs) in recent 

decades have led to significant depletion of 

global stratospheric ozone, with particularly 

high losses in polar regions. The Montreal 

Protocol has been established to protect the 

ozone layer by reducing the global production 

and consumption of ODSs.

The complex interrelationship between ozone 

and climate change is illustrated in Figure 4.1. 

Multiple radiative, chemical, and dynamical 

processes control ozone amounts and their 

distribution in the troposphere and stratosphere. 

Production and loss cycles of ozone involve 

many chemical species, as well as aerosols, and 

are influenced by atmospheric parameters such 

as solar insolation and temperature. Natural 

processes and human activities influence ozone 

through changes in atmospheric composition 

and climate parameters. The chemical loss rate 

of ozone leads to an atmospheric lifetime that 

is relatively short compared to carbon dioxide, 

for example. As a result, dynamical processes 

such as planetary waves and the Brewer-Dobson 

circulation have important roles in establishing 

the observed non-uniform distribution of ozone 

in the atmosphere. 

Figure 4.1 Schematic depiction of how climate change and ozone abundances are linked to each other in the 

stratosphere and troposphere and to natural and human activities at Earth’s surface (Isaksen, 2003.)

Connections between Climate and Ozone
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

The ODSs are also greenhouse gases. The 

radiative effect of accumulated ODS emissions 

is partially offset by the reduction in global 

ozone amounts caused by these emissions (see 

Section 4.2). Systematic and long-term ozone 

depletion can change atmospheric circulation 

patterns and contribute to climate change. 

Furthermore, changes in climate can alter ozone 

amounts. Changes in temperature, amounts 

of trace gases such as methane, nitrous oxide, 

and water vapor, and atmospheric circulation 

can all potentially lead to ozone changes in 

the stratosphere and troposphere. Finally, 

large volcanic emissions can alter both ozone 

and climate for temporary periods of several 

years.

This chapter assesses these interconnections, 

schematically shown in Figure 4.2, by outlining 

what is known about ODSs and volcanic 

emissions and the processes through which 

ozone influences climate and through which 

climate change will influence ozone amounts. 

Further detail on the coupling of climate and 

ozone changes can be found in recent scientific 

assessments (IPCC/TEAP, 2005; WMO, 

2007).

4.2 Radiative Forcing of

Climate by Ozone-depleting 

Substances and Ozone

Changes

Ozone and ODSs are greenhouse gases. 

The accumulation of ODS emissions in the 

atmosphere causes a direct radiative effect in a 

manner similar to that from emissions of CO2, 

the leading anthropogenic greenhouse gas. 

ODS emissions also cause an indirect radiative 

effect, because ODSs cause the destruction of 

stratospheric ozone. These radiative effects are 

discussed in the following sections. 

4.2.1 Direct Radiative Forcing by 

Ozone-Depleting Substances

The accumulation of ozone-depleting 

substances in the atmosphere contributes to 

the direct radiative forcing of climate. The ODS 

efficiencies as ozone-depleting substances (i.e., 

ODPs) and as greenhouse gases (i.e., GWPs) 

are contrasted in Figure 4.3 (also see Box 4.1). 

A large range is found in both metrics for the 

principal gases. The continuous measurements 

of ODS abundances in the atmosphere over 

the last two to three decades allow an accurate 

evaluation of their contributions to ozone 

depletion and climate change. Projections of 

emissions allow the future contribution of ODSs 

to ozone depletion and climate change to be 

estimated. ODP-weighted and GWP-weighted 

The accumulation 

of ozone-depleting 

substances in 

the atmosphere 

causes a direct 

radiative effect, 

similar to that from 

carbon dioxide.

Figure 4.2 Schematic of specific processes addressed in this assessment that interconnect and 

influence atmospheric ozone amounts and climate parameters.

Connections between Climate and Ozone
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Chapter 4

long-lived greenhouse gases and 21 percent of 

the total anthropogenic RF in 2005. 

ODSs account for 94 percent (+0.32 W per 

m2) of the halocarbon term in Figure 4.5. The 

balance (6 percent) is due to the accumulation 

of hydrofluorocarbon (HFC) emissions, which 

are included in the Kyoto Protocol (UNFCCC, 

1997). Emissions of HFCs, which are substitutes 

for ODSs in many applications, are increasing. 

HFCs do not deplete ozone (ODP = 0) but can 

have substantial GWPs (Figure 4.3).

A comprehensive evaluation of the protection 

of climate afforded by increases or reductions 

in ODS emissions must take into account 

two indirect effects or offsets. The case of 

reduced ODS emissions from actions under 

the Montreal Protocol were analyzed recently 

by Velders et al. (2007). As ODS emissions 

are reduced, global stratospheric ozone levels 

are restored from their depleted state. Since 

ozone is a greenhouse gas, ozone RF increases 

as ozone levels are restored, thereby offsetting 

the reductions in RF from ODS reductions. 

The second offset is the increase in emissions 

of HFCs, all potent greenhouse gases, which 

ODS emissions grew substantially in recent 

decades but peaked in the late 1990s, as shown 

in Figure 4.4 (IPCC/TEAP, 2005; Velders et 

al., 2007). The decline is in response to the 

provisions of the Montreal Protocol, which 

requires a staged phase-out of all principal ODS 

use in developed and developing nations. The 

direct radiative forcing (RF) contribution from 

ODSs likely would have been approximately 

twice as large in 2010 in the absence of the 

Montreal Protocol or other regulation (Figure 

4.4, right panel) (Velders et al., 2007). 

The radiative forcing of individual ODSs 

varies because of differences in emissions, 

lifetimes, and radiative efficiencies. The RF 

values attributable to individual ODSs for the 

period 1970 to 2000 are shown in Table 4.1 

along with values for CO2 and CH4. CFCs as 

a group form the largest contribution to RF 

amongst all ODSs. A comparison of the RF 

from halocarbon gases as a group with values 

associated with other aspects of natural and 

anthropogenic climate forcing is shown in 

Figure 4.5. The RF from halocarbon gases in 

2005 is +0.34 [±0.03] W per m2 (warming), 

which represents 13 percent of the RF from all 

Figure 4.3 Comparison of the Ozone Depletion Potentials (ODPs) and Global Warming Potentials (GWPs) for principal 

ozone-depleting substances (ODSs) and hydrofluorocarbons (HFCs) (see Box 4.1). The contributions of emissions to ozone 

depletion and climate change increase with the ODP and GWP values, respectively. HFCs are ODS substitute gases that do not 

destroy ozone (i.e., ODP = 0). The comparison is for emissions of equal mass. The GWPs are evaluated for a 100-year period 

after emission. The ODPs of CFC-11 and CFC-12, and the GWP of CO2 are defined to have values of 1.0 (Daniel et al., 1995; 

IPCC/TEAP, 2005; WMO, 2007).
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Layer Recovery, and Implications for Ultraviolet Radiation Exposure

BOX 4.1: Ozone Depletion Potential and Global Warming Potential1

Ozone Depletion Potential: Ozone-depleting substances are compared in their effectiveness to destroy 

stratospheric ozone using the “Ozone Depletion Potential” (ODP), as shown in Figure 4.3. A gas with a larger 

ODP has a greater potential to destroy ozone over its lifetime in the atmosphere. The ODP is calculated 

on a “per mass” basis for each gas relative to CFC-11, which has an ODP defined to be 1. Halon-1211 and 

halon-1301 have ODPs significantly larger than CFC-11 and most other emitted gases, because bromine is 

about 60 times more effective overall on a per-atom basis than chlorine in chemical reactions that destroy 

ozone in the stratosphere. The gases with small ODP values generally have short atmospheric lifetimes 

or fewer chlorine and bromine atoms. The production and consumption of all principal ozone-depleting 

substances (ODSs) in human activities are regulated under the provisions of the Montreal Protocol. 

Global Warming Potential: The climate impact of a given mass of a halocarbon emitted to the atmosphere 

depends on both its radiative properties and its atmospheric lifetime. The two can be combined to compute 

the Global Warming Potential (GWP), which is a proxy for the climate effect of a gas relative to the 

emission of a pulse of an equal mass of CO2. Multiplying emissions of a gas by its GWP gives the emission 

of that gas over a given time horizon. A value of 100 years is often chosen as a reference time horizon for 

intercomparisons of GWPs.

GWPs are most useful as relative measures of the climate response due to direct radiative forcing of wellmixed greenhouse gases whose atmospheric lifetimes are controlled by similar processes, which includes 

most of the halocarbons. 

1

Adapted from WMO (2007) and IPCC/TEAP (2005).

Figure 4.4 ODP-weighted emissions (left panel), GWP-weighted emissions (100-year) (middle panel), and radiative forcing (right 

panel) for ODS and CO2 scenarios for 1960–2020. Four scenarios are used: (1) the baseline which represents ODS observations to 

date and projections for the future (black line); (2) the emissions that plausibly would have occurred in absence of the Molina and 

Roland warning that ODSs deplete ozone (MR74) (green shaded and striped region), (3) the emissions that plausibly would have 

occurred in absence of the implementation of the Montreal Protocol (NMP87) (blue shaded and striped region), and (4) the IPCC 

SRES scenario for CO2 emissions beyond 2003 (red line). ODS emissions are normalized by their direct GWPs to form units of 

equivalent gigatons (Gt) CO2 per year. The shaded regions reflect uncertainties in projecting ODS growth rates in the MR74 and 

NMP87 scenarios. The striped regions indicate larger uncertainties are associated with the scenarios past 2010. The CO2 emissions 

for 1960–2003 are from global fossil fuel and cement production. All RF values represent net changes from the start of the industrial 

era (1750) to present (2005). The magnitude of the reduction target of the first commitment period of the Kyoto Protocol (2008 to 

2012) is shown in the middle panel for reference. The target represents average emission reductions in six key greenhouse gases by 

Annex-1 countries (5.8 percent below 1990 values) plus the expected growth in emissions between 1990 and 2008 to 2012 (adapted 

from Velders et al., 2007).
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Chapter 4

is intrinsically tied to ODS reductions because 

HFCs are key substitute gases for ODSs. Thus, 

the net gain from reducing the RF contribution 

of ODSs must include negative offsets due to the 

reversal of some ozone depletion and increased 

abundances of other greenhouse gases. In 2010, 

these factors were estimated to offset about 

25 percent of the RF decrease attributable to 

reductions of ODSs under the Montreal Protocol 

since 1990 (Velders et al., 2007). There is a 

large uncertainty in this offset in radiative 

forcing (see below). As ODS abundances 

continue to decline in the atmosphere after 

2010, the relative size of the ozone offset is 

likely to remain unchanged while the HFC 

offset might increase depending on growth 

in production and use of HFCs, which are not 

regulated by the Montreal Protocol.

Figure 4.5 Radiative forcing values for the principal contributions to climate change from anthropogenic activities and natural 

processes. Each numerical value listed and indicated with a bar is a global mean value representing the best estimate of the change 

between preindustrial times (ca, 1750) and the present (2005). The error bars indicate the uncertainty ranges. The spatial scale 

and level of scientific understanding (LOSU) is also indicated for each value (adapted from IPCC, 2007, Figure SPM-2).

1

 For accumulated emissions in the period 

1970-2000. (Adapted from IPCC/TEAP, 2005, 

Table TS-1.)

Table 4.1 Direct radiative forcing of CO2, 

CH4, and principal ODSs1

Gas Radiative Forcing 

(W per m2

)

CO2 0.67

CH4 0.13

N2O 0.068

CFC-11 0.053

CFC-12 0.136

CFC-113 0.023

CFC-114 0.003

CFC-115 0.002

HCFC-22 0.0263

HCFC-141b 0.0018

HCFC-142b 0.0024

Halon-1211 0.0012

Halon-1301 0.0009

CCl4 0.0029
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4.2.2 Radiative Forcing 

From Ozone Changes

Stratospheric ozone depletion and increases 

in tropospheric ozone both contribute to the 

radiative forcing of climate. Stratospheric ozone 

depletion is an indirect radiative effect of ODS 

emissions. The response of surface climate to 

ozone changes is complex, in general, because 

it depends on the balance between shortwave 

and longwave radiative effects. For example, 

when ozone is increased in the troposphere 

or lower stratosphere, surface temperatures 

tend to increase due to increased longwave 

forcing (Forster and Shine, 1997; IPCC/

TEAP, 2005) (see Box 4.2). Overall, surface 

temperatures are most sensitive to changes in 

ozone concentrations near the tropopause. 

Stratospheric ozone depletion has occurred 

primarily at extratropical latitudes with 

substantially larger changes in the Southern 

Hemisphere. Southern Hemisphere ozone 

values over the period from 2000 to 2003 are 

on average 6 percent below pre-1980 values, 

while Northern Hemisphere values are 3 

percent lower. The net RF change from these 

The radiative forcing is defined as the change in net radiative flux at the tropopause after allowing for stratospheric 

temperatures to readjust to radiative equilibrium (IPCC, 2001, Chapter 6). The details of this definition are crucial 

for stratospheric ozone, and are explained in the figure.

Box Figure 4.2 The instantaneous effect of stratospheric ozone depletion (left-hand side of schematic) is to increase the 

shortwave radiation from the Sun reaching the tropopause (because there is less ozone to absorb it), and to slightly reduce 

the downward longwave radiation from the stratosphere, as there is less ozone in the stratosphere to emit radiation. This 

gives an instantaneous net positive radiative forcing. However, in response to less absorption of both shortwave and longwave 

radiation in the stratosphere, the region cools, which leads to an overall reduction of thermal radiation emitted downward 

from the stratosphere (right-hand side of schematic). The size of this adjustment term depends on the vertical profile of 

ozone change and is largest for changes near the tropopause. For the observed stratospheric ozone changes the adjustment 

term is larger than the positive instantaneous term, thus the stratospheric ozone radiative forcing is negative.

1

Excerpted from IPCC/TEAP (2005).

BOX 4.2: Why is the radiative forcing of stratospheric ozone negative?1
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Chapter 4

Stratospheric 

ozone depletion is 

estimated to cause 

a net reduction in 

radiative forcing at 

the tropopause and 

lead to a cooling effect 

on the atmosphere.

observed depletions has been assessed by the 

Intergovernmental Panel on Climate Change 

(IPCC) to be −0.05 [±0.1] W per m2

 (cooling)

(Figure 4.5) (IPCC, 2007; Hansen et al., 2005). 

The instantaneous response to ozone depletion 

in the lower stratosphere is a positive forcing 

because solar flux significantly increases below 

the tropopause and downwelling longwave 

radiation decreases slightly. In addition, less 

solar and longwave radiation is absorbed in 

the lower stratosphere when ozone amounts are 

reduced, thereby cooling the region and further 

reducing the downwelling longwave flux from 

ozone and other gases. The calculation of the 

net change in forcing is associated with large 

uncertainties because net cooling is sensitive 

to the location where ozone depletion occurs 

(see Box 4.2). When all effects and feedbacks 

are taken into account, stratospheric ozone 

depletion is estimated to cause a net reduction 

in RF at the tropopause and lead to a cooling 

effect on the atmosphere.

Stratospheric ozone depletion is included in 

the IPCC best estimate of the RF from global 

stratospheric ozone changes that have occurred 

between 1750 and 2005 (IPCC, 2007). The 

net global ozone RF estimates used in the 

IPCC include the possibility of a positive 

RF contribution due to a modeled increase 

in stratospheric ozone in some regions since 

preindustrial times, despite losses related to 

ODSs. The RF value is particularly sensitive to 

ozone changes in the tropical lower stratosphere, 

which are small compared to changes at high 

latitudes. Observational and modeling studies 

indicate how Northern Hemisphere ozone 

amounts are also influenced by changes in 

atmospheric dynamics, such as changes in 

tropopause heights (Pyle et al., 2005), in 

addition to increased amounts of ODSs. The 

model results show large differences in how 

ozone column amounts have responded at 

midlatitudes to changes in ODSs and other 

parameters, such as circulation (Gauss et al., 

2006).

The total RF associated with ODS emissions 

is the sum of three effects: the direct warming 

summed over the accumulation of the 

contributing ODS gases (+0.34 [±0.03] W per 

m2

); the direct warming from HFC emissions 

that occurred as a result of ODSs reductions; 

and the indirect cooling associated with 

stratospheric ozone changes (−0.05 [±0.1] W 

per m2

). The warming due to HFCs is currently 

small (approximately +0.02 W per m2) but is 

expected to grow in the future as ODS warming 

declines (Velders et al., 2007). The uncertainty 

in the indirect effect is large compared to the 

direct effects. As previously discussed, in 

some cases the indirect effect as evaluated 

in atmospheric models includes effects not 

related to ODS emissions. In addition, the 

regional distribution of the radiative forcing 

from ODSs generally is different from that 

due to ozone changes. Thus, there are currently 

large uncertainties in the magnitude of global 

ozone cooling effects and the degree to which 

they offset the warming due to ODSs, as well 

as where and how these offsets would vary 

regionally. 

Tropospheric ozone since preindustrial times 

has increased as a result of increased emissions 

of human pollutants, primarily nitrogen oxides, 

carbon monoxide, and organic compounds, 

including methane. Photochemical and radiative 

transfer models are used to calculate ozone 

changes and the associated RF, respectively. 

The changes include the net transport of ozone 

from the stratosphere to the troposphere, 

which can be altered by climate change and 

stratospheric ozone depletion. The tropospheric 

ozone RF (+0.35 W per m2

) from human 

activities is larger than the stratospheric ozone 

term and associated with large uncertainties 

(Figure 4.5).

4.3 The Response of

Ozone to Climate Change

Parameters

Ozone responds to climate change parameters in 

a variety of ways because ozone is continually 

photochemically produced and destroyed in 

the atmosphere and thus dependent on the 

abundance of other gases emitted by natural 

processes and human activities. The complexity 

of the interaction of ozone with climate change 

parameters (Figure 4.1) requires the use of 

chemistry climate models (CCMs) (see Box 4.3) 

to diagnose the sensitivity of ozone to climate 

change parameters and to predict future ozone 

amounts in a changing climate.
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4.3.1 Calculating the Response 

of Ozone to Climate Change 

Parameters with CCMs

The approach to CCM use is schematically 

shown in Figure 4.6. Transport, radiation, 

dynamics, and chemistry and microphysics are 

the four principal aspects of a CCM. A CCM 

requires as input specific knowledge of natural 

process and their trends, such as in emissions, 

solar irradiance, and volcanic eruptions; and of 

human activities and their trends, primarily for 

emissions. These inputs define and constrain the 

current and future state of climate parameters. 

The CCM output includes a wide array of 

parameters and diagnostics in addition to ozone 

that can be compared to observations and other 

models.

Chemistry climate models are complex 

because simulating the atmosphere requires 

interdependence and interaction between the 

core aspects of the model. Important examples 

include the coupling between transport and 

radiation. Transport depends in part on 

atmospheric temperature gradients that are 

established by the distribution of radiative 

heating. Radiative heating is determined, in 

part, by long-lived greenhouse gases and ozone. 

Photochemical reaction rates also depend on 

ambient temperatures. Thus, the photochemical 

balance controlling the abundances of ozone 

and other species depends substantially on the 

atmospheric abundances of greenhouse gases. 

Sea surface temperatures, land-sea temperature 

differences, and other factors influence wave 

propagation into the stratosphere, thereby 

affecting meridional transport rates. These 

couplings are discussed in more detail in WMO 

(2007) and Eyring et al. (2005).

The validation of CCM output for ozone and 

other parameters has become a focus topic 

because of the heightened need to project future 

ozone abundances with reliable uncertainty 

estimates (Eyring et al., 2005; 2006). Reasonable 

agreement is found between many CCMs and 

global ozone trends, but for polar ozone trends 

the CCMs show a large spread in results. 

Uncertainties in CCM results reflect limitations 

in our understanding of how to represent 

atmospheric processes and their feedbacks in 

model simulations and, therefore, limit the 

precision and accuracy of our projections of 

future ozone amounts and the influence of 

climate change.

Atmospheric General Circulation Model (AGCM): A three-dimensional model of large-scale (spatial 

resolution of a few hundred km) physical, radiative, and dynamical processes in the atmosphere over years 

and decades. An AGCM is used to study changes in natural variability of the atmosphere and for investigations 

of climate effects of radiatively active trace gases (greenhouse gases) and aerosols (natural and human-made), 

along with their interactions and feedbacks. Usually, AGCM calculations employ prescribed concentrations of 

radiatively active gases, e.g., carbon dioxide (CO2), methane (CH4), nitrous oxide (N2O), chlorofluorocarbons 

(CFCs), and ozone (O3). Changes of water vapor (H2O) concentrations due to the hydrological cycle are 

directly simulated by an AGCM. Sea surface temperatures (SSTs) are prescribed. An AGCM coupled to an 

ocean model, commonly referred to as an AOGCM or a climate model, is used for investigation of climate 

change. More recently, climate models may also include other feedback processes (e.g., carbon cycle, interaction 

with the biosphere). 

Chemistry Climate Model (CCM): An AGCM that is interactively coupled to a detailed chemistry module. In 

a CCM, the simulated concentrations of the radiatively active gases are used in the calculations of net heating 

rates. Changes in the abundance of these gases due to chemistry and advection influence heating rates and, 

consequently, variables describing atmospheric dynamics such as temperature and wind. This gives rise to a 

dynamical-chemical coupling in which the chemistry influences the dynamics (via radiative heating) and vice 

versa (via temperature and advection). Not all CCMs have full coupling for all chemical constituents; some 

radiatively active gases are specified in either the climate or chemistry modules. Ozone is always fully coupled, 

as it represents the overwhelmingly dominant radiative-chemical feedback in the stratosphere. 

2

Excerpted from WMO (2007).

BOX 4.3: Models Used to Study Climate Processes2

A chemistry climate 

model requires 

as input specific 

knowledge of natural 

processes and their 

trends, such as in 

emissions, solar 

irradiance, and 

volcanic eruptions. 

It also requires 

knowledge of human 

activities and their 

trends, primarily 

for emissions.
The U.S. Climate Change Science Program

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Chapter 4

Stratospheric 

temperature has 

decreased by about 

0.5 K per decade in 

the lower stratosphere 

and by about 1 to 2 

K per decade in the 

middle stratosphere.

4.3.2 Stratospheric 

Temperature Changes

Stratospheric temperatures have decreased over 

the last three to four decades. Observations from 

satellites beginning in 1979 and radiosonde 

observations from about 1960 both reveal the 

cooling. The trend is about −0.5K per decade in 

the lower stratosphere and about −1 to −2K per 

decade in the middle stratosphere (about 25 to 30 

km in altitude). The latitude dependence of the 

temperature trends is not fully consistent across 

the various datasets, especially in the tropics, 

and remains a topic of research (WMO, 2007). 

The time series of temperatures reveals a nonmonotonic decrease in the lower stratospheric 

temperatures (Figure 4.7). Volcanic aerosols 

formed in the aftermath of a volcanic eruption 

that injects sulfur dioxide in to the stratosphere 

lead to a warming of the stratosphere for a 

few years following an eruption. Both the 

El Chichón (1982) and Mt. Pinatubo (1991) 

volcanic eruptions increased stratospheric 

aerosol amounts (Figure 4.7) (McCormick et 

al., 1995; Pawson et al., 1998). In the evolution 

of the global lower stratospheric temperature, 

a sharp increase, lasting for approximately two 

years, is found immediately following the El 

Chichón (1982) eruption and is followed by a 

period of quasi-steady temperatures that are 

lower than the pre-eruption value. After the 

eruption of Mt. Pinatubo (1991), temperatures 

again increased sharply and were followed 

by a steady period in which the temperatures 

became lower than before this eruption. There 

is a slightly reduced cooling towards the end 

of 1990s and beginning of 2000s (Mears et 

al., 2003). 

Climate model simulations also show that the 

combined influences of the agents that are 

known to “drive” the climate system offer 

a reasonable quantitative explanation of the 

observed non-monotonic decrease of the 

temperatures in the global lower stratosphere 

(Seidel and Lanzante, 2004; Dameris et al., 

2005; Ramaswamy et al., 2006). The global 

stratospheric temperature trends over the 

past two to three decades are attributed in 

modeling studies to a combination of increases 

in well-mixed greenhouse gases and water 

vapor, and decreases in ozone (Ramaswamy 

and Schwarzkopf, 2002; Schwarzkopf and 

Ramaswamy, 2002; Langematz et al., 2003; 

Shine et al., 2003; Santer et al., 2006). The 

above studies indicate that attribution of the 

cooling trend is possible on the global-annual 

and zonal-annual scales, and for the springtime 

Antarctic, but smaller spatial scales and 

Figure 4.6 The processes in a Chemistry Climate Model (CCM) are represented by four basic 

groups: transport, dynamics, radiation, and stratospheric chemistry and microphysics. Significant 

interactions occur between aspects within each group. The CCM requires inputs describing human 

activities and natural processes. The CCM provides as output projections of future ozone abundances 

and their distribution, along with a large variety of other parameters and diagnostics. See Box 4.3 

(Eyring et al., 2005).
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seasonal behavior pose problems in attribution 

owing to the dynamical variability. The results 

from an ensemble of climate models are shown 

in Figure 4.7 for stratospheric temperature 

anomalies calculated as global and monthly 

means (Santer et al., 2006). In addition to 

stratospheric ozone depletion, the models all 

include climate forcings from changes in wellmixed greenhouse gases (WMGHGs), sulfate 

aerosol, volcanic aerosol, and solar irradiance. 

The temperature anomalies are differences 

from a 1979 to 1999 reference period. The 

models, in general, account for the long-term 

decrease in stratospheric temperatures and the 

short-term increases caused by the two large 

volcanic eruptions. 

Hansen et al. (2005) examined the effects on 

climate of a wide range of forcings, showing 

that different forcings produce different 

response patterns in the vertical temperature 

profile. Results from climate model simulations 

as outlined in CCSP (2006) show:

• Increases in greenhouse gases warm the 

troposphere and cool the stratosphere

• Volcanic aerosols warm the stratosphere 

and cool the troposphere

• Increase in solar forcing warms most of 

the atmosphere

• Increases in tropospheric ozone warm the 

troposphere

• Decreases in stratospheric ozone cool the 

stratosphere

• Sulfate aerosols cool the troposphere and 

slightly warm the stratosphere.

The projections for the twenty-first century by 

coupled atmosphere-ocean general-circulation 

models (AOGCMs) using IPCC emissions 

scenarios show that average global temperatures 

will continue to decrease in the stratosphere 

and increase in the troposphere (Figure 4.8). 

This result is primarily a consequence of 

increases in WMGHGs (mainly CO2). Changes 

in the thermal gradients in the stratosphere 

and troposphere, initiated by greenhouse 

gas and aerosol changes, could additionally 

For the twenty-first 

century, average 

global temperatures 

are projected to 

continue to decrease 

in the stratosphere 

and increase in the 

troposphere, primarily 

as a consequence of 

increases in well-mixed 

greenhouse gases.

Figure 4.7 Temperature anomalies in the lower stratosphere as calculated by an ensemble of climate models. All anomalies 

are expressed as departures from a 1979 to 1999 reference period average. These models were chosen because they satisfy 

certain minimum requirements in terms of the forcings applied in the model run. All were driven by changes in well-mixed 

greenhouse gases (WMGHGs), sulfate aerosol direct effects, tropospheric and stratospheric ozone, volcanic aerosols, and 

solar irradiance. Model results are compared with observations derived from satellite datasets (RSS and UAH). Further 

details are discussed in Section 4.3.2 (From Santer et al., 2006: Figure 5.2 and WMO, 2007: Figure 5-3). 
The U.S. Climate Change Science Program

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Chapter 4

alter stratosphere-troposphere interactions 

and the state of the stratosphere. Changes of 

water vapor in the stratosphere arising from 

tropospheric warming, and possible changes in 

convective activity and transport of water to the 

stratosphere, also can affect the stratospheric 

thermal state. Volcanoes can be expected to 

substantially alter the climate and chemistry 

of the stratosphere for a few years through 

the particulates produced and impacts on 

atmospheric circulation if they erupt more 

frequently and/or if they inject much more 

material into the stratosphere than the Mt. 

Pinatubo eruption.

4.3.2.1 Response of ozone to

stratospheric temperature

changes

With the coupling of ozone with climate 

parameters as outlined above, the effect of 

temperature changes on ozone is difficult 

to isolate. However, model simulations 

reveal some strong tendencies arising due to 

temperature changes. In the upper stratosphere, 

ozone amounts are controlled primarily by 

photochemical processes rather than transport, 

and these processes are considered well 

understood. When temperatures decrease, 

ozone loss slows in the dominant photochemical 

cycles (reactive nitrogen [NOx], reactive 

chlorine [ClOx], and reactive hydrogen [HOx]) 

(Figure 4.9). For example, 15-20 percent 

ozone increases were calculated in the upper 

stratosphere for a climate with doubled 

(CO2) concentrations (Jonsson et al., 2004). 

In the lower stratosphere, the destruction 

rate decreases for lower temperatures, but 

production and destruction rates are lower than 

in the upper stratosphere and transport plays a 

more important role. As a result, temperature 

changes have less influence on steady state 

ozone values in the lower stratosphere than in 

the upper stratosphere.

In the polar lower stratosphere, the reduction in 

photochemical loss with lower temperatures can 

be completely offset by increased activation of 

reactive chlorine and bromine, which increases 

ozone loss. Lower temperatures promote the 

formation of polar stratospheric clouds (PSCs), 

which facilitate heterogeneous reactions 

that form reactive halogens from reservoir 

gases. In the Arctic region, increased reactive 

halogens have the largest effect in controlling 

the ozone response to lower temperatures. For 

Northern Hemisphere winters from 1993 to 

present, a strong linear relationship is found 

between winter/early spring ozone depletion 

and the volume of air containing PSCs during 

the winter. The ODS abundances are nearly 

constant during this time period (Rex et al., 

2004). Arctic ozone depletion might increase if 

further reductions occur in Arctic stratospheric 

temperatures because temperature decreases 

can lead to increases in the duration and 

frequency of PSCs (Douglass et al., 2006).

In the Antarctic lower stratosphere, winter 

temperatures are well below the thresholds for 

heterogeneous conversion of halogen reservoirs 

for much longer periods and for much larger 

fractions of the polar vortex than found in the 

Arctic. Antarctic ozone depletion is currently 

much more extensive and complete, and 

decreasing temperatures would have less of 

an effect (Tilmes et al., 2006). Under current 

conditions, seasonal Antarctic ozone depletion 

is more sensitive to reductions in ODS amounts 

than to small decreases in temperature (see 

When temperatures 

decrease in the 

upper stratosphere, 

ozone loss slows 

in the dominant 

photochemical cycles.

Figure 4.8 Temperature trends in the troposphere and stratosphere calculated 

as global and annual means for the twenty-first century using atmosphere-ocean 

GCMs (AOGCMs, with no ozone chemistry). Temperatures are expected to 

continue to increase in the troposphere and decrease in the stratosphere. The 

calculations were made for two IPCC emission scenarios: A2 (high) and B1 (low). 

The symbols indicate the average trend computed for all models, while the thin 

horizontal lines indicate the range. The vertical scale represents altitude from the 

surface (1000 hectopascal, hPa) to approximately 30 km (10 hPa). The pressure 

of the tropopause, which divides the stratosphere above from the troposphere 

below, varies between 100 and 300 hPa over most of the globe (WMO, 2007).
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

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Chapter 3) (Newman et al., 2004). As ODS 

abundances decrease in the coming decades, 

polar ozone destruction due to reactions with 

halogen species ultimately will decrease in 

both hemispheres regardless of changes in the 

frequency and duration of PSCs.

4.3.3 Stratospheric Water 

Vapor Changes

Atmospheric water vapor is the most important 

and abundant greenhouse gas. Change in 

the global distribution of water vapor is one 

of the important responses to the humancaused climate forcings summarized in Figure 

4.5. Water vapor enters the stratosphere 

primarily through the tropical tropopause. 

The water vapor abundance is reduced in 

dehydration processes involving low tropical 

tropopause temperatures and the formation 

and sedimentation of ice particles. Methane, 

released in the troposphere and oxidized in 

the stratosphere, is the underlying cause of 

the water-vapor component of human-caused 

radiative forcing as summarized in Figure 4.5. 

Stratospheric water vapor has been measured 

by a wide variety of instruments and platforms, 

including balloons, aircraft, and satellites. The 

longest time series of continuous measurements 

is from small-balloon observations beginning 

in 1980. These measurements show that water 

vapor has increased at all levels between 15 and 

26 km altitude. At 17 to 22 km, the increase can 

be expressed as a trend of about 5-10 percent 

per decade (Figure 4.10) (Oltmans et al., 2000; 

Rosenlof et al., 2001). Other stratospheric 

observations up to 30 to 35 km also show 

increasing trends, but over shorter time periods 

and with a high degree of variability (SPARC, 

2000; Rosenlof et al., 2001). Part of the longterm increase in water vapor is attributable to 

increases in methane abundances due to human 

emissions. Methane, which has increased by 

about 0.55 parts per million by volume (ppmv) 

since the 1950s, is oxidized in the stratosphere, 

producing two water molecules for each 

molecule of methane. The methane water vapor 

source in the stratosphere increases radiative 

forcing from water vapor by an estimated 0.1 

W per m2

 (Myhre et al., 2007). 

Water vapor has 

increased about 5 

to 10% per decade 

at a height of 17 to 

22 km since 1980.

Figure 4.9 Comparison of ozone production and loss rates as a function of stratospheric 

temperature for 40-km altitude (left panel) and 20-km altitude (right panel) conditions at 

45°N at equinox (end of March). The colored regions indicate the contribution from the 

principal loss cycles of ozone: odd-oxygen (Ox), reactive chlorine (ClOx), reactive bromine 

(BrOx), reactive nitrogen (NOx), and reactive hydrogen (HOx). The fractional contribution 

of each cycle varies with temperature differently in the two regions. The top trace in each 

panel is the ozone value at the end of 20-day runs of a chemical box model starting from 

climatological values for ozone, other constituents, and temperature (250K at 40 km; 215K 

at 20 km). At 40 km, the production rate coincides with the sum of loss rates because ozone 

is in photochemical balance at all temperatures shown. At 20 km, ozone production can be 

higher or lower than total loss depending on temperature because transport plays a more 

important role. The changes in ozone after 20-day runs are much smaller at 20 km than at 

40 km, confirming that small temperature changes in the upper stratosphere will significantly 

alter ozone abundances (adapted from IPCC/TEAP, 2005).
The U.S. Climate Change Science Program

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Chapter 4

4.3.3.1 Response of ozone to

stratospheric water vapor

changes

Increases in stratospheric water lead to increases 

in reactive hydrogen (HOx), which catalyze the 

chemical destruction of ozone (Wennberg et 

al., 1994; Brasseur and Solomon, 1986). Ozone 

destruction is chemically buffered with a 

combination of loss cycles so that the response 

to increased HOx is generally not linear and 

varies with temperature and location in the 

stratosphere (Figure 4.9). Model simulations 

show that a 1 percent increase per year, longterm trend in water vapor would increase 

ozone loss due to increases in HOx and delay 

the recovery of the ozone layer (Dvortsov and 

Solomon, 2001). Increased water vapor also 

increases the temperature threshold for PSC 

formation in both polar regions because PSC are 

formed, in part, from water-vapor condensation. 

A higher threshold increases heterogeneous 

conversion of chlorine and extends the time 

period over which PSCs can form in the winter 

season (Stenke and Grewe, 2005). Both effects 

lead to increased ozone destruction in polar 

regions at constant ODS amounts. However, 

the sensitivity of ozone to PSCs will decrease 

as ODS amounts decrease, because less chlorine 

and bromine will be available to participate in 

ozone destruction reactions.

4.3.4 Changes in Ozone From 

Increases in Long-Lived Gases 

in the Stratosphere

The atmospheric concentrations of the three 

long-lived greenhouse gases, CO2, CH4, and 

N2O, have increased significantly due to 

human activities since 1750 and are expected 

to continue increasing in the 21st century 

(IPCC, 2007). These continuing increases 

have consequences for ozone amounts and, 

hence, indirectly influence climate through 

the changes they produce in ozone (Portmann 

and Solomon, 2007). Calculations with a twodimensional, chemical-radiative-dynamical 

model illustrate the sensitivity of ozone to each 

of these gases (Figure 4.11). Carbon dioxide 

increases, as previously discussed, reduce 

stratospheric temperatures and ozone loss rates, 

and consequently, increase ozone amounts in the 

mid- to upper stratosphere. The increased ozone 

in the upper stratosphere can lead to reduced 

ozone in the lower stratosphere because of the 

Since about 2000, the water vapor in key 

balloon and satellite observations in the midto lower stratosphere have shown significant 

decreases (Randel et al., 2004). As a possible 

explanation, an analysis of the tropical 

tropopause temperatures for 1992 to 2005 

shows that satellite water vapor amounts are 

consistent with interannual changes in the cold 

point temperatures and with the occasions of 

anomalously low tropopause temperatures 

(Randel et al., 2004; 2006). Tropopause 

temperatures modulate the dehydration of 

air entering the lower stratosphere from the 

troposphere. In contrast, the earlier, longer 

water record from balloon measurements is not 

fully consistent with the record of tropopause 

temperatures (Seidel et al., 2001). In general, 

the attribution of the causes of observed water 

vapor changes and trends in the stratosphere 

is incomplete, suggesting that projections of 

future amounts are uncertain.

Increases in reactive 

hydrogen (HOX) and 

the temperature 

threshold for polar 

stratospheric cloud 

(PSC) formation lead 

to increased ozone 

destruction in polar 

regions at constant 

ozone-depleting 

substance (ODS) 

amounts. However, the 

sensitivity of ozone 

to PSCs will decrease 

as ODS amounts 

decrease because less 

chlorine and bromine 

will be available to 

participate in ozone 

destruction reactions.

Figure 4.10 Time series of stratospheric water vapor mixing ratios (ppm) 

for the period 1980 to 2005 showing increases up to about 2001 followed by 

a decreasing trend. The measurements were made with a balloon-borne frost 

point hygrometer over Boulder, Colorado (40°N, 105°W). The data points 

are averages over 17 to 22 km altitudes. The thin line is a smoothed fit to the 

measurements. HALOE satellite observations for 1992-2005 are shown with 

the heavy line for the same altitude near Boulder (latitude 35°N to 45°N, longitude 80°W to 130°W). Preliminary revisions to the frost-point data reveal a 

slightly smaller trend (Scherer et al., 2007). Updated from Randel et al. (2004) 

(WMO, 2007).
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

reduced penetration of solar UV into the lower 

stratosphere. Increases in N2O lead to increases 

in the NOx catalytic loss cycle for ozone in 

the mid- to upper stratosphere, because N2O 

decomposes to form NOx in the stratosphere. 

The effect of increased NOx is less in the lower 

stratosphere, because the NOx loss cycle plays 

a less prominent role, competing with the HOx

and ClOx catalytic loss cycles (Figure 4.7) 

(Wennberg et al., 1994). Finally, the oxidation 

of CH4 increases H2O and ozone losses in the 

HOx catalytic cycle in the upper stratosphere 

and lower mesosphere. In the troposphere, 

ozone is increased because oxidation of CH4

catalyzed by NOx produces ozone. 

4.3.4.1 Changes in ozone from

stratospheric circulation

changes

The net mass exchange between the troposphere 

and stratosphere is associated with the largescale Brewer-Dobson circulation (Holton 

et al., 1995), with a net upward flux in the 

tropics balanced by a net downward flux in 

the extratropics. Model studies indicate that 

climate change will impact the mass exchange 

rates across the tropopause. For a doubled 

CO2 concentration, all 14 climate-change 

model simulations analyzed by Butchart et al.

(2006) showed an increase in the annual mean 

troposphere-to-stratosphere exchange rate, with 

a mean trend of about 2 percent per decade. 

Consequences of such an increase include 

shorter lifetimes and more rapid removal from 

the atmosphere for long-lived gases, including 

CFCs, CH4 and N2O (Butchart and Scaife, 

2001), and increased mass flux of ozone from 

the stratosphere to the troposphere at mid- and 

high latitudes. A model simulation for 2100 with 

projected climate change (Zeng and Pyle, 2003) 

shows that a strengthened Brewer-Dobson 

circulation would increase ozone amounts in 

the lower stratosphere and the flux of ozone 

to the troposphere. A larger flux results from 

increased transport downward across the 

tropopause and enhanced ozone amounts in the 

extratropical lower stratosphere. The enhanced 

ozone results from the strengthened circulation 

and decreases in ODSs and temperatures.

4.4 The Effect of Ozone

Changes on Climate

Parameters

Ozone and climate change are highly coupled, 

as illustrated in Figures 4.1 and 4.2. The 

response of ozone to changes in stratospheric 

temperature and water vapor is discussed 

in Section 4.3. In this section, changes in 

atmospheric temperatures and circulation are 

described as examples of the responses in the 

climate system to ozone depletion. 

Model studies indicate 

that climate change 

will impact the mass 

exchange rates across 

the tropopause.

Figure 4.11 Comparison of perturbations to ozone amounts for increases in three principal greenhouse gases, N2O, CH4, and 

CO2, as a function of altitude in the troposphere, stratosphere, and mesosphere. The changes are computed for six 30°-wide 

latitude bands with the NOCAR 2-D model and expressed as the percent change from 2000 to 2100. The ozone changes vary 

widely in magnitude and show increases as well as decreases. The changes result from a variety of chemical and radiative processes. The global mean change is shown with the black line in each panel. The halogen changes follow the WMO (2003) scenario 

and greenhouse gas increases follow the IPCC A2 scenario. The latter is high compared to other scenarios but was chosen to 

maximize the ozone response in the model (adapted from Portmann and Solomon, 2007).
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Chapter 4

4.4.1 Response of Stratospheric and 

Tropospheric Temperatures to Ozone 

Depletion

Temperatures have decreased throughout the 

stratosphere in recent decades as described 

above in Section 4.3.2. Furthermore, model 

simulations show that a combination of 

increases in greenhouse gases and water 

vapor and decreases in ozone can account for 

observed temperature changes. A more detailed 

examination of the influence of ozone on 

temperature was carried out with the SKYHI 

GCM for ozone decreases observed in the 

period from 1979 to 1997 (Ramaswamy and 

Schwarzkopf, 2002). The results in Figure 4.12 

indicate that in the lower to middle stratosphere 

(100 to 5 hPa; 21 to 38 km), ozone changes 

create a larger decrease in temperature than 

increases in WMGHGs. In this case, these 

include CO2, CH4, N2O, CFC-11, CFC-12, 

CFC-113, and HCFC-22. However, above 

about 5 hPa (about 38 km) changes in both 

ozone and WMGHGs contribute significantly 

to temperature decreases. Thus, depletion in 

stratospheric ozone plays a significant role 

throughout the stratosphere in creating a 

reduction in stratospheric temperatures in GCM 

simulations. 

The response of temperature to stratospheric 

ozone depletion extends into the upper 

troposphere. The Reading Narrow Band Model 

was used to calculate temperature changes for 

observed ozone depletion with the assumption 

of fixed dynamical heating (Forster et al., 

2007). Model cooling occurs in the 70 to 30 

hPa (13 to 25 km) region due to ozone depletion. 

Shortwave absorption and upwelling longwave 

radiation are both reduced and contribute 

comparably to the cooling in this region. The 

missing ozone also causes a decrease in the 

downwelling longwave radiation that causes 

reduced temperatures at altitudes below the 

ozone depletion region (150 to 100 hPa; 14 to 21 

km). This response or coupling of temperatures 

in different altitude regions is found at all 

latitudes in the model and may be a cause of 

upper tropospheric temperature trends.

4.4.2 Response of Surface 

Temperatures to Antarctic 

Ozone Depletion

The largest depletion in stratospheric ozone is 

found over the Antarctic in late winter/early 

spring. Studies of Antarctic ozone depletion 

have revealed strong evidence for responses 

in the temperatures and circulation of the 

Antarctic troposphere (Gillet and Thompson, 

2003; Thompson and Solomon, 2002). Severe 

ozone depletion strengthens the circumpolar 

winds of the Antarctic winter vortex in many 

model simulations. Recent observations show 

that strengthened circumpolar winds extend 

to the surface, especially in the summer 

months, with changes in geopotential heights 

(a vertical coordinate system referenced to 

Earth’s mean sea level) serving as a proxy. A 

model with high vertical resolution was used 

to show that anomalies in geopotential height 

Figure 4.12 The response of stratospheric temperatures calculated with 

a global climate model (GCM) using observed ozone depletion and changes 

in well-mixed greenhouse gases (WMGHGs) between 1979 and 1997. The 

upper panel shows modeled temperature response to ozone changes alone 

and the lower panel shows the response to ozone changes plus increases 

in greenhouse gases over the same period. Ozone changes alone cause a 

significant temperature response in the period. Solid lines enclose regions 

of statistical significance. The vertical scale shows pressure (left) and altitude (right) (adapted from Ramaswamy and Schwarzkopf, 2002).

A combination 

of increases in 

greenhouse gases 

and water vapor 

and decreases in 

ozone can account 

for observed 

temperature changes 

in the stratosphere 

in recent decades.
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

in the troposphere could be well simulated in 

intensity and seasonality (Gillet and Thompson, 

2003). Changes in surface circulation also 

lead to cooling over most of the Antarctic 

continent and modest warming of the Antarctic 

Peninsula. Figure 4.13 shows model results 

compared to observed changes in 500-hPa 

geopotential height over a 22-year period and 

in surface temperature over a 32-year period 

(1969 to 2000), both averaged over December 

to May. The observed and simulated patterns 

in geopotential height and surface temperature 

show strong similarities, reinforcing the 

conclusion that stratospheric circulation 

influences tropospheric circulation patterns 

and, hence, that intense stratospheric ozone 

depletion can bring about changes in surface 

climate parameters. 

4.5 Importance of

Volcanoes

4.5.1 The Effect of Volcanic 

Aerosol on Ozone

Large volcanic eruptions are those that inject 

significant quantities of sulfur dioxide (SO2) 

into the stratosphere. The SO2 is subsequently 

oxidized to sulfuric acid, which condenses 

onto preexisting aerosols, causing significant 

increases in aerosol surface area and volume 

in the lower stratosphere. As a consequence, 

heterogeneous reactions occurring on these 

surfaces gain prominence in the chemical 

production and loss balance of ozone, leading 

to decreased ozone amounts (WMO, 2007: 

Figures 3-26). These reactions convert NOx to 

a more stable form, nitric acid (HNO3). In the 

lower stratosphere, reduced NOx increases the 

role that reactive halogen compounds (ClOx) 

play in destroying ozone. Analysis following 

the most recent large volcanic eruption, Mt. 

The intense ozone 

depletion above 

Antarctica brought 

about changes in surface 

climate parameters 

of that continent.

Figure 4.13 Comparison of model results (left column) with observations (right column) 

for changes in 500-hPa geopotential height (m) (upper row) and near-surface temperature 

(K) (lower row) in the Southern Hemisphere in response to stratospheric ozone depletion 

between 1979 and 1997. The observed and simulated patterns in geopotential height and 

surface temperature show strong similarities, reinforcing the conclusion that stratospheric 

circulation influences tropospheric circulation patterns and, hence, that intense stratospheric 

ozone depletion can effect changes in surface climate parameters (adapted from Thompson 

and Solomon, 2002).
The U.S. Climate Change Science Program

130

Chapter 4

 A period of frequent 

volcanic eruptions 

in the next century 

could enhance ozone 

depletion from 

ozone-depleting 

substances (ODSs) 

but the potential for 

enhancement will 

lessen as global ODS 

abundances decline.

entering the stratosphere from the troposphere 

and influence ozone amounts through the 

sensitivity of ozone chemical reaction rates. 

Climate-chemistry model simulations of the 

temperature perturbations after the eruptions 

of El Chichón and Mt. Pinatubo often show 

larger increases than observed (Figure 4.7). 

The elevated temperatures are evident for 

several years and are followed by an overall 

slow cooling. The strength of the volcanic 

temperature response varies substantially 

between the different CCM and climate models 

(see Eyring et al., 2006).

Volcanic eruptions, while not predictable, are 

expected to occur in the future atmosphere. A 

large volcanic eruption is likely to occur in the 

next 30 years, based on the historical record 

(Roscoe, 2001). Infrequent large volcanic 

eruptions would affect ozone with timescales 

as observed for previous large eruptions (Figure 

4.7). A period of frequent large eruptions in the 

next century could enhance ozone depletion 

from ODSs for many years but the potential 

for enhancement will lessen as global ODS 

abundances decline in the coming decades 

(Figure 2.12). Whenever the stratosphere is 

cleansed of volcanic aerosol, ozone abundances 

are expected to recover fully from volcanic 

effects. 

4.6 Summary

Stratospheric ozone and climate change are 

linked through a variety of processes. Radiative 

forcing of climate occurs from depletion of 

stratospheric ozone, as well as increases in 

ozone-depleting substances. Global ozone 

depletion is a principal cause of decreasing 

temperature trends in the stratosphere and 

upper troposphere. Severe ozone depletion 

over Antarctica has changed the circulation 

over the continent in both the stratosphere and 

troposphere and altered surface temperatures. 

Other important components of human-caused 

climate change arise from emissions of longlived greenhouse gases, such as carbon dioxide. 

Observed and anticipated changes in climate 

parameters include decreases in stratospheric 

temperatures and increases in stratospheric 

water vapor, carbon dioxide, methane, and 

nitrous oxide. Lower stratospheric temperatures 

reduce ozone loss rates in the mid- to upper 

Pinatubo in 1991, shows that ozone amounts 

reached record lows and that halogen reactions, 

aided by temperature variability, could explain 

the observed losses (Solomon et al., 1998; Tie 

and Brasseur, 1995).

As ODS amounts decrease to pre-1980 levels 

in the coming decades, the sensitivity of 

ozone to depletion caused by volcanic aerosol 

reactions will also decrease. Global ozone levels 

decreased by about 2 percent following the 

1991 Mt. Pinatubo eruption. Sedimentation and 

transport removal of volcanic aerosol occurs 

over a two to three year period following an 

eruption, so the effects are short lived compared 

to ODS atmospheric lifetimes, which are 45 to 

100 years for principal species (e.g., CFC-11 and 

CFC-12). Thus, expectations for the long-term 

recovery of ozone are not significantly affected 

by episodic volcanic eruptions.

The plumes of large volcanic eruptions contain 

significant amounts of hydrochloric acid (HCl) 

that are removed in the troposphere by uptake 

and sedimentation of the liquid aerosols formed 

(Tabazadeh and Turco, 1993). However, some 

eruptions inject non-negligible amounts of 

HCl into the stratosphere, adding to inorganic 

chlorine. For example, temporary increases 

in HCl column amounts of up to 40 percent 

were observed after the Mt. Pinatubo eruption 

(Coffey, 1996). Overall, the frequency of 

explosive volcanic eruptions has been low 

in the past two decades, thereby precluding 

significant, long term, volcanic enhancements 

in global stratospheric chlorine.

Volcanic aerosols have a direct radiative impact 

by scattering incoming solar radiation and 

absorbing solar infrared radiation. Absorption 

increases stratospheric temperatures while 

decreasing surface temperat ures. The 

tropospheric cooling that results can be expected 

to change the tropospheric circulation, as well as 

the interaction between the stratosphere and the 

troposphere. Lower stratospheric temperatures, 

for example, following the eruptions of El 

Chichón and Mt. Pinatubo, were observed to 

increase by about 1K near 20 km altitude (Figure 

4.7). The loss of ozone following an eruption also 

adds to the temperature perturbation. Lower 

stratospheric temperatures influence water 

vapor amounts through dehydration of air parcels 
131

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

stratosphere, thereby aiding the recovery from 

ozone depletion. Enhanced water vapor alters 

ozone destruction rates in reactive hydrogen 

photochemistry and can increase the frequency 

and extent of polar stratospheric clouds, which 

aid ozone destruction. These varied composition 

changes contribute to circulation changes in the 

stratosphere and between the stratosphere and 

troposphere that can cause significant changes 

in the ozone distribution. The increases in 

stratospheric aerosols that follow explosive 

volcanic eruptions create several-year changes 

in climate parameters in the stratosphere and 

troposphere and increase ozone depletion.

The complexity of the interactions between 

ozone and climate involving changes in 

atmospheric composition pose a challenge to 

our understanding of basic stratospheric and 

tropospheric processes. Tools of the complexity 

of chemistry climate models (CCMs) are required 

to combine stratospheric transport, dynamics, 

radiation, and chemistry and microphysics 

to analyze past ozone amounts and project 

future amounts. CCMs guided by atmospheric 

observations help define the sensitivity of 

ozone to future climate changes and reduce the 

uncertainties in our understanding of ozone and 

climate interactions. As ozone depletion slows 

and ozone amounts recover from ODSs in the 

coming decades, expected changes in climate 

parameters will increase in importance in 

influencing stratospheric ozone amounts.

4.6.1 Relevance for the United States

Human activities have led to changes in 

ozone abundances and climate parameters. 

Ozone depletion is attributed primarily to the 

accumulation of ozone-depleting substances 

and climate change is attributed to increases 

in long-lived greenhouse gases, changes in 

aerosols and clouds, and surface albedo changes 

(Figure 4.5). Ozone is further influenced 

by changes in climate parameters such as 

stratospheric temperatures and composition, 

and atmospheric circulation. Since activities 

in the United States have caused significant 

emissions of greenhouse gases and ozonedepleting substances, the changes in ozone and 

climate attributable to human activities are, in 

part, attributable to the United States.

Decisions initiated or supported by U.S. 

policymakers have great potential to influence 

ozone and climate in the future. Important 

decisions could be taken on the following topics 

or issues:

• Increased stringency of Montreal Protocol 

regulations. The Montreal Protocol 

regulates production and consumption of 

ODSs in developed and developing nations. 

Stratospheric ODS amounts will decline 

to pre-1980 values around the middle of 

this century based on current regulations. 

More stringent regulations could accelerate 

this decline. For example, recent unratified 

regulation, supported by the United States, 

accelerates hydrochlorof luorocarbon 

(HCFC) production in developed and 

developing nations. 

• Increased destruction or capture of ODS 

banks. Banks of ODS compounds represent 

large sources of future ODS emissions (see 

Chapters 2 and 5).

• Increased climate protection under the 

Montreal Protocol. ODS compounds are 

also greenhouse gases. Reducing ODS 

production and consumption under the 

Montreal Protocol has led to significant 

reductions in ODS atmospheric abundances 

and their associated radiative forcing 

of climate. Further reductions in ODS 

production, as well as emissions, will 

further protect climate. The accelerated 

HCFC phase-out under the Montreal 

Protocol represents a large potential 

benefit to climate. In addition, promotion 

of low-GWP compounds as replacements 

for ODSs in widespread applications can 

help minimize the climate consequences 

of new and existing Montreal Protocol 

regulations.

• Reductions in the future growth rates of 

carbon dioxide, methane, and nitrous oxide 

emissions. Ozone and climate are strongly 

influenced by carbon dioxide, methane and 

nitrous oxide emissions. Emissions of these 

gases ultimately affect the photochemical 

production and loss of ozone in the 

troposphere and/or stratosphere. All are 

greenhouse gases that have increased 

significantly due to human activities.
The U.S. Climate Change Science Program

132

Chapter 4
133

 

The Future and Recovery

CHAPTER

5

Convening Lead Author: Malcolm Ko, NASA 

Lead Authors: John S. Daniel, NOAA; Jay R. Herman,

 NASA; Paul A. Newman, NASA; V. Ramaswamy, NOAA

 

KEY ISSUES

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

This chapter presents results on how future halogen loading is expected to affect the behavior 

of total column ozone and the prospect for the detection/validation of the anticipated recovery 

trend. In a hypothetical argument, if circulation, climate, and the background atmosphere were to 

remain unchanged as of the present-day, the projection of ozone could be based essentially upon 

future halogen loading. However, the reality is that the concentrations of other trace gases that 

affect ozone (e.g., methane [CH4], nitrous oxide [N2O], and water vapor) are changing because 

of changes in emissions; climate, which also affects ozone abundances, is changing as well. The 

model-simulated results show that the ozone increase expected between now and 2025 is largely 

due to the anticipated decrease in halogen loading, and ozone-depleting substances (ODSs) will 

remain one of the drivers of human-caused ozone depletion up until the middle of the twentyfirst century, when the halogen loading is expected to approach its 1980 value. Reductions in 

emissions of these chemicals represent the only known currently acceptable method to reduce 

this depletion in this period. The effects of climate change (largely driven by increases in carbon 

dioxide, CO2) and changes in other trace gases (e.g., CH4, N2O, and hydrofluorocarbons) will 

play an increasing role in the ozone behavior. 

Ozone is only one of many factors that affect UV (ultraviolet radiation) at the surface. Future 

changes in UV will be discussed in this chapter in the context of the projected ozone change in 

the stratosphere. The Equivalent Effective Stratospheric Chlorine (EESC) is used to compare 

the relative impacts of various ODS emissions scenarios on future ozone. Included in this discussion is the radiative forcing associated with the halocarbons as well as the hydrofluorocarbons 

(HFCs) used as replacements for the ODSs. The contribution from the United States to the 

future halocarbon loading will be addressed in the context of EESC and radiative forcing. 

The key issues, in the form of questions, that are addressed in this chapter include:

• What is the future behavior of ozone as predicted by numerical models?

• What is the future behavior of ultraviolet radiation at the Earth’s surface?

• Are there any new findings concerning projected future emissions of ODSs?

• What is the radiative forcing associated with ODSs and HFCs emitted as replacement 

chemicals for the ODSs and how will it likely change in the future?

• To the extent that the emissions from a specific country can be used to estimate its contribution to global ozone depletion and radiative forcing by ODSs and their replacements, 

what is the United States’ contribution? 
The U.S. Climate Change Science Program Chapter 5

134

Key Findings

Equivalent Effective Stratospheric Chlorine (EESC) is a useful index for comparing relative merits of different halocarbon emission scenarios in minimizing ozone depletion. 

Current scenarios being considered include projected emissions of long-lived source gases 

only and do not include emissions of very short-lived source gases. 

• The time for EESC to return to the 1980 level (the EESC recovery date) and the integrated 

EESC values (to the EESC recovery date) provide useful metrics to compare the relative 

merits of various emissions scenarios. 

• There have been suggestions that the mean age of air, and age-dependent release factors 

should be used in the calculation of EESC. This is potentially useful for calculating EESC 

values that are more representative of polar ozone depletion, in particular. The new recipe 

will change the absolute values of the metric for global ozone depletion, but should not 

qualitatively affect the relative benefit estimates of the different scenarios. 

Two-dimensional chemistry transport models (both with and without climate feedback) 

and three-dimensional climate chemistry models (3-D CCMs) were used to simulate 

the behavior of ozone in the twenty-first century using the WMO (2003) baseline scenario.

The scenario includes changes arising from natural and anthropogenic activities other than ODS 

emissions. Thus, it is expected that the model-simulated ozone recovery date will differ from 

the EESC recovery date. The model-simulated ozone results will be discussed in the context of 

the EESC recovery date between 2040 and 2050 for midlatitudes, and between 2060 and 2070 

for the polar regions. Analyses of simulation results indicate that:

For the model-simulated ozone content between 60°N and 60°S :

• Between now and 2020, the simulated total ozone content will increase in response to the 

decrease in halogen loading. 

• Some 3-D CCMs predict that stratospheric cooling and changes in circulations associated 

with greenhouse gas emissions will enable global ozone to return to its 1980 value up to 

15 years earlier than the EESC recovery date. 

• Based on the assumed scenario for the greenhouse gases (which includes CH4 and N2O), 

the ozone content between 60°N and 60°S is expected to be 2% above the 1980 values by 

2100. Values at midlatitudes could be as much as 5% higher.

For the model-simulated Antarctic ozone:

• The model-simulated ozone recovery date (the year when ozone returns to its 1980 value) 

for Antarctic ozone behavior depends on the diagnostics chosen. The minimum ozone value 

is projected not to start increasing until after 2010 in several models, while a decrease in 

ozone mass deficit in most models has occurred by 2005.

• Model simulations show that the ozone amount in the Antarctic will reach the 1980 values 

10 to 20 years earlier than the 2060 to 2070 time frame.

For the model-simulated Arctic ozone:

• Ozone in the Arctic region is expected to increase. Because of large interannual variability, 

the simulated results do not show a smooth monotonic recovery. The dates of the minimum 

column ozone from different models occur between 1997 and 2015. 

• Most CCMs show ozone values at 2050 larger than the 1980 values, with the recovery date 

between 2020 and 2040.

• Results from the majority of the models indicate that Arctic ozone depletion will not be 

significantly worse than what has already occurred. 

With the current scenarios, anthropogenic halogens identified in the Montreal Protocol 

should have a minimal effect on midlatitude ozone beyond 2050. In order to predict 

the future trend of ozone in that time frame, one must consider projections for climate 

changes and changes in trace gases such as other halogens, CH4, and N2O.
135

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

Applying ozone trend analysis techniques to future observations should enable one to 

confirm the time when halogen loading has minimal influence on global ozone within 

five to ten years after it occurs.

The future UV trend at the surface is likely to be more dominated by changes in clouds, 

aerosols, and tropospheric air quality than by ozone changes in the stratosphere. Equivalent Effective Stratospheric Chlorine (EESC) will still be a useful predictor for the relative 

effects of ODSs on future UV in terms of evaluating the different scenarios. 

Future halocarbon emissions are derived using a new bottom-up approach for estimating 

bank sizes. The new method gives future chlorofluorocarbon (CFC) emissions that are 

higher than previously estimated in WMO (2003).

• Current projected concentrations for EESC in the twenty-first century are higher than 

reported in WMO (2003) because the most recent CFC bank estimates, which are believed 

to be more accurate, are larger and lead to larger emissions, and the estimated emissions 

due to future production of hydrochlorofluorocarbons (HCFCs) from Article 5(1) countries 

are also larger.

• The EESC in the baseline scenario returns to the 1980 value in the year 2049, about five 

years later than the date based on the WMO (2003) baseline scenario.

• Compared to the WMO (2007) baseline scenario, cessation of all future emissions will 

bring the EESC recovery date earlier by 15 years to 2034. Integrated EESC (from 2007 to 

the EESC recovery date) from ODSs already in the atmosphere as of 2007 is 58% of the 

integrated EESC for the baseline scenario.

• If no future ODS production is assumed, the date when EESC returns to the 1980 level is 

moved earlier by six years to 2043. The integrated EESC from ODSs produced after 2007 

is 17% of the integrated EESC for the baseline case. 

Direct radiative forcing from ODSs and HFC replacement chemicals is approximately 

0.34 W per m2

 in 2005 and is expected to stay below 0.4 W per m2 through 2100 (according to the WMO (2007) scenario for ODSs and the IPCC Special Report on Emission 

Scenarios (SRES) scenarios considered for HFCs). This is to be compared with forcing 

from CO2 of 1.66 W per m2 in the 2005 atmosphere, increasing to as high as 5 W per m2

by 2100 for the SRES A1B scenario.

• Direct forcing from CFCs will decrease to 0.1 W per m2 by 2100. Direct forcing from HCFCs 

and other anthropogenic ODSs is expected to be negligible by 2100.

• Forcing from HFCs is 0.15 W per m2 and 0.24 W per m2 in 2050 and 2100, respectively, for 

the SRES A1B scenario, whereas other scenarios indicate that it will be lower. However, 

current observations suggest that the present atmospheric radiative forcing of the HFCs 

has been larger than computed for the SRES scenarios, primarily due to higher HFC-23 

concentrations.

The forcing associated with the observed ozone depletion was estimated to be about 

–0.05±0.10 W per m2 in 2005, corresponding to one-sixth of the direct forcing due to 

ODSs.

If one assumes that all of the observed ozone depletion is due to ODSs, that would imply that 

the indirect effect of ODSs due to ozone depletion is one-sixth of the direct effects for the mix 

of ODSs present in the atmosphere at that time. Current estimates assume that the indirect 

forcing from ODSs will decrease to zero approximately when EESC returns to its 1980 levels, 

while the direct forcing (mainly from CFCs and HFCs remaining in the atmosphere) will continue. 

The indirect forcing remains highly uncertain and is discussed in more detail in Chapter 2.

Using available historical and projected United States and global emissions estimates, 

we find that emissions from the United States contribute between about 15% and 37% to 

global EESC due to man-made emissions at 2030. For the same year, the United States’ 

contribution to direct radiative forcing from ODSs, HFCs, and PFCs (perfluorocarbons) 

is 19% to 41%.
The U.S. Climate Change Science Program Chapter 5

136

5.1 INTRODUCTION 

This chapter presents results on how future 

halogen loading will affect the future behavior 

of total column ozone and the prospect for the 

detection/validation of the expected recovery 

trend. In a hypothetical argument, if the 

transport circulation, the climate, and the background atmosphere were to remain unchanged 

as of the present-day, the projection of ozone 

could be based essentially upon future halogen 

loading. Chapter 2 discussed the concept of 

Equivalent Effective Stratospheric Chlorine 

(EESC) and how the values for midlatitude 

EESC and polar EESC could be used as metrics 

to approximate the effects of ODSs on ozone 

behavior. Daniel et al. (1995) pointed out that 

global stratospheric ozone losses became apparent from observations around the late 1970s to 

around 1980 (see also Figure 3.2 in this report). 

They also considered that the total ozone over 

Halley Bay dropped nearly linearly from 1980 

to 1990. From these facts, they assumed that 

the ozone trend due to halogen loading was 

close to zero for EESC smaller than the 1980 

value. If one assumes that the same holds for 

the future and nothing else is different, ozone 

is expected to recover to the 1980 value at the 

same time as EESC. The time for EESC to 

return to its 1980 value thus has been adopted 

as a metric to compare various control options. 

For this reason, the date when the future EESC 

value returns to its 1980 value is given some 

significance, and is referred to as the EESC 

recovery date (EESC RD). 

Since policy decisions are being made based on 

EESC, it would be prudent to perform analyses 

to see how well the EESC-based predictions 

correlate with model simulations of ozone 

recovery. Unfortunately, the scenarios used in 

the model simulations reported in WMO (2007) 

include changes in trace gases other than the 

halogens that affect ozone directly through 

chemical interactions, or indirectly through climate change. Thus, the model-simulated ozone 

recovery date (MS ORD, i.e., the date ozone 

returns to its 1980 value) for these simulations is 

expected to be different from the EESC RD. 

The results of numerical simulations of the 

future behavior of ozone as reported in the 

WMO (2007, Chapter 6) report are presented 

in Section 5.2. A discussion of how future observed ozone behavior can be used to detect and 

confirm the ozone recovery date is presented 

in Section 5.2.3. Section 5.3 discusses how 

future ozone may affect UV. Sections 5.4 and 

5.5 focus on expected future trends of the halocarbons through 2050. The future emissions 

and abundances of the CFCs and HCFCs are 

discussed in Section 5.4. The EESC will be used 

to compare the relative impacts of various ODS 

emissions scenarios on future ozone in Section 

5.5. Included in this section is a discussion of 

the radiative forcing associated with the ODSs 

as well as the HFCs used as replacements for 

them. The contribution from the United States 

to the global future halocarbon loading will be 

addressed in the context of EESC and radiative 

forcing in Section 5.6. 

5.2 MODEL SIMULATIONS 

AND ANALYSES OF OZONE 

BEHAVIOR 

Analyses of the over 40-year time series of 

global ozone data between 1964 and 2006 (see 

discussion in Chapter 3, Figure 3.2) indicate 

that it is possible to attribute the observed ozone 

behavior to several processes that affect ozone. 

These include the responses to the seasonal 

cycle, to the quasi-biennial oscillation (QBO), 

to the 11-year solar cycle, to episodic volcanic 

eruptions, and to halogen loading from halocarbons. In particular, most of the decreasing trend 

in ozone during this period can be correlated 

with EESC and attributed to the increase in 

halogen loading. It is anticipated that the decrease in halogen loading in the next 20 years 

will have a large influence on the decadal trend 

of ozone. To predict the future trend of ozone, 

one must identify all processes that may affect 

ozone, determine how the driving mechanisms 

may change (i.e., the scenarios), and employ 

numerical models to simulate the ozone behavior. The projected behavior will depend on the 

adopted scenario. The results presented in this 

section also show that different models predict 

different results for the same scenario. This 

indicates that there is still disagreement on how 

processes are represented in the models, and 

one must depend on further comparison with 

observations to resolve these issues. Finally, 

the purpose for presenting the model results in 

this chapter is to illustrate, in general terms, 

The decreasing trend 

in ozone during 

the period 1964 to 

2006 can be mostly 

attributed to the 

increase in halogen 

loading. It is anticipated 

that the decrease in 

halogen loading in the 

next 20 years will have 

a large influence on the 

decadal trend of ozone.
137

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

how the expected ozone behavior differs from 

the parameterized behavior based on EESC. It 

is beyond the scope of this report to address 

the various outstanding issues associated with 

simulating ozone behavior. Such attempts 

would greatly benefit from studies of changes 

in local ozone as functions of altitude.

5.2.1 Processes and Scenarios 

Used in Model Simulations 

It is clear that the model simulations must include 

the effects from changes in halogen loadings. 

The model simulations use prescribed surface 

concentrations of the halocarbons derived from 

projected emissions. The method for deriving 

the surface concentrations from emissions 

will be discussed in more detail in Section 5.4. 

The current best-estimate scenario for future 

halocarbon emissions (A1) is discussed in 

Table 8-4 and the surface concentrations are 

summarized in Table 8-5 of the WMO (2007, 

Chapter 8) report. Note that this scenario, as 

with all other scenarios previously considered 

in the WMO reports, considers only relatively 

long-lived (lifetime > 0.5 years) chlorine and 

bromine source gases. However, it has become 

clearer that very short-lived (VSL) ODSs also 

contribute to stratospheric ozone depletion, 

particularly short-lived bromocarbons. A 

more detailed discussion of the contribution of 

VSL compounds to stratospheric chorine and 

bromine loading can be found in Chapter 2 of 

WMO (2007). Note that the standard procedure 

for estimating EESC from emissions of longlived source gases (Box 2.7 in Chapter 2 of this 

report) should not be applied to VSL source 

gases in its current form. It was estimated 

in WMO (2007) that VSL compounds might 

contribute 50 ppt (parts per trillion by mole) 

of stratospheric chlorine and 3 to 8 ppt of 

stratospheric bromine. It was unclear whether 

any trend in these VSL compounds should 

be expected in the future or has occurred in 

the recent past. Enhancement in convective 

activities associated with future climate 

changes may increase the Ozone Depletion 

Potentials (ODPs) of the VSL species. 

Because the chapters in the WMO (2007) 

report were prepared in parallel, there was not 

sufficient time to use this most updated scenario 

in the model simulations. The model results 

presented in Chapter 6 of the WMO (2007) 

report were simulated using the scenario (Ab) as 

summarized in Table 4B-2 in the WMO (2003, 

Chapter 4) report. Using assumed values for 

the atmospheric lifetimes, the release factors 

of the halocarbons, and the transport lag from 

the tropopause, one can compute the date when 

midlatitude and polar EESC will reach its 1980 

value. For scenario Ab, the midlatitude EESC 

RD is calculated to be 2044. However, because 

of the uncertainties associated with the lifetimes 

and the release factors, Chapter 6 of the WMO 

report chose to discuss the results relative to 

an EESC RD range between 2040 and 2050. 

The model-simulated ozone recovery date (MS 

ORD) could be earlier than the EESC RD if 

the net effect from other non-ODS factors (see 

below) causes an increase in ozone relative to 

the 1980 value. Finally, the MS ORD for ozone 

at specific latitudes is likely to be different for 

different latitudes.

Note that even in simulations that keep all 

other parameters the same except for halogen 

loading, the EESC RD and MS ORD is still 

likely to be different because the lifetimes, 

the release factors, and the bromine efficiency 

factor in the model are likely to be different 

from what are assumed in the EESC calculation. 

To resolve this issue, comparison and validation 

of model-simulated atmospheric lifetimes and 

release factors should be a priority. Until this 

is accomplished, direct comparison between 

the EESC RD with the MS ORD is not as 

meaningful as it could be. 

Variations in natural factors such as changes 

in the Sun’s energy output and volcanic events 

will continue to have impacts on the ozone 

abundances. Changes in solar UV between 

solar cycles are assumed to be small. Effects on 

ozone from variations within each 11-year cycle 

can be isolated as demonstrated in Figure 3.2. 

Once identified, the effect can be removed in 

interpreting the observed ozone changes. Thus, 

it is not crucial whether the solar cycle effect 

is included in the simulations. Effects from 

volcanoes are not included as there is no reliable 

way to predict volcanic eruptions in the future. 

Their effects can be removed in the analyses 

several years after it occurs. The philosophy 

here is that, like the solar cycle, the effect can be 

removed from the observation before they are 

compared with the model simulated trends.

Variations in natural 

factors such as 

changes in the Sun’s 

energy output and 

volcanic events will 

continue to have 

impacts on the 

ozone abundances.
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Chapter 4 discussed how climate change 

due to increased CO2 (and other well-mixed 

greenhouse gases), changes in water vapor 

in the stratosphere, and changes in long-lived 

source gases (CH4 and N2O) could affect 

ozone. Climate change can affect ozone 

through changes in temperature and transport 

circulation. Cooling of the stratosphere 

associated with greenhouse gases is expected 

to slow gas-phase ozone loss reactions and 

increase ozone. This is particularly effective 

in the upper stratosphere. Water vapor in the 

stratosphere plays a particularly interesting 

role. It affects ozone concentration through 

the odd-hydrogen (HOx) chemistry, as well as 

contributing to the cooling of the stratosphere. 

Its concentration can be changed due to changes 

in methane and changes in climate. In the 

scenario calculations, changes in water are not 

prescribed. They are calculated from the CH4

increase and from changes associated with 

climate in chemistry climate models (CCMs). 

The scenario for CO2, CH4, and N2O used in 

the simulations is summarized in Table 5.1. 

Based on sensitivity simulations from twodimensional (2-D) chemistry transport models 

(CTMs) reported in WMO (1999, Chapter 12), 

a 15% increase in CH4 at 2050 from its 2000 

values would have added about 0.5% in column 

ozone at midlatitudes. A 15% N2O increase 

would have decreased ozone by about 1%. 

Thus, in the scenario shown, the combined 

effects from CH4 and N2O increase ozone 

around 2050.

The WMO reports also discussed changes 

in aerosol and oxides of nitrogen (NOx) 

from aviation (WMO, 2003, Chapter 4), 

emissions from rocket launches (WMO, 2003, 

Chapter 4), and changes in molecular hydrogen 

(H2) (WMO, 2007, Chapter 6). The effects 

from these processes are not included in the 

WMO simulations. Emission of NOx from 

subsonic airplanes increases ozone in the upper 

troposphere. The IPCC (1999) estimates that a 

0.4% increase in column ozone at midlatitudes 

in the current atmosphere can be attributed to 

en route emissions from aircraft. Anticipated 

doubling to tripling of emissions by 2050 could 

add another 1%. Detailed projections of future 

emissions based on demands and technology 

advances are not yet available. Previous 

estimates suggest that the current rocket launch 

schedule may have caused a small (less than 

1%) column decrease due to loss in the lower 

stratosphere. Future trends will depend on 

growth and mix of solid fuel and liquid fuel 

propellants. Estimates for changes in H2 are 

based on the assumption that liquid hydrogen 

may become an important energy source 

for the economy and leakage from storage 

and usage may cause a dramatic increase in 

H2. Not enough is known to do any reliable 

projections. 

5.2.2 Results From Model Simulations

Three types of models were used to simulate 

the future behavior of ozone in WMO (2007, 

Chapter 6): 

1. Two-dimensional chemistry-transport 

models (2-D CTMs) use fixed temperature 

and circulation. They are most useful for 

isolating the effects of different source 

gases; 

2. Interactive 2-D models partially account 

for the changes in transport associated 

with climate change by calculating the 

residual circulation from heating rates 

allowing for interaction of planetary waves 

with the mean circulation. However, the 

feedback from changes in wave forcing is 

not simulated; and 

3. Three-dimensional climate chemistry 

models (3-D CCMs) incorporate all the 

identified feedbacks and are generally 

Climate change can 

affect ozone through 

changes in temperature 

and transport 

circulation. Cooling 

of the stratosphere 

associated with 

greenhouse gases 

is expected to slow 

gas-phase ozone 

loss reactions and 

increase ozone.

Table 5.1 Future concentrations of CO2, CH4, and N2O used in the model simulations. The CO2

values are from the Integrated Science Assessment Model (ISAM) as listed in Appendix II of 

IPCC (2001).

Year 2000 2010 2020 2030 2040 2050 2060 2070 2080 2090 2100

CO2 (ppm) 369 391 420 454 491 532 572 611 649 685 717

CH4 (ppb) 1760 1871 2026 2202 2337 2400 2386 2301 2191 2078 1974

N2O (ppb) 316 324 331 338 344 350 356 360 365 368 372
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better able to represent the key processes 

related to 3-D transport in the atmosphere 

(particularly the polar regions). 

In the following discussion, both the observations 

and the model results will be displayed as 

annual mean or monthly anomalies expressed 

as a percentage of the pre-1980 conditions. The 

midlatitude EESC RD is expected to occur 

sometime between 2040 and 2050. In looking 

at the MS ORD, it has proven convenient to 

examine the spatial aspects of the problem 

in terms of the phenomena in the two polar 

regions (Arctic and Antarctic) and that in 

the tropics plus midlatitudes (approximately 

60°N to 60°S). This separation accounts for 

the distinct stratospheric circulation patterns 

prevailing in the climate system, is relevant 

for compartmentalizing approximately the 

ozone chemical-dynamical interactions, and 

represents a convenient way to look at the “big” 

global picture. 

 5.2.2.1 Tropics and midlatitudes

Figure 5.1 shows the simulated future behavior 

of column ozone from interactive 2-D models 

(solid lines) and non-interactive 2-D CTMs 

(dashed lines). The models’ hind-cast predictions 

are compared with observations as a way to 

screen the 2-D CTMs. All 2-D models show that 

ozone amount increases with time between 2007 

and 2050. The model spread among the noninteractive 2-D CTMs for northern midlatitudes 

is about 3% at 2050. The WMO (2007, Chapter 

6) report did not discuss how changes in N2O 

and CH4 contributed to the individual model 

results. Based on the estimates given above and 

the scenario stated in Table 5.1, it would appear 

that CH4 is adding about 1% while the effect 

of N2O is to decrease ozone by about 0.5% in 

2050. It is also evident from the figure that the 

sensitivities in these models differ. 

Results from the interactive 2-D models show 

that the ozone anomaly is larger by about 2% in 

2050 and 4% in 2100. The effect at midlatitudes 

is larger at about 3% in 2050. This is consistent 

with the expected ozone increase due to cooling 

in the stratosphere. There is no clear indication 

on the effect of increased upwelling in the 

tropics, though this could have been masked 

by the ozone increase in the upper stratosphere 

due to the cooling. 

Results from 3-D CCMs are shown in Figure 

5.2. Several tests were used to identify models 

that successfully simulate parameters important 

for ozone response to halogen loading (Eyring 

et al., 2006). Models that perform better in 

those tests are identified using solid lines in 

Figure 5.2. For our purpose, we concentrate on 

the three models (CCSRNIES, CMAM, and 

WACCM) that “earned” the solid line rating 

and performed the REF2 simulations from 1980 

to 2050. Other models performed the REF2 

simulation starting in 1990 or 2000 making 

it difficult to compare the ozone anomaly at 

2050 to the anomaly at 1980 to determine the 

All two-dimensional 

models used to 

simulate the future 

behavior of ozone 

show that ozone 

amount increases 

with time between 

2007 and 2050.

Figure 5.1 Simulated annual mean ozone anomaly from 2-D models for 

different latitude bands. Results from interactive models are designated by 

solid lines. The figure is based on Figure 6-9 in WMO (2007). See Eyring et 

al. (2006) for details on how the annual mean anomaly is computed. The 

black line with the grey shade represents the observed mean values and 

the range. The grey vertical band marks the time period when midlatitude 

EESC is expected to recover to the 1980 value.
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MS ODR. For ozone content between 60°S and 

60°N, the MS ORD are 2030 for WACCM and 

2040 for CMAM and CCSRNIES. All three 

models show little ozone increase beyond the 

1980 values in the tropics, consistent with 

the expectation that increase in upwelling 

is suppressing ozone. This is evident in the 

model-simulated decrease in tropical ozone 

below 20 hectopascals (hPa) (see Figure 6[b] 

in Eyring et al., 2007). The MS ORDs for 

northern midlatitudes are 2010 for WACCM, 

2020 for CMAM, and 2030 for CCSRNIES. 

The MS ORDs for the southern midlatitudes are 

all between 2030 and 2040. The CMAM model 

presented results through 2100. 

To isolate the effects of climate change, three 

CCMs performed a simulation where the 

surface concentrations of the greenhouse 

gases (GHGs) were kept fixed at their 1970 

value (Figure 5-25 in Chapter 5, WMO, 2007). 

The results from WACCM show that in the 

absence of these GHG forcings, the MS ORD 

for 60°S to 60°N is around 2040 and the ozone 

amount in 2050 is about 1% smaller compared 

to the baseline scenario that includes the GHG 

forcings. Unfortunately, the run also kept the 

surface concentrations of CH4 and N2O fixed. 

Thus, the 1% effect results from both climate 

change and the direct chemical effects of CH4

and N2O.

5.2.2.2 Polar region

Figure 5.3 shows the model-simulated ozone 

anomalies for the Arctic and the Antarctic 

regions. Most models show larger anomalies 

in the Antarctic, consistent with the fact that 

the temperature is colder, leading to formation 

of more polar stratospheric clouds (PSCs), and 

the vortex is more confining. Within almost all 

models, the Arctic ozone recovery date is much 

earlier than the Antarctic ozone recovery date. 

The polar EESC RD is estimated to be between 

2060 and 2070. We will again concentrate 

on the results from CCSRNIES, CMAM, 

and WACCM for the reason discussed in the 

previous section. The MS Arctic polar ORDs 

are 2000 for CCSRNIES, 2010 for CMAM, and 

2015 for WACCM. Once the ozone anomaly 

reaches the 1980 value, it increases smoothly 

to beyond the 1960 anomaly.

The exact time evolution of the Antarctic ozone 

hole is different depending on the diagnostics 

chosen. These include ozone amount in October, 

minimum ozone in September and October, 

ozone mass deficit, and maximum Antarctic 

ozone hole area between September and 

October. The minimum ozone value is projected 

not to increase until after 2010 in several models, 

while the decrease in ozone mass deficit in most 

models has occurred by 2005. If we use the 

ozone content poleward of 60° calculated by the 

three models as the metric, the MS ORDs are up 

to 30 years earlier. The CMAM and WACCM 

Almost all 3-D models 

predict that the Arctic 

ozone recovery 

date is much earlier 

than the Antarctic 

recovery date.

Figure 5.2 Results from 3-D CCMs. The figure is adapted from Figure 5 in 

Eyring et al. (2007) which includes additional model results computed after 

publication of the WMO (2007) report. See Eyring et al. (2006) for details 

on how the annual mean anomaly is computed. The solid line with the grey 

shade represents the observations with uncertainty. The grey vertical band 

marks the time period when midlatitude EESC is expected to recover to 

the 1980 value.
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models produced ensemble results. The three 

simulations from the WACCM model produced 

polar ozone recovery dates between 2030 and 

2040, while those from CMAM are between 

2040 and 2060. The CMAM results also showed 

that the value for the Antarctic ozone anomaly 

stays closed to zero for about 20 years after the 

initial recovery before taking off.

5.2.3 Stages of Ozone 

Recovery From ODSs

The model-simulated results presented in 

Figure 5-25 in Chapter 5 of WMO (2007) 

suggest that the ozone increase between now 

and 2025 is largely due to the decrease in 

halogen loading. ODSs will remain a driver of 

human-caused ozone depletion up until 2040, 

and reductions in emissions of these chemicals 

represent the only known acceptable method to 

reduce the associated ozone depletion expected 

in this period. The halogen loading is expected 

to approach its 1980 value toward the middle 

of the century at midlatitudes. In the decades 

that follow, the effects of climate change and 

changes in other trace gases will determine the 

ozone behavior. 

A good portion of Chapter 6 in the WMO 

(2007) report was devoted to discussion of 

the detection and attribution of the expected 

ozone recovery based on future observations. 

Because of interannual variability, it is not 

possible to identify an ozone recovery date from 

observation as soon as it occurs. If one were 

able to isolate the trend due to halogen loading 

from the observed ozone time series, one would 

expect this trend to first stop decreasing when 

EESC peaks (around 1997 for midlatitudes), 

followed by an increasing trend until the trend 

becomes zero again when halogen loading no 

longer has an effect on ozone. This is the trendderived halogen-induced ozone recovery date 

(TD H-ORD). The work of Weatherhead et al.

(2000) and Yang et al. (2006) clearly show that 

the length of observations required to detect 

such change depends on the quality of the data. 

Given the current results, it is anticipated that 

we should be able to confirm whether ozone is 

increasing due to decrease in halogen loading 

in the next five or six years. We are not in a 

position to recognize the TD H-ORD as soon 

as it occurs. Nonetheless, the simulations give 

confidence that one should be able to confirm 

the ozone recovery after the fact by waiting 

several years to analyze the observations and 

removing the interannual variability. 

In the context of the rest of this chapter 

(Sections 5.4 and 5.5), the most important 

attribution issue is whether EESC is a good 

proxy for future ozone behavior so that one 

could have confidence that policy decisions 

based on reducing EESC would achieve the goal 

of similar reductions in ozone depletion. Indeed, 

there are concerns (e.g., Hadjinicolaou et al., 

2005) that improper interpretation of the recent 

Figure 5.3 Zonal mean monthly ozone anomalies for the Arctic in 

March (panel a) and Antarctic in October (panel b). The figure is based 

on Figure 7 in Eyring et al. (2007), which was updated to include additional results after the publication of the WMO (2007) report. The 

observations are shown as black dots and a smooth curve representing 

the mean value. The grey vertical band marks the time period when 

polar EESC is expected to recover to the 1980 value.
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observed ozone increase after the late 1990s 

may give the wrong impression that the effects 

of halogens on ozone have been overestimated 

and one should relax the reduction strategy.

Chapter 6 in WMO (2007) identifies other 

factors that could complicate the attribution of 

the observed changes. These include changes 

in atmospheric composition of ozone-relevant 

compounds other than the halogens, changes 

in temperature and transport circulation, 

changes in solar cycle, and volcanic eruptions. 

The largest effect on the short-term (five to ten 

years) trend is expected to come from changes 

in transport circulation. The study of Yang et 

al. (2006) concluded that half of the observed 

increase in ozone between the late 1990s and 

2005 could be attributed to changes in transport 

circulation in the lower stratosphere. This is not 

unexpected because, while EESC has stopped 

increasing, it essentially remained unchanged 

during this period. Hadjinicolaou et al. (2005) 

had a similar conclusion using a very different 

approach. The authors use a 3-D CTM to 

calculate the ozone from 1979 to 2003. The 

CTM uses the transport circulation from the 

ERA-40 analyses of the European Centre for 

Medium-range Weather Forecasting (ECMWF). 

The ozone chemistry is parameterized with 

the local loss frequency fixed at the 1980 

conditions. The conclusion is that the ozone 

trends between 1994 and 2003 derived from the 

modeled and observed ozone agree, indicating 

that change in transport is the main driver in 

this time period. The paper also concluded 

that the model-calculated ozone (again with 

fixed loss frequency) showed a decreasing 

trend between 1979 and 1993, and the trend 

is around one-third of the trend derived from 

observation. More analyses (such as additional 

model results using full chemistry to show that 

the combined trend due to changes in transport 

and halogen loading is not significantly larger 

than the observed trend) are needed to support 

this last conclusion that changes in transport are 

responsible for one-third of the observed ozone 

trend between 1979 and 1993. 

After 2050, effects from other changes (changes 

in CH4, N2O, and CO2) will dominate. If the 

desire is to understand future ozone behavior 

beyond the effects of halogens, one must focus 

on the trends of the other ozone-relevant source 

gases and try to determine to what extent one 

could separate the effects on ozone in the future 

observations. 

5.3 EXPECTED RESPONSE IN 

SURFACE UV

Ozone column in the atmosphere is one of 

many factors that affect UV at the ground. The 

UV community emphasizes the importance 

of variations in aerosol, clouds, and surface 

albedo on UV. The effect of ozone change on 

clouds through climate feedback has not been 

quantified at this point, but is expected to be 

small.

If everything else is assumed to be constant, the 

future UV trend will depend on the anticipated 

ozone change. Within the limitation that applies 

to EESC as a proxy for future ozone behavior, 

it can likewise be used as a predictor for UV. 

However, most UV predictions are done locally 

at specific latitudes, thus the relationship 

between EESC and typical midlatitude ozone 

depletion is not particularly useful due to 

the other factors affecting UV. In practice, 

model-simulated ozone changes at specific 

latitudes are fed into a radiative transfer model 

to compute the change in UV irradiance. An 

example of such a calculation is shown in Figure 

5.4, which shows the calculated noon-time 

erythemal irradiance at several latitudes. Note 

that the recovery at the southern polar latitude 

occurs much later than the midlatitude values, 

reflecting similar behavior of midlatitude and 

polar ozone columns as indicated in the results 

from the AMTRAC model in Figure 5.2. 

5.4 FUTURE SCENARIOS 

FOR ODSs AND THEIR 

REPLACEMENTS

5.4.1 Baseline Scenario

In general, the historical portion of the baseline 

(A1) scenario is based on the observed mixing 

ratio time series, while future emissions are 

estimated using the most current information 

regarding expected future demand of ODSs, 

future banks (the amount of a chemical that 

has been produced but not yet emitted or 

chemically altered; see Box 2.5 in Chapter 2 

for a more detailed definition), and current 

constraints placed by the Montreal Protocol. 

If all other factors 

remain constant, the 

future ultraviolet (UV) 

trend will depend 

on the anticipated 

ozone change.
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While this scenario consists of reasonable 

assumptions about the future, it does not 

represent a prediction, and future levels could 

be higher or lower depending on, for example, 

future policy actions and consumer choices. 

However, it represents a useful projection 

that is used to examine the sensitivity of ODS 

abundances to choices concerning future 

production, banks, and emissions.

The mixing ratios used to calculate the 

historical emissions are obtained primarily 

from atmospheric observations made by the 

NOAA Earth System Research Laboratory/

Global Monitoring Division (ESRL/GMD) 

(formerly Climate Monitoring and Diagnostics 

Laboratory, [CMDL]), the Advanced Global 

Atmospheric Gases Experiment (AGAGE), and 

the University of East Anglia (for halon-1211). 

South Pole firn observations are also considered 

for methyl chloride (CH3Cl) and methyl 

bromide (CH3Br) emissions before 1996. A 

box model is used to determine the emissions 

of the species for each year through 2005 using 

the observed mixing ratio time series and its 

current best estimate of the steady state global 

lifetime. Hence, when the same box model 

and lifetimes are used to calculate the surface 

mixing ratios from the derived emissions, they 

produce mixing ratios in the baseline scenario 

that are exactly equal to the observationally 

based time series given in Table 8-5 of WMO 

(2007). The same box model and lifetimes are 

used to derive the mixing ratio of each species 

after 2005 based on projected emissions.

Projections of future demand and sizes of 

banks are taken from IPCC/TEAP (2005), and 

play a major role in the calculation of future 

emissions. Details relating to the production 

and emission projections can be found in 

WMO (2007). The use of future demand and 

bank sizes from IPCC/TEAP (2005) in WMO 

(2007) represents an important departure from 

the approach used in previous WMO reports. 

Previously, the evolution of the estimated bank 

sizes was calculated solely using the difference 

between estimated annual production and 

emission. This approach had the potential to 

lead to accumulating large errors in the bank 

Bank estimates of 

ozone-depleting 

substances are based 

on inventories of 

equipment containing 

these substances, an 

approach called the 

“bottom-up” method.

Figure 5.4 Estimated changes in erythemally weighted surface UV irradiance at local noon in response 

to projected changes in total column ozone as calculated by the AMTRAC CCM (see blue dashed 

curve in Figure 5.2) for the period 1970 to 2099, using zonal-averages in total ozone in the latitude 

bands 35°N-60°N, 35°S-60°S, and 60°S-90°S, and the solar zenith angle corresponding to 45°N in 

July, 45°S in January, and 65°S in October, respectively. At each latitude, the irradiance is expressed 

as the ratio to the 1970 to 1980 average. The results have been smoothed with a five-year running 

mean filter to remove some of the year-to-year variability in the ozone predictions in the model.
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sizes because the bank often represents a small 

difference between the two relatively large 

production and emission values. The IPCC/

TEAP (2005) bank estimates, however, are 

based on inventories of equipment, an approach 

often referred to as a “bottom-up” method. 

Hence, these estimates are independent of 

systematic errors in production or emission. 

It is believed that this new approach has led to 

better projections of future emissions.

Comparisons between future emissions 

projections of WMO (2003) and WMO (2007) 

demonstrate that the most substantial differences 

arise for CFC-11, CFC-12, carbon tetrachloride 

(CCl4), and the HCFCs. The increase in the 

CFC emissions in WMO (2007) is primarily 

due to larger bank estimates of the bottom-up 

approach than were estimated by WMO (2003). 

The greatest HCFC emission difference is for 

HCFC-22 and is due to the substantially larger 

estimated future consumption of this compound 

by Article 5(1) countries. Current and projected 

future carbon tetrachloride emissions are now 

estimated to be higher than those projected in 

WMO (2003) based on observed mixing ratios 

consistent with a smaller decrease in emissions 

over the last few years and the continued 

inability to account for all CCl4 emissions. 

The resulting differences in mixing ratios are 

discussed in Section 5.4.3.

5.4.2 Alternate Scenarios 

Alternative scenarios and test cases were 

examined in WMO (2007) to quantify the 

relative effects of making various production 

and/or direct emissions reductions on EESC. 

Three cases for different ODS groups are 

designed to address three issues: (1) no future 

emission; (2) no future production; and (3) 

eliminations of the 2007 bank. Results from 

the “no future emission” case provide the 

future mixing ratios due to the decay of the 

ODSs already in the atmosphere only. This case 

represents the greatest theoretically possible 

reduction in the future atmospheric burden of 

the particular compound (short of processing 

the air to remove the ODSs). The “no future 

production” case quantifies the importance of 

new production relative to future emissions, 

while the “bank elimination” quantifies the 

benefit of the one-time sequestration and 

destruction of the 2007 global bank for future 

emissions. Additional cases are presented 

here that examine the effect of recovering and 

destroying the total estimated U.S. bank and 

the U.S. accessible bank in 2009. Estimates of 

these bank sizes and the technique used by the 

U.S. Environmental Protection Agency (EPA) 

to calculate these estimates are discussed in 

Chapter 2.

WMO (2007) also examined three alternative 

cases involving CH3Br. Two cases involved 

removing quarantine and pre-shipment uses 

from 2015 onward and continuing Critical 

Use Exemptions at 2006 levels into the future. 

The third case explored the importance of 

the assumption that the 1992 anthropogenic 

emissions represented 30% of the total. Recent 

mixing ratio observations have suggested that 

this might be an overestimate with a more 

accurate percentage falling somewhere between 

20% and 30%. These results are presented later 

in Section 5.5 and in Table 5.2. 

A scenario based on the mitigation scenario 

described in IPCC/TEAP (2005) is also 

examined to quantify the effect of this carefully 

considered set of future policy options. The 

mitigation scenario only has a substantial 

effect on the bank of HCFC-22 in the scenario 

considered here.

After the WMO (2007) report and IPCC 

(2007) reports were written, the Parties to 

the Montreal Protocol voted to strengthen the 

HCFC regulations on both Article 5(1) and 

non-Article 5(1) countries. Approximations for 

the effect of this strengthening are discussed in 

Section 5.5.1.1.

5.4.3 Time Series of Source Gases

The mixing ratios of the current baseline 

scenario are compared with those of the WMO 

(2003) baseline scenario, and the “no future 

production” and “no future emission” test 

cases in Figure 5.5. The differences between 

the WMO (2007) and WMO (2003) baseline 

scenarios are apparent for several gases, with 

the differences for HCFC-22 particularly 

apparent. More modest, but also important are 

the differences for CFC-11 and CFC-12. The 

HCFC-22 difference is due to the increase in 

the expected future consumption of Article 5(1) 

countries of the Montreal Protocol, while the 
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increase in the CFCs is due to the larger bank 

size estimates of IPCC/TEAP (2005). HCFC141b, HCFC-142b, and halon-1301 show reduced 

mixing ratios in the short term compared to 

WMO (2003) owing to the reduced observed 

growth rates between 2001 and 2004 and the 

expectation of lower future emissions. 

The importance of future projected production 

and bank sizes to future emissions is also 

apparent for the various compounds. For 

example, the “no future production” curve 

for CFC-12 is only slightly different from the 

baseline curve; hence the bank of CFC-12 is 

expected to dominate future emissions, with 

its effect represented by the difference between 

the “no future emission” and “no future 

production” curves. The relative importance of 

future production compared to the amount in 

the banks varies strongly among the ODSs, with 

the future abundances of CFC-11 depending 

primarily on its bank size and HCFC-22 future 

abundances depending primarily on future 

production. No bank is considered in the future 

projections of methyl chloroform (CH3CCl3), 

CCl4, and CH3Br.

5.5 CHANGES IN INTEGRATED 

EESC AND RADIATIVE FORCING

5.5.1 Time Series of EESC

5.5.1.1 Midlatitudes

The evolution of ODS mixing ratios cannot, 

by themselves, be used to accurately quantify 

the ozone destruction due to those ODSs. The 

established relationship between stratospheric 

ozone depletion and inorganic chlorine and 

bromine abundances suggest that the temporal 

evolution of inorganic chlorine- and brominespecies in the stratosphere is an important 

indicator of the potential damage of human 

activity on the health of stratospheric ozone. 

Equivalent Effective Stratospheric Chlorine 

(EESC) was developed to relate this halogen 

evolution to tropospheric source gases in a 

simple manner (Daniel et al., 1995; see also Box 

2.7 in Chapter 2). This quantity sums ODSs, 

accounting for a transit time to the stratosphere, 

for the number of halogen atoms in the ODS, 

for the greater per-atom potency of stratospheric 

bromine (Br) compared to chlorine (Cl) in its 

ozone destructiveness with a constant factor (α), 

and also includes the varying rates with which 

Cl and Br will be released in the stratosphere 

from different source gases. EESC has been 

used as a proxy for the effect of humanproduced ODS abundances on future ozone 

depletion (WMO, 1995, 1999, 2003, 2007). 

The values for midlatitude EESC discussed 

here were calculated for WMO (2007) from the 

global averaged surface mixing ratios for the 

ODSs, using a constant lag time of three years 

Figure 5.5 Mixing ratio comparisons of the WMO (2007) baseline scenario 

(solid black) with the baseline scenario from WMO (2003) (green), the “no 

future emission” test case (dashed), and the “no future production” case 

(dotted curve). The figure is based on Figure 8-2 in WMO (2007). Note that 

different vertical scales are adopted for sub panels and some of the plotted 

values do not start from zero. For several of the gases, the solid black curve 

obscures the dotted or dashed curves.
The U.S. Climate Change Science Program Chapter 5

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and release factors given in the same report. 

Recent development on how to apply EESC to 

polar ozone and refinements in using the mean 

age of air will be discussed in Sections 5.5.1.2 

and 5.5.2, respectively. 

The relative contribution of various ODSs and 

ODS groups to midlatitude EESC are shown 

as a function of time from 2000 to 2100 on the 

left-hand side of Figure 5.6. The prominent role 

of CFCs today and into the future is apparent. 

The slow decline of the contribution from 

CFCs is primarily due to the relatively long 

atmospheric lifetimes of the species in this 

group of compounds and what is already in the 

atmosphere, and not to continued emission, 

although continued emission does play a 

small role. The importance of the halons and 

CH3Br, all bromine-containing source gases, 

is also clear even though their atmospheric 

concentrations are substantially smaller than 

those of the chlorine-containing ODSs. This 

is because stratospheric bromine is much 

more effective per atom than chlorine for 

stratospheric ozone destruction. As stated in 

Chapter 2, WMO (2007) has estimated that 

an atom of bromine is 60 times more effective 

than an atom of chlorine for global stratospheric 

ozone destruction. The lower panels show 

the change in EESC due to the elimination of 

production and emission for CFCs, HCFCs, 

halons, and CH3Br.

In the past, EESC estimates have been used 

to evaluate various ODS emissions scenarios 

primarily using two metrics. They are: (1) a 

comparison of the times when EESC returns to 

1980 levels, the EESC RD; and (2) the relative 

integrated changes in EESC between 1980 or 

the current time and the corresponding EESC 

RD. Figure 5.7 demonstrates that the return of 

midlatitude EESC to the 1980 level is currently 

expected to occur around 2049 for the baseline 

scenario, five years later than projected in 

WMO (2003). This later return was primarily 

ascribed to higher estimated future emissions 

of CFC-11, CFC-12, and HCFC-22. The increase 

in CFC emissions is due to larger estimated 

current bank sizes, while the increase in 

HCFC-22 emissions is due to larger estimated 

future production. The soonest that a complete 

theoretical elimination of emissions could lead 

to a return to 1980 levels is 2034. Elimination 

of all future ODS production is expected to lead 

to a return to 1980 EESC levels in 2043, while 

an elimination of the 2007 bank is expected to 

lead to a return in 2041. 

A detailed partitioning of the effects of 

reductions in the various ODS groups is shown 

in Table 5.2. The years when EESC is expected 

to return to 1980 levels are also included in 

the table for midlatitudes and for the Antarctic 

vortex. The Antarctic ozone response to 

EESC will be discussed in more detail in 

Section 5.5.1.2. The table illustrates that the 

elimination of the future emissions of CFCs, 

HCFCs, and halons represents the greatest 

potential for reducing future ozone depletion. 

To accomplish this elimination for CFCs and 

halons, banks would have to be captured and 

destroyed because future emission is expected 

to be dominated by the release from banks. For 

HCFCs, future production plays a larger role 

in future emission than do the current banks, 

so emissions from both banks and future 

production would have to be eliminated. The 

technical difficulty and expense involved with 

capturing banks depends on the nature of the 

banks, while the feasibility of reducing future 

production will depend on replacement options 

for the pertinent applications. More details 

concerning the nature of the various ODS uses 

and bank types and the options available for 

reducing future ODS emissions can be found 

in other reports, including IPCC/TEAP (2005) 

and UNEP/TEAP (2007). It should also be 

recognized that these full bank recovery and 

zero production and emission test cases shown 

in Table 5.2 are meant as hypothetical cases 

against which more realistic scenarios can be 

compared. This procedure is used in Section 

5.6 to evaluate the significance of the United 

States ODS banks.

The results for the scenario representing the 

IPCC/TEAP (2005) mitigation scenario are 

not shown in Table 5.2, but this scenario leads 

to an EESC response that is approximately 

20% of the zero-emission case for the HCFCs, 

due primarily to actions that would reduce the 

HCFC-22 bank emission. 

Future emissions of CH3Br also have the 

potential to be as important as each of these 

three classes of compounds. The continuation 

In late 2007, the 

Montreal Protocol 

HCFC restrictions for 

both production and 

consumption were 

strengthened, partly 

in response to the 

renewed awareness 

of the importance of 

HCFCs to climate 

forcing in addition 

to ozone depletion.
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Table 5.2 Comparison of scenarios and casesa: the year when EESC drops below the 1980 value (i.e., 

EESC RD) for both midlatitude and polar vortex cases, and integrated EESC differences (midlatitude 

case) relative to the baseline (A1) scenario. 

Scenario

Percent Difference in 

integrated EESC relative 

to baseline scenario for the 

midlatitude case

Year (x) when EESC is expected to 

drop below 1980 value

Midlatitude Antarctic Vortexb

∫ EESC dt ∫ EESC dt

Scenarios

A1: Baseline Scenario 2049 2065

Casesa

 of zero production from 2007 onwards of:

All ODSs −8.0 −17.1 2043 2060

CFCs only −0.1 −0.3 2049 2065

Halons only −0.2 −0.5 2049 2065

HCFCs only −5.5 −11.8 2044 2062

Anthropogenic CH3Br only −2.4 −5.1 2048 2063

Casesa

 of zero emissions from 2007 onwards of:

All ODSs −19.4 −41.7 2034 2050

CFCs only −5.3 −11.5 2045 2060

CH3CCl3 only −0.1 −0.2 2049 2065

Halons only −6.7 −14.4 2046 2062

HCFCs only −7.3 −15.7 2044 2062

CCl4 only −1.3 −2.9 2049 2065

Anthropogenic CH3Br only −2.4 −5.1 2048 2064

Casesa

 of full recovery of the 2007 banks of:

All ODS −12.9 −27.8 2041 2057

CFCs only −5.2 −11.3 2045 2060

Halons only −6.7 −14.3 2046 2062

HCFCs only −1.9 −4.1 2048 2065

CH3Br sensitivity:

Same as A1, but CH3Br anthropogenic emissions set to 20% 

in 1992c

3.1 6.6 2051 2068

Same as A1, but zero QPS production from 2015 onwards −1.5 −3.2 2048 2064

Same as A1, but Critical Use 

Exemptions continued at 2006 

level

1.9-2.2 4.0-4.7 2050 2067

a

 Importance of ozone-depleting substances for future EESC were calculated in the hypothetical “cases” by setting 

production or emission of all or individual ODS groups to zero in 2007 and subsequent years or the bank of all ODS or 

individual ODS groups to zero in the year 2007 alone. These cases are not mutually exclusive and separate effects of 

elimination of production, emissions and banks are not additive.

b

 Calculated using a lag time of six years and the same release factors as in midlatitudes.

c

 In the baseline scenario this fraction was assumed to be 30% in 1992 with a corresponding emission fraction of 0.88 

of production. In this alternative scenario an anthropogenic fraction was assumed to be 20% with an emission fraction 

of 0.56 of production. In both scenarios the total historic emission was derived from atmospheric observations and a 

lifetime of 0.7 years.

x

2007

x

1980
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Figure 5.6 EESC and direct radiative forcing estimates from 2000 to 2100 for the A1 baseline 

scenario in WMO (2007) (upper panels), and expected decreases relative to the baseline scenario 

due to a cessation of emission (dashed curves) and production (dotted curves) in 2007 for CFCs, 

HCFCs, halons, and anthropogenic CH3Br. The figure is adapted from Figure 8-5 in WMO (2007) 

except for the addition of the indirect forcing in the upper right hand panel. Note the difference in 

the vertical scales between the top panel and the bottom four panels. The “no production” curve 

for CFCs and halons lies almost on the zero line, indicating that future productions of these ODSs 

play a very small role in the baseline scenario. In contrast, the contribution from the HCFCs is 

mostly due to future productions. The “no emission” and “no production” curves are identical for 

CH3Br because no bank was considered in its projections. The HFC forcing is shown for the B1 

and A1B SRES scenarios. The indirect forcing due to ozone depletion caused by ODSs is included 

for comparison, but should be considered only a rough approximation owing to reasons discussed 

in the text of this chapter and of Chapter 2.
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Trends in Emissions of Ozone-Depleting Substances, Ozone 

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of the Critical Use Exemption at the 2006 level 

and the continuation of QPS uses both have a 

substantial impact on global EESC. 

In late 2007, the Montreal Protocol HCFC 

restrictions for both production and consumption 

were strengthened, partly in response to the 

renewed awareness of the importance of 

HCFCs to climate forcing in addition to ozone 

depletion. While restrictions were tightened for 

both Article 5(1) and non-Article 5(1) countries, 

the changes for the Article 5(1) countries are 

much more significant for stratospheric ozone. 

In Figure 5.8, the former Protocol HCFC 

restrictions for Article 5(1) countries are 

compared to the newly adopted ones, as well as 

to the United States proposal that contributed 

to the strengthened restrictions. An additional 

curve is also shown that represents the closest 

scenario calculated in UNEP/TEAP (2007). 

For the TEAP scenario, it is estimated that 

integrated EESC is reduced by 2.6% and 5.6%, 

respectively, for the integration from 1980 to the 

return of EESC to 1980 levels and from 2007 to 

the return to 1980 levels. This is a substantial 

reduction even when compared to the zero 

emissions case for HCFCs in Table 5.2. The 

baseline HCFC emissions are slightly higher 

in UNEP/TEAP (2007) than those assumed 

in WMO (2007), making the effect of HCFC 

reductions correspondingly slightly higher. This 

TEAP report also examines other “practical 

options” that could be usefully employed to 

reduce future emissions of 

HCFCs and other ODSs. These 

include emission reduction 

measures during the use phase 

of applications and equipment, 

design and material section 

alter natives, end-of-life 

ma nagement , a nd ea rly 

retirement of equipment. These 

measures were submitted by 

the Parties to the Montreal 

Protocol and organized at the 

26th Open-ended Working 

Group Meeting of the Parties 

to the Montreal Protocol. 

The TEAP report finds that 

a combination of earlier 

HCFC phase-out described 

above with these additional 

“practical measures” leads 

to an integrated EESC reduction of 7.4% and 

16.0% percent, respectively, for the integration 

from 1980 to the return of EESC to 1980 levels 

and from 2007 to the return to 1980 levels.

5.5.1.2 Polar regions

Compared to midlatitude EESC, Arctic EESC is 

less useful as a proxy for polar ozone depletion 

because interannual variability in meteorology 

has a much larger impact on the ozone response 

For midlatitudes, the 

test case involving 

elimination of all 

future ozone-depleting 

substance production 

is expected to lead 

to 1980 Equivalent 

Effective Stratospheric 

Chlorine levels in 2043.

Figure 5.7 Midlatitude EESC estimates from 1980 through 2050 for 

the baseline scenario and the three comparative test cases considered in 

WMO (2007). The horizontal line represents the 1980 EESC level. This 

figure is redrawn from Figure 8-4 of WMO (2007).

Figure 5.8 Comparison of alternate scenarios for future emissions of HCFCs.
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150

to inorganic halogen loading. In the core of 

the Antarctic vortex during early spring, the 

interannual variability is smaller, suggesting 

that EESC could provide a useful proxy for 

ozone hole recovery (Newman et al., 2006). The 

far right column in Table 5.2 shows the results 

calculated using a time lag of six years and 

the same release factors based on midlatitude 

measurements. Because of the larger time-lag, 

the polar EESC value in 1980 is smaller than 

the 1980 midlatitude value. In addition, the 

larger lag time also makes the polar EESC value 

larger than the midlatitude EESC value in 2050. 

Therefore, the polar EESC RD is 15 to 17 years 

later than the midlatitude EESC RD. 

5.5.2 EESC and Mean Age of Air 

Previous EESC calculations have not included 

a distribution of transport times from the 

tropopause into the stratosphere (the so called 

age-of-air spectrum) or any dependence of the 

fractional chlorine release value on the age-ofair. Newman et al. (2006) reformulated EESC 

to account for both an age-of-air spectrum and 

age-dependent fractional release rates. Those 

results were discussed in Box 8-1 of WMO 

(2007). In this section, we will summarize how 

the EESC estimates derived from Newman et 

al. (2006) differ from the results in Section 

5.5.1.2.

The dashed lines of Figure 5.9 show EESC for 

mean ages of three years (blue) and six years 

(red) as estimated using the Newman et al.

(2006) technique, while the solid lines show 

EESC for constant age shifts (time-lags) of three 

years (blue) and six years (red). The solid lines 

duplicate the EESC used in WMO (2007). The 

main difference in the recovery dates between 

the two methods in each case (midlatitude and 

polar) is a result of the differences in fractional 

release values. In the WMO (2007) case, the 

release factors are fixed values, while the release 

factors in the Newman et al. (2006) curves are 

mean-age dependent. Note that the Newman et 

al. (2006) release factors at midlatitudes lead 

to smaller EESC values relative to 1980 levels 

than the values used in WMO (2007) and that 

an earlier EESC RD is projected. In contrast, 

the Newman et al. (2006) release factors at the 

Figure 5.9 Comparison of EESC values calculated using a lag time and fixed fractional release values 

(solid) vs. those calculated using a mean age with an age spectrum and fractional release values parameterized as functions of mean age (dashed). The blue curves are for midlatitudes and are calculated with 

a mean age and lag time of three years. The curves for the polar region are in red and are calculated 

with a mean age and lag time of six years.
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pole for the six-year mean age are larger than 

the WMO (2007) values, and result in a later 

EESC RD. 

Newman et al. (2006) raised the issue of 

the uncertainty in predicting the EESC RD 

associated with the choice of mean age and 

release factors to represent midlatitude and 

polar conditions. While the use of different 

mean ages would change the absolute timing of 

the recovery for the baseline case and other test 

cases, we are reasonably confident that it would 

not change the conclusion about the relative 

effects of different test cases. 

5.5.3 Time Series of Radiative Forcing

To adhere to the requirements of the Montreal 

Protocol, several courses of action have been 

adopted, including not-in-kind replacements 

of ODSs and changes in operations that 

reduce emissions. Applications that previously 

used CFCs are now performed with CFC 

replacements, HCFCs and HFCs, with HFCs 

likely to play a larger role in the future as 

HCFCs are phased out by the Montreal Protocol. 

Because HFCs contain no chlorine, bromine, or 

iodine, they do not destroy stratospheric ozone 

and therefore are not considered in WMO 

(2007). Furthermore, because future HFC 

emissions and production are not regulated by 

the Montreal Protocol, as are HCFCs, future 

projections of HFC concentrations are generally 

much more uncertain than those of ODSs and 

are heavily dependent on future economic 

growth assumptions and policy decisions. The 

forcing from HFCs will be included here as 

part of the discussion. However, it should be 

pointed out that the replacement strategy may 

also involve changes in other greenhouse gas 

emissions associated with the life cycle analyses 

(IPCC/TEAP, 2005). The change in forcing 

associated with those greenhouse gases is not 

included in the discussion.

Once the mixing ratio time series has been 

determined, it is a simple matter to estimate 

the future direct radiative forcing due to the 

various compound groups from the radiative 

efficiencies of each ODS (WMO, 2007). This 

forcing time series, calculated by multiplying 

the mixing ratio time series by the radiative 

forcing efficiencies of the particular ODSs, 

is shown on the right-hand side of Figure 5.6. 

Direct radiative forcing from ODSs and HFC 

replacement chemicals is approximately 0.34 W 

per m2

 in 2005 and is expected to stay below 0.4 

W per m2 through 2100 (according to the WMO 

(2007) scenario for ODSs and SRES A1B and 

B1 scenarios for HFCs). This is to be compared 

with forcing from CO2 of 1.66 W per m2 in the 

2005 atmosphere, increasing to as high as 5 W 

per m2 by 2100 for the SRES A1B scenario. The 

continued importance of the CFCs, along with 

the importance of the HCFCs, are perhaps the 

most striking features of this figure. The direct 

forcing contributions of the halons and CH3Br 

are small because of their small atmospheric 

mixing ratios. The effect of the bromocarbons 

on ozone depletion is enhanced because of 

the higher per-atom efficiency of bromine 

compared to chlorine in destroying ozone; such 

a chemically-caused enhancement does not 

apply to radiative forcing. 

The potential reduction in direct forcing due 

to the elimination of future production and 

emission is shown for CFCs, HCFCs, halons, 

and CH3Br in the lower four panels of Figure 

5.6. Overall, direct forcing from CFCs is 

expected to decrease to 0.1 W per m2 by 2100. 

Direct forcing from HCFCs and other ODSs are 

expected to be negligible by 2100. It is evident 

that elimination of future HCFC emission has 

the largest effect on radiative forcing of the 

ODSs among the test cases considered here. 

This peak forcing reduction of almost 0.07 W 

per m2 in 2040 represents slightly less than 5% 

of the CO2 radiative forcing in 2000 and less 

than half of the N2O radiative forcing in 2000. 

In comparison, elimination of future CFC 

emissions will bring about a reduction of 0.015 

W per m2 in 2040.

The forcing of the HFCs, generally used as 

replacements for the ODSs, are included in 

the figure as the orange shaded region for the 

SRES (Nakićenović et al., 2000) A1B scenario. 

The line within the orange region represents 

the alternate forcing of the HFCs in the B1 

SRES scenario. Atmospheric observations 

suggest that the 2000 forcing due to the HFCs 

is slightly larger than that estimated by the 

SRES scenarios, primarily due to the higher 

abundances of HFC-23 observed. Projected 

forcing from HFCs is 0.15 W per m2

 and 0.24 

W per m2 in 2050 and 2100, respectively, for 

To adhere to the 

requirements of the 

Montreal Protocol, 

several courses of 

action have been 

adopted, including notin-kind replacements 

of ozone-depleting 

substances and 

changes in operations 

that reduce emissions.
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the SRES A1B scenario, while other scenarios 

indicate that it will be lower. 

The direct forcing of the ODSs provides 

only a partial story of their total radiative 

forcing. Their destruction of ozone leads to an 

additional, indirect forcing that is thought to 

be significant but remains highly uncertain. 

An estimate of the indirect forcing from the 

ODSs is shown in Figure 5.6 as the red hatched 

region. It represents an uncertainty range of 

±100% around the best estimate for the period 

1979-1998 (i.e., −0.05±0.05 W per m2), taken 

from IPCC (2007). While the ozone forcing itself 

is highly uncertain, additional uncertainties 

exist that are not quantified here associated 

with, for example, the simplifying assumption 

of a linear relationship between ozone depletion 

and EESC above the 1980 threshold and even 

the amount of the observed ozone trend that is 

due to ODSs. The central forcing value (−0.05 

W per m2) suggests that ozone depletion offsets 

about one-sixth of the total direct ODS forcing 

around 2000. This figure also shows that the 

indirect forcing will gradually diminish in 

the coming decades, and is expected to return 

to near zero around 2050 as ozone becomes 

insensitive to the level of ODSs. This results 

from the assumption that the EESC value in 

1980 represents a threshold, below which ozone 

depletion does not respond to changing EESC 

levels. While such a picture is likely imperfect, 

global ozone data do suggest that the response 

of ozone to EESC may have changed at this 

EESC level. Finally, there have been studies 

suggesting that ozone radiative forcing may 

lead to a substantially different temperature 

response than does the same radiative forcing 

change from CO2 (Joshi et al., 2003 and 

references therein). Such a situation would 

imply that direct and indirect radiative forcing 

comparisons, as shown in Figure 5.6, could 

be misleading in estimating climate response. 

Additional details concerning the indirect 

forcing of the ODSs due to ozone depletion can 

be found in Chapter 2, particularly in Box 2.2. 

Accurate quantification of this forcing and of 

indirect GWPs currently represents a scientific 

gap due to the significance of the associated 

uncertainties. 

5.6 United States

Contributions to

EESC and to Radiative

Forcing by odss and their

replacements

Because of the long-lived nature of the ODSs, 

EESC and the radiative forcing arising from 

emissions of these compounds should be 

thought of as global quantities. This allows 

the contribution to midlatitude EESC and 

global radiative forcing to be apportioned to 

individual countries if their emissions are 

accurately known. As discussed in Chapter 2 in 

this Report, ODS production and consumption 

amounts for the United States are reported to 

the United Nations Environment Programme 

(UNEP), as required by the Montreal Protocol 

(UNEP, 2007). Data are also compiled by the 

Alternative Fluorocarbons Environmental 

Acceptability Study (AFEAS, 2007) for certain 

ODSs and for HFC-134a, although the amount 

reported to AFEAS has represented a smaller 

fraction of global production in the last decade 

or so when compared with the UNEP data. 

A discussion of the comparison of the data 

from these two compilations with observed 

atmospheric mixing ratio observations is 

provided in Chapter 2. Also in response to a 

requirement associated with the United States 

being a signatory to the Montreal Protocol, 

the U.S. EPA uses a vintaging model to 

estimate annual emissions from the estimated 

production and consumption values after 

1985. Chapter 2 discusses the results from the 

U.S. EPA’s vintaging model through the past 

and the assumptions made to estimate United 

States emissions prior to 1984. Here, we use 

these assumptions along with the U.S. EPA’s 

projections to estimate future levels of source 

gases attributable to the United States and 

their contributions to both EESC and radiative 

forcing. 

5.6.1 Contribution to EESC

It is useful to note that the EESC in 2030 will 

be 2400 ppt, with about one-third from the 

natural CH3Cl and CH3Br. For the remaining 

two-thirds attributed to manmade emissions, 

about 15% is due to emissions prior to 1975, 

and about 20% is due to emissions between 

1975 and 1984. The contributions from United 

In addition to direct 

forcing by ozonedepleting substances, 

the destruction of 

ozone by ozonedepleting substances 

also leads to an 

indirect forcing that 

is thought to be 

significant but remains 

highly uncertain.
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States emissions to the loading in 2030 due to 

manmade emissions are 4.5 to 9% from United 

States pre-1975 emissions, 2 to 9% from United 

States emissions between 1975 and 1984, and 9 

to 19% from United States emissions after 1985. 

Summing the contributions, we estimate that 

in 2030 the midlatitude anthropogenic EESC 

amount resulting from United States emissions 

represents about 15 to 37% of the EESC 

amount resulting from all global anthropogenic 

emissions. 

5.6.2 Contribution to Radiative Forcing by ODSs and Their Replacements

The United States’ contribution to global direct 

radiative forcing from ODSs, HFCs, and PFCs 

is expected to be 19 to 41% by 2030. The lower 

end of this range remains roughly constant 

until 2030, while the upper end gradually 

declines from about 47% in 2010. As was done 

for the previous EESC contribution results, 

these forcing estimates are calculated only 

considering anthropogenic emissions. Future 

United States perfluorocarbon (PFC) and 

sulfur hexafluoride (SF6) emissions, which 

are projected to contribute very little to future 

The United States’ 

contribution to global 

direct radiative forcing 

from ODSs, HFCs, 

and PFCs is expected 

to be 19 to 41% by 

the year 2030.

Table 5.3 Comparison of global and United States bank elimination projections in terms of 

integrated EESC and EESC RD. The global test cases are taken from WMO (2007) and consider 

elimination of the global bank in 2007. U.S. cases assume elimination of the full U.S. bank, or 

the accessible U.S. bank in 2009. 

Scenario

Percent Difference in integrated 

EESC relative to baseline scenario 

for the midlatitude case

Year (x) when EESC is 

expected to drop below 

1980 value

Midlatitude Antarctic 

Vortex

∫ EESC dt ∫ EESC dt

Scenarios

A1: Baseline Scenario 2048.9 2065

Cases of full recovery of the 2007 banksb of:

B0: All ODS (global) −12.9 −27.8 2040.8 2056.7

 CFCs (global) −5.2 −11.3 2045.1 2060.4

CFC-11 (U.S., accessible) −0.0a −0.0a 2048.1 2064.1

CFC-12 (U.S., accessible) −0.0a −0.1 2048.9 2065.0

CFC-11 (U.S., total) −1.1 −2.3 2048.1 2064.1

CFC-12 (U.S., total) −0.3 −0.7 2048.7 2064.8

Halons (global) −6.7 −14.3 2045.7 2062.0

halon-1211 (U.S., accessible) −0.1 −0.2 2048.9 2065.0

halon-1301 (U.S., accessible) −0.3 −0.6 2048.7 2064.8

halon-1211 (U.S., total) −0.1 −0.2 2048.9 2065.0

halon-1301 (U.S., total) −0.3 −0.6 2048.7 2064.8

HCFCs (global) −1.9 −4.1 2048.4 2064.8

HCFC-22 (U.S., accessible) −0.3 −0.6 2048.8 2065.0

HCFC-22 (U.S., total) −0.3 −0.7 2048.8 2065.0

HCFC-141b (U.S., total)c −0.4 −0.8 2048.8 2065.0

HCFC-142b (U.S., total)c −0.1 −0.1 2048.9 2065.0

a Values reported as −0.0 are smaller than 0.05% in magnitude. b Note that the U.S. numbers are for 2009 banks.

c Accessible bank values for HCFC-141b and HCFC-142b are not provided because the U.S. EPA estimates a zero 

accessible bank size for these compounds.

x

2007

x

1980
The U.S. Climate Change Science Program Chapter 5

154

radiative forcing (less than 10−5 W per m2

through 2030), are estimated from the U.S. 

EPA vintaging model, while the future global 

abundances of these compounds are taken from 

the A1B and B1 SRES scenarios (Nakićenović 

et al., 2000). If global PFCs and SF6 were not 

considered in the radiative forcing calculation, 

United States emissions are projected to 

contribute about 20 to 43% of global radiative 

forcing by ODSs and their replacements in 

2030.

5.6.3 Options for United 

States ODS Banks

The accessible and total bank size estimates 

for United States equipment and applications 

in 2005 are compared to global bank estimates 

from WMO (2007) in Table 5.3. Additional 

reduction cases are shown in Table 5.3 that 

quantify the importance of the recovery 

and destruction of United States total banks 

and United States accessible banks. The 

U.S. EPA has defined “accessible” banks 

as the quantity of ODSs that is contained in 

equipment (i.e., fire protection equipment 

and refrigeration/air conditioning systems). 

Furthermore, it is assumed that the amount of 

ODS recoverable from this equipment is equal 

to the full equipment charge minus the average 

annual loss rate from leakage and servicing. 

It is possible that some of the non-accessible 

banks could be recovered and destroyed 

with appropriate policy measures, marketbased incentives, and/or certain technological 

advances. Table 5.3 shows that the halon-1301 

and HCFC-22 United States accessible banks 

are the most substantial in terms of contributing 

to potential future integrated EESC reductions. 

If the total United States banks are considered, 

CFC-11, HCFC-141b, CFC-12, HCFC-22, and 

halon-1301 banks are most important. The 

calculations for the United States halon banks 

do not include stockpiles, and so should be 

considered underestimates of the full potential 

benefit of recovering and destroying these 

banks.
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Implications for the United States

CHAPTER

6

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

Lead Authors: John S. Daniel, NOAA; David W. Fahey, NOAA; 

Jay R. Herman, NASA; Stephen A. Montzka, NOAA; Malcolm Ko, 

NASA; Paul A. Newman, NASA; Richard Stolarski, NASA

Convening Lead Authors: A.R. Ravishankara, NOAA; 

Michael J. Kurylo, NASA

6.1 INTRODUCTION

The depletion of the stratospheric ozone layer due to human-produced ozone-depleting substances (ODSs) is a global phenomenon. Emissions of ODSs from around the globe contribute 

to depletion of the ozone layer throughout much of the stratosphere. ODSs emitted from 

different locations are well mixed within most of the lower atmosphere before they reach the 

stratosphere, where they contribute to chemical ozone depletion. Consequently, ozone depletion above a specific location is caused collectively by ODS emissions from different locations 

around the globe.

The observed pattern of the ozone depletion is not uniform around the globe; depletion above 

one region may differ from that above another region. However, this is not because of variations 

in emissions among regions, but because chemical and dynamical processes in the stratosphere 

cause regional variations in ozone and ozone loss rates. The extent of ozone depletion over a 

given region also varies with season and its overall magnitude changes with time. Consequently, 

the increase in ultraviolet (UV) light at the Earth’s surface due to the depletion of the stratospheric ozone layer also varies with region and time.

Because of these factors, a simple connection cannot be drawn between emissions of ODSs 

from a country or a region and the depletion of stratospheric ozone above that country or 

region. For example, there is substantial ozone depletion each austral spring over Antarctica, a 

continent with essentially no emissions of ozone-depleting substances. In contrast, the decrease 

in stratospheric ozone at northern midlatitudes, where the dominant emissions of ODSs occur, 

is significantly less.

The decades of release of ODSs, the associated decreases in stratospheric ozone and increases 

in surface UV radiation, along with the influence of ozone depletion and of ODSs on climate, 

have important implications for the United States. These implications can be viewed by examining three areas: impacts, accountability, and potential management options. Each area will be 

summarized in the following sections. The discussion of impacts will highlight past, present, and 

future changes in stratospheric ozone, surface UV radiation, and globally averaged radiative forcing. The section on accountability will address the United States’ contributions to the emissions 

of ODSs and the associated contributions to Equivalent Effective Stratospheric Chlorine (EESC) 

and radiative forcing of the Earth’s climate. In the section on potential management options, we 

will attempt to provide a scientific perspective of ODS issues that can be used for supporting 

future management decisions. These issues will be discussed generically without addressing any 

specific option.

Ozone depletion 

above a specific 

location is caused 

by the emissions 

of ozone-depleting 

substances from 

around the globe.
Ozone above the 

continental United 

States and other 

northern latitudes 

decreased to a 

minimum in 1993 

and has increased 

since that time.

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6.2 IMPACTS

The changes in stratospheric ozone and surface 

UV radiation vary considerably among regions 

of the United States, which stretches over a wide 

range of latitudes in the Northern Hemisphere. 

Arctic ozone losses impact Alaska most 

prominently, while subtropical ozone changes 

affect Hawaii as well as Guam, Puerto Rico, 

and other United States territories.

6.2.1 Changes in Ozone 

over the United States

The changes in total column ozone amounts 

for various regions around the globe have 

been derived from observations made by 

both satellite and ground-based instruments. 

The ozone trends reported here are derived 

primarily from the satellite data.

• Ozone decreases above the continental 

United States have essentially followed 

those occurring over the northern midlatitudes: a decrease to a minimum in about 

1993, followed by an increase since that 

time. The minimum total column ozone 

amounts over the continental United States 

reached in 1993 were about 5 to 8 percent 

below those present prior to 1980. The decreases roughly followed the increases in 

atmospheric abundance of ozone-depleting 

substances, which reached its maximum 

in roughly 1995. The ozone minimum occurred earlier than the maximum in ozonedepleting substances in large part due to 

the atmospheric changes brought about by 

the eruption of Mt. Pinatubo in 1991 (as 

discussed in Chapter 3). Column ozone 

increases since 1993 have brought the 

ozone deficit back to about 2 to 5 percent 

below the pre-1980 amounts. Model calculations suggest that these midlatitude ozone 

changes may have a significant contribution from the mixing of lower stratospheric 

ozone-depleted air from the northern polar 

latitudes during the spring period.

• Ozone depletion over northern high latitudes, such as over northern Alaska, is 

strongly influenced by Arctic springtime 

total ozone values, which have been significantly lower than those observed in the 

1980s. However, these Arctic springtime 

ozone depletions are highly variable from 

year to year.

• There has been no significant ozone depletion at the lower latitudes of the tropics 

and subtropics around the globe. Hence, 

column ozone over the parts of the United 

States in these regions has been essentially 

unchanged from their 1980 values.

6.2.2 Changes in UV over 

the United States

Changes in UV levels over the United States 

have been obtained from ground-based 

and satellite-based measurements. Surface 

UV levels are strongly affected by clouds, 

atmospheric fine particles, and air pollution, 

in addition to stratospheric ozone depletion, 

making it difficult to attribute changes in UV 

to long-term changes in stratospheric ozone 

alone. This difficulty is particularly acute 

since stratospheric ozone depletion during the 

past three decades has been rather small (less 

than 10 percent), with the exception of the 

high latitudes. In a world without the Montreal 

Protocol, stratospheric ozone changes would 

have been much larger than have actually 

occurred, and the associated UV increases 

would have been so large as to stand out from 

other variability and be easily measured over 

a wide range of latitudes. In addition, groundbased records are of limited use in relating UV 

increases to ozone decreases that occurred 

during the 1980s and 1990s since many stations 

did not initiate measurements until the late 

1990s, when ozone had already reached its 

minimum. A reliable way to derive the current 

changes in UV associated with ozone depletion 

is to use satellite measurements of atmospheric 

backscattered UV and the small amount 

reflected from the surface.

• Direct surface-based observations do not 

show significant UV trends for the United 

States over the past three decades because 

effects of clouds and atmospheric fine 

particles have likely masked the increase 

in UV due to ozone depletion.

• Estimates of UV based on satellite measurements of column ozone and reflectivity 

of the surface suggest that the clear-sky 

erythemal irradiance (a weighted combination of UVA and UVB wavelength 

ranges based on skin sensitivity) over the 
The total contribution 

of human-produced 

ozone-depleting 

substances and 

substitutes to direct 

radiative forcing is 

about 15% as large 

as the contribution 

from other 

greenhouse gases.

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Layer Recovery, and Implications for Ultraviolet Radiation Exposure

continental United States increased from 

1979 to the mid-1990s by about 7 percent. 

Currently, this irradiance is about 4 percent 

higher than it was at the start of the record 

in 1979. Year-to-year seasonal variations 

ranged from only a few percent to about 

20 percent. 

• Barrow, Alaska, has experienced UVB 

increases in March and April related to 

springtime ozone depletion. While these 

increases are larger than those observed 

at midlatitudes in the mid-1990s, they 

are roughly ten times smaller than those 

observed at the southern high latitudes due 

to the Antarctic ozone hole. 

6.2.3 Changes in Radiative Forcing

Globally averaged radiative forcing is a good 

metric for the relative contributions to climate 

change. Positive (negative) values for radiative 

forcing lead to warming (cooling). It is a 

reasonably good assumption that the globalaverage impacts from long-lived greenhouse 

gases scale with the magnitude of the globally 

averaged forcing. Many ODSs are themselves 

greenhouse gases, and hence ODSs contribute 

to radiative forcing.

The combined direct radiative forcing from ODSs 

and substitutes, including hydrofluorocarbons 

(HFCs), is still increasing, but at a slower 

rate than in the 1980s because the use of 

many ODSs (particularly chlorofluorocarbons 

[CFCs]) has been curtailed by the Montreal 

Protocol. This continued increasing trend 

in radiative forcing arises from continued 

increases in the atmospheric mixing ratios of 

hydrochlorofluorocarbons (HCFCs) and HFCs, 

which are being used as substitutes for CFCs in 

various applications.

• The total contribution of human-produced 

ODSs and substitutes to direct radiative 

forcing is approximately 0.34 W per m2

(representing the change between preindustrial times, ca. 1750, and the present), 

which is about 15 percent as large as the 

contribution from other greenhouse gases 

(1.66 W per m2 from carbon dioxide [CO2] 

plus 0.6 W per m2 from methane [CH4] and 

nitrous oxide [N2O] together). The bulk 

of the direct forcing from halocarbons in 

2005 was from CFCs (about 80 percent); 

other contributors include 10 percent from 

HCFCs, 7 percent from other ODSs, and 

3 percent from HFCs. Projections of these 

contributions to 2100 under the SRES 

A1B emission scenario can be found in 

Chapter 5.

Changes in atmospheric ozone abundances 

contribute to climate change by modifying 

atmospheric radiative properties and atmospheric 

temperatures. Changes in stratospheric ozone 

are often considered to be indirect climate 

forcings of ODSs, though other processes have 

also influenced changes in stratospheric ozone 

over time.

• Depletion of stratospheric ozone since 

about 1980 is estimated to have caused 

a negative radiative forcing (cooling) on 

climate (of approximately −0.05 W per 

m2

), corresponding in absolute magnitude 

to about 15 percent of the total direct 

positive forcing by ODSs alone. However, 

the uncertainties on this forcing are large 

enough (±0.1 W per m2, i.e., a range of 

−0.15 to +0.05 W per m2) to suggest that it 

could even be a positive radiative forcing 

(warming). Twentieth century increases in 

tropospheric ozone from pollution chemistry have caused a positive radiative forcing 

(of approximately +0.35 W per m2). ODS 

emissions and abundances have very little 

influence over ozone abundances in the 

lower atmosphere.

6.2.4 Future Ozone and UV 

Changes over the United States

As stated earlier, changes in ozone over the 

United States should be similar to the changes 

occurring over similar latitudes around the 

globe. Ozone-depleting substances addressed 

by the Montreal Protocol and its amendments 

should have a declining effect on stratospheric 

ozone between now and 2050, and a small effect 

on stratospheric ozone beyond 2050, assuming 

compliance with the Montreal Protocol and if 

all other factors remain roughly the same. In 

order to predict the future trend of ozone in that 

time frame, one must consider projections for 

climate changes and changes in trace gases such 

as other halogens, CH4, and N2O (in addition to 

any changes in solar UV irradiance). 
By the end of this 

century, direct 

forcing from CFCs is 

expected to decrease 

to roughly 40% of 

its current value.

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• Based on the prescribed surface concentrations of halocarbons used in the WMO 

(2007) baseline scenario (the scenario that 

was consistent with the Montreal Protocol provisions as of 2006), atmospheric 

halogen loading is estimated to recover to 

its 1980 value between 2040 and 2050 for 

midlatitudes, and between 2060 and 2070 

for the polar regions.

• Between now and 2020, the simulated total 

ozone content between 60°N and 60°S will 

increase in response to this decrease in 

halogen loading.

• Three-dimensional climate chemistry 

models (3-D CCMs) predict that stratospheric cooling and changes in circulation 

associated with greenhouse gas emissions 

may enable global ozone to return to its 

1980 value up to 15 years earlier than the 

expected halogen recovery date. Based on 

the assumed scenario for the greenhouse 

gases (which include CH4 and N2O), the 

ozone content is expected to be 2 to 5 percent above the 1980 values by 2100.

• Because of large interannual variability, 

the dates of the minimum in Arctic ozone 

from different models occur between 1997 

and 2015. Most CCMs show Arctic ozone 

values at 2050 larger than the 1980 values, 

with recovery between 2020 and 2040. 

Results from the majority of the models 

indicate that future Arctic ozone depletion 

will not be significantly worse than what 

has already occurred.

The future trend in surface UV is likely to be 

controlled more by changes in cloud cover, 

atmospheric fine particles, and tropospheric 

air quality than by changes in stratospheric 

ozone. Nevertheless, Equivalent Effective 

Stratospheric Chlorine (EESC) will still be a 

useful predictor for the relative effects of ODSs 

on future UV in terms of evaluating different 

scenarios.

6.2.5 Future Changes in 

Radiative Forcing

The radiative forcing from CO2 increases to +5 

W per m2 by 2100 for the IPCC Special Report 

on Emission Scenarios (SRES) A1B scenario 

(Nakićenović, 2000), a scenario involving rapid 

economic growth and balanced energy sources. 

Forcing from halocarbons and their substitutes 

will decline in the future, assuming continued 

compliance with the Montreal Protocol, and is 

summarized below.

• Direct forcing from CFCs, which constitute 

a significant fraction of total ODS forcing in today’s atmosphere, is expected to 

decrease from the current value of about 

+0.26 W per m2

 to about +0.1 W per m2

by 2100. Direct forcing from HCFCs and 

other ODSs is expected to be negligible 

by 2100.

• The indirect forcing of ozone depletion 

is expected to approach zero when EESC 

returns to its 1980 levels, while the direct 

forcing due mainly to CFCs remaining in 

the atmosphere and the CFC substitutes 

that do not contain either chlorine or bromine (e.g., HFCs) will continue.

• In the SRES A1B scenario, forcing from 

HFCs, which do not deplete stratospheric 

ozone, is predicted to increase to +0.15 W 

per m2 and +0.24 W per m2 by 2050 and 

2100, respectively, while other scenarios 

result in smaller forcings from these chemicals. However, current observations suggest that the present atmospheric radiative 

forcing of the HFCs has been larger than 

computed for the SRES scenarios, primarily due to higher HFC-23 concentrations. 

Therefore, additional uncertainty perhaps 

should be attached to the SRES HFC projections.

• Changes in ozone due to changes in other 

trace gases (CH4 and N2O) and to changes 

in climate will also contribute to future 

forcing. For example, increases in atmospheric circulation due to climate change 

could increase the flux of ozone from the 

stratosphere to the troposphere, resulting in 

an additional positive radiative forcing.

6.3 ACCOUNTABILITY

As stated earlier, the amount of stratospheric 

ozone depletion at any location is, in large part, 

a result of long-lived ODSs emitted from all over 

the globe. In addition, ozone depletion in a given 

location is not simply and linearly related to 

ODS amounts in the atmosphere. Accordingly, 

there is no simple relationship between a 

country’s contribution to global ODS emissions 

with the local ozone depletion occurring in 
From 1986 to 

1994, the United 

States accounted 

for about 25% of 

total production 

and consumption 

of ozone-depleting 

substances. From 

2001 to 2005, 

this fraction had 

declined to 10%.

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that region. To extend this association to local 

changes in UV radiation at the ground is further 

complicated by the dependence of UV on many 

local factors. Acknowledging this complexity, 

we can estimate the United States’ contribution 

to the global emissions of ODSs to derive the 

United States’ contribution to the atmospheric 

abundances of ODSs. We can then obtain a first 

approximation of the United States’ contribution 

to ozone depletion at any location, and thereby 

estimate the United States’ contribution to that 

portion of UV increase at that same location. In 

a similar manner, we can estimate the United 

States’ contribution to changes in radiative 

forcing due to the emissions of ODSs.

This chapter uses several metrics to estimate 

the United States’ ODS contributions. Ozone 

Depletion Potential (ODP) and Global Warming 

Potential (GWP) weighting is used to sum 

annual emissions to represent the United 

States’ contribution in that particular year to 

future ozone depletion and radiative forcing. 

The same could be used for ODS “banks” since 

they represent potential future emissions. The 

atmospheric burden of a specific halocarbon 

calculated from historical United States 

emissions, when compared with the observed 

total burden, provides a measure of the United 

States’ contribution to that halocarbon. The 

individual abundances can be combined 

using the formalism of Equivalent Effective 

Chlorine (EECl) and EESC to be used as a 

measure of ozone depletion contribution. 

Finally, radiative forcing calculated from the 

individual halocarbon abundances provides a 

measure of the United States’ contribution to 

climate forcing.

6.3.1 Contribution of the United States 

to the Global Abundance of ODSs

It is difficult to accurately quantify the United 

States’ contribution to the current atmospheric 

loading of ozone-depleting substances because 

of uncertainties associated with United States 

emission data prior to 1985. However, estimates 

of the United States’ contributions to global 

consumption and emissions of ODSs for recent 

periods can be derived, respectively, from 

information compiled by the United Nations 

Environment Programme (UNEP) or from 

estimates made by the U.S. Environmental 

Protection Agency (EPA). It should be noted 

that consumption of ODSs can have different 

emission patterns (in location and time) 

depending on the particular end use of the 

ODS. The U.S. EPA vintaging model calculates 

ODS emissions based on a variety of factors 

associated with the use or product application 

of the ODS. 

Production / Consumption

• Global production and consumption of 

ODSs have declined substantially since the 

late 1980s in response to the Montreal Protocol, its amendments and adjustments, and 

United States policy decisions. By 2005, 

annual global ODP-weighted production 

and consumption had declined 95 percent 

from the peak values of the late 1980s. By 

2005, annual ODP-weighted production 

and consumption in the United States had 

declined by 97 to 98 percent based on 

UNEP data.

• During 1986 to 1994 the United States 

accounted for 25(±2) percent of the total 

annual global production and consumption 

of ODSs reported by UNEP when weighted 

by ODPs. From 2001 to 2005, this fraction 

has been 10(±2) percent. This decline has 

been maintained despite a slower decline 

in U.S. CH3Br consumption reported to 

UNEP than in other nations owing to 

increased U.S. Critical Use Exemptions, 

CUEs (critical uses that are exempted 

from the Montreal Protocol). Increased 

CUEs caused the annual U.S. contribution 

to global CH3Br consumption to increase 

from 23(±4) percent between 2000 to 2003, 

to 36(±1) percent during 2004 to 2005.

Burdens and EECl

• Taking into account the uncertainties in 

United States emissions estimates for past 

years, atmospheric chlorine from United 

States emissions accounted for 17 to 42 

percent of global chlorine from regulated 

ODSs and substitute chemicals in 2005. 

Atmospheric chlorine from United States 

and global emissions has declined since 

the mid-1990s.

• The U.S. EPA vintaging model suggests 

that in 2005 the United States accounted 

for approximately 17 to 35 percent of the 

global atmospheric bromine burden arising 

from industrially produced CH3Br and ha-
United States 

banks account for 

approximately 28 

percent of total 

global banks of 

ozone-depleting 

substances.

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Chapter 6

lons, similar to that calculated for the peak 

year, 1998. Changes in total tropospheric 

bromine from the United States emissions 

of ODSs regulated by the Montreal Protocol mimicked global trends until 2002. 

Further, the vintaging model suggests that 

bromine emissions from the United States 

began increasing in 2002, due primarily to 

increased emissions of CH3Br.

• The decrease in tropospheric EECl since 

1994 has been about 20 percent of what is 

needed to return EECl to 1980 values (i.e., 

before substantial ozone depletion was 

observed). Atmospheric EECl calculated 

from United States emissions declined between 1994 and 2004; however, it declined 

much more slowly from 2004 to 2005. The 

United States accounted for 15 to 36 percent of EECl from industrially produced 

chemicals measured in the troposphere 

in 2005.

Banks and Future Emissions

United States emissions of some ODSs in the 

future, like those from other developed nations, 

will be determined to a large extent by the size 

of the U.S. ODS “banks,” i.e., those ODSs that 

are already produced but not yet released to 

the atmosphere due to old devices, structures, 

and stockpiles that exist in the United States. 

The magnitude of halocarbon banks has been 

derived using a new bottom-up approach. This 

new method leads to larger CFC banks and 

yields potential future CFC emissions that are 

higher than previously estimated (WMO, 2003). 

The U.S. EPA has divided total banks into 

“accessible” and “non-accessible” categories, 

with accessible banks consisting of ODSs in 

refrigeration and air conditioning equipment 

and fire fighting equipment.

• If released in a single year, the ODS banks 

in the United States in 2005 would have 

been equivalent (in terms of their contribution to stratospheric ozone depletion) to 7 

to 16 times the actual United States emissions of ODSs during that year.

• The U.S. EPA estimates that United 

States banks account for approximately 

28 percent of global banks, whether they 

are accessible or not, of all ODSs (ODPweighted). CFCs accounted for the largest 

fraction of the 2005 banks in the United 

States as well as throughout the globe.

• Approximately one-quarter of United 

States banks in 2005 were classified as 

being accessible (consisting of 210 ODPkilotons, Kt) and these accessible banks 

were comprised predominantly of halons 

(roughly 140 ODP-Kt), CFCs (~38 ODPKt), and HCFCs (~31 ODP-Kt). CFCs 

accounted for 18 percent of the accessible 

banks of ODSs as defined currently by the 

U.S. EPA.

• Banks play an important role in current 

HCFC emission rates. Future emissions 

of HCFCs will also be determined by 

the magnitude of any additional HCFC 

production.

6.3.2 Contribution of the United States 

to Climate Change via Emission of 

Ozone-Depleting Substances and 

the Resulting Ozone Changes

The increased abundances of ODSs, as well 

as the associated depletion of stratospheric 

ozone, contribute to the radiative forcing of 

climate. Since activities in the United States 

have caused significant emissions of ozonedepleting substances and other greenhouse 

gases, the changes in climate attributable to 

human activities are, in part, attributable to the 

United States.

• Globally, the direct radiative forcing from 

ODSs and substitutes was approximately 

+0.34 W per m2

, roughly 20 percent of that 

from CO2 in 2005. When indirect forcing 

associated with stratospheric ozone depletion is included, the net forcing from ODSs 

and substitutes is between +0.18 and +0.38 

W per m2. These values were estimated 

using the 100-year direct GWPs.

• The United States’ contribution to this direct forcing amounted to between +0.068 

and +0.16 W per m2, or between roughly 20 

and 50 percent of the global direct radiative 

forcing from these chemicals. This contribution has been fairly constant over the past 

decade. When net GWPs (i.e., that includes 

direct and indirect forcings) are considered, 

the range for the U.S. contribution is +0.04 

to +0.18 W per m2.

• Considering ODSs acting as climate forcing agents and using a 100-year direct 

GWP weighting, the U.S. EPA estimates 
Global ozone will 

return to its pre-1980 

levels around 2050. 

The Antarctic ozone 

hole is expected to 

disappear around 

2060 to 2070.

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Layer Recovery, and Implications for Ultraviolet Radiation Exposure

that United States banks in 2005 account 

for about 32 percent of global banks. The 

range for net GWP weighting is (31-60) 

percent. These banks, if released to the 

atmosphere, would represent the equivalent of 6.2 gigatons (Gt) CO2 emissions. 

Approximately one-quarter of United 

States banks in 2005 were classified as 

being accessible and they were comprised 

of HCFCs (0.9 to 1.1 Gt CO2-equivalents), 

HFCs (0.4 Gt CO2-equivalents), and CFCs 

(0.2-0.4 Gt CO2-equivalents). 

6.4 POTENTIAL MANAGEMENT 

OPTIONS

To provide quantitative information for 

assessing the societal benefit of potential 

future regulatory action, the future levels of 

ozone-depleting substances can be estimated 

for a variety of scenarios based on the findings 

noted above. These include scenarios to assess 

the influence of currently unrestricted uses, 

such as methyl bromide in quarantine and preshipment (QPS) applications, and unrestricted 

emissions from most banks and stockpiles. 

Equivalent Effective Stratospheric Chlorine 

(EESC) is a useful index for comparing the 

merits of different emissions scenarios. While 

changes in EESC do not relate in a simple way 

to stratospheric ozone levels that vary with 

location and time (due to the non-linearities that 

were mentioned earlier), it is clear that EESC 

changes are representative of the relative ozone 

depletion impacts. Based on projected EESC 

values and our understanding of atmospheric 

chemical and dynamical processes, we conclude 

the following:

• Amounts of atmospheric ozone-depleting 

substances, measured in terms of EESC, 

will be comparable to pre-1980 levels 

around 2050 if future emissions evolve in 

a manner consistent with current Montreal 

Protocol regulations. It is anticipated that, 

given the proven connection between 

ozone-depleting substances and stratospheric ozone loss, global ozone will also 

return to the pre-1980 levels roughly around 

the same time, assuming no other climate 

or atmospheric composition changes. 

However, as stated earlier, factors such as 

climate change and changes in other trace 

gases are predicted to accelerate global 

ozone recovery to pre-1980 values. 

• The ozone-depleting substances (ODSs), 

measured in terms of EESC, in the Antarctic ozone-hole region will return to 

pre-1980 levels around 2060 to 2070. Thus, 

the Antarctic ozone hole will essentially 

disappear around this date assuming full 

compliance with the Montreal Protocol 

and its amendments and barring major 

influences by climate change and other 

factors.

The date at which the atmospheric abundances 

of ODSs return to their 1980 levels is 2049 for 

a baseline scenario (a scenario that is consistent 

with Montreal Protocol provisions as of 2006). 

This return can be accelerated under several 

scenarios.

• The hypothetical cessation of all future 

emissions of ozone-depleting substances (such as hydrochlorofluorocarbons 

(HCFCs), and chlorofluorocarbons (CFCs) 

from banks) starting in 2007 would hasten 

the decline of ozone-depleting substances 

to their 1980 level by roughly 15 years (to 

2034).

• Under the scenario where no future production is assumed but emissions still arise 

from ODS banks, the EESC recovery date 

(i.e., to the 1980 level) is moved up by 

roughly six years (to 2043).

• Under the scenario where all ODS banks 

were recovered and destroyed in 2007, but 

future production is allowed to continue 

as in the baseline case, the EESC recovery 

date is moved up by eight years (to 2041).

• The significance of various United States 

ODS banks has been evaluated in terms of 

their effect on integrated EESC and compared with the significance of the global 

banks. Of accessible ODS banks in the 

U.S., halon-1301 and HCFC-22 represent 

the greatest potential contribution to integrated EESC above 1980 levels. If total 

United States banks are considered, not just 

accessible ones, CFC-11 banks are the most 

significant potential contributors to integrated EESC. Banks deemed inaccessible 

may still be recovered with appropriate 

policy measures, market-based incentives 

and/or certain technological advances.
The coming decades 

will be a period of 

changing atmospheric 

levels of ozonedepleting substances 

superimposed on 

changing climate, 

climate variability, 

and other factors.

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There are some uses of methyl bromide for 

which production and consumption are not 

limited or restricted under the currently 

amended Montreal Protocol.

• Global consumption of methyl bromide 

(CH3Br) for all fumigation-related uses 

declined by a factor of two from 1997 to 

2005 despite substantial consumption in 

applications not regulated by the Montreal 

Protocol. Nearly half (43 percent) of the 

global, industrially derived emissions of 

CH3Br during 2005 arose from QPS consumption not regulated by the Montreal 

Protocol. 

• United States consumption of CH3Br for all 

fumigation uses declined 40 percent from 

1997 to 2005 despite enhanced Critical Use 

Exemptions (CUEs) and QPS consumption 

since 2001. Enhanced CUEs caused the 

annual United States contribution to global 

CH3Br consumption for regulated uses to 

increase from 23(±4) percent during 2000 

to 2003 to 36(±1) percent during 2004 to 

2005. In the United States during 2001 to 

2006, consumption of CH3Br for fumigation not regulated by the protocol (QPS 

use) was, on average, 57(±20) percent of 

the amounts used and reported to UNEP 

in restricted applications and had increased 

by 13 percent per year, on average, from 

2001 to 2005.

The expected increase in stratospheric ozone 

over the coming decades will decrease surface 

UV. However, the future UV trend at the surface 

is likely to be more dominated by changes 

in cloud cover, atmospheric fine particle 

abundances, and tropospheric air quality than 

by changes in ODS abundances projected in 

accordance with the provisions of the Montreal 

Protocol.

Little further reduction in radiative forcing 

from ODSs can be achieved by 2100 beyond 

that predicted under the current provisions of 

the Montreal Protocol. Emissions reductions, 

however, could lower radiative forcing in the 

coming decades. Reductions in HFC emissions 

could also have a modest effect in this area.

• Action could be taken to limit the release 

of CFCs and HCFCs from banks and thus 

reduce their future emissions beyond what 

the current Montreal Protocol is expected 

to accomplish. If the entire estimated 

global CFC and HCFC banks had been 

recovered and destroyed in 2007, the direct 

radiative forcing is expected to be reduced 

by about 0.015 W per m2

 and 0.07 W per 

m2, respectively, in 2040, compared with 

the radiative forcing calculated assuming future emissions consistent with the 

Montreal Protocol regulations. However, 

a complete assessment of any benefits of 

such action would also need to include consideration of indirect influences associated 

with ozone depletion changes. 

6.4.1 The World Avoided

The dramatic decrease in emissions of ODSs 

since the late 1980s, called for by the Montreal 

Protocol, has been achieved in the United 

States through a combination of regulations 

restricting the use, handling, and labeling 

of specific compounds, a robust program to 

evaluate alternative compounds, voluntary 

industry initiatives, and outreach and education 

programs that have raised awareness of the 

threats caused by ODSs and the need to 

control them. Various emissions scenarios 

have been used to compare the ozone and UV 

levels of today with what might have occurred 

in the absence of the Montreal Protocol and 

the associated actions, as a way to assess the 

effectiveness and value of the Protocol to the 

United States and the world.

• Without the implementation of the Montreal Protocol, EESC levels around 2010 

likely would have been more than 50 

percent larger than currently expected. 

The abundances during the remainder of 

the twenty-first century would have depended on any subsequent policy actions 

taken. These increases in ODSs would 

have caused a corresponding substantially 

greater global ozone depletion. The Antarctic ozone hole would have persisted longer 

and may have been even larger than that 

currently observed.
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• The contributions of the United States to 

the ozone depletion via emission of ODSs 

to date have been significant. However, 

the United States has also contributed 

significantly to achieving the expected 

recovery of the ozone layer and associated 

surface UV changes, and reductions in 

direct climate forcing caused by ODSs by 

the phase-out of CFCs.

• The decline since the late 1980s of the United States’ emissions of ODSs, considered 

on a CO2-equivalent basis, corresponds to 

a climate benefit whose magnitude is large 

compared with the Kyoto Protocol’s targets 

during its first commitment period. This 

benefit includes an offset of the −0.05 W 

per m2 estimated in the Fourth Assessment 

Report of Working Group I of the Intergovernmental Panel on Climate Change 

(IPCC, 2007) due to the destruction of 

stratospheric ozone by ODSs (as discussed 

in Chapter 4). 

The coming decades will be a period of 

changing atmospheric ODS levels superimposed 

on changing climate, climate variability, and 

other factors. Box 6.1 outlines the key gaps in 

scientific understanding that can be identified 

at this time and that could help inform future 

decisions regarding the continued recovery 

of the ozone layer back to a state that is not 

affected by ozone-depleting substances. 

In this document we have synthesized and assessed what is known about the depletion of the stratospheric 

ozone layer by ozone-depleting substances, the associated changes in surface UV radiation, and our 

expectations for the recovery of the ozone layer to pre-1980 values. We have described aspects of the 

interrelationship between stratospheric ozone depletion and climate change, such as the contribution 

of ozone-depleting substances to climate forcing, the impact of climate change on stratospheric ozone, 

and the effects of ozone depletion on climate. We have also outlined the importance of understanding 

the ozone-climate interrelationship, including variability of climate, in making accurate projections of 

future ozone as it recovers to pre-1980 values.

Evolving societal and decision-making imperatives arising from the continued global commitment to 

shepherd the ozone layer back to “good health” will drive future research on the stratospheric ozone 

issue. For example, the Parties to the Montreal Protocol recently made adjustments to the phase-out 

schedule to phase-out hydrochloro-fluorocarbons (HCFCs) earlier than scheduled; this adjustment 

agreement takes effect in mid-2008 and will be implemented over the coming few years as scheduled 

in the agreement. Questions still remain about topics such as the control of chlorofluorocarbon (CFC) 

bank emissions and the use of methyl bromide for exempted and unregulated purposes. Accurate 

predictions of the consequences of near-term decision options will require the United States and 

international scientific communities to acquire new observational data, to develop an improved 

understanding of the physical and chemical processes involved in ozone depletion and ozone-climate 

interactions, and to incorporate this understanding in global models used to project the future state 

of the ozone layer. Further, it will also require some reporting and documentation on production and 

use of ODSs and their substitutes.

At present, the scientific and regulatory communities are in the “accountability” phase of the ozone layer 

issue, because science-based regulation to protect ozone has been in place for nearly two decades. 

Decision makers are increasingly interested in having answers to the bottom-line questions: Are our 

actions having the desired and expected effect? Is the ozone layer recovering? Are there other actions that 

would hasten ozone layer recovery? As outlined in this document, scientists have addressed and/or partially 

answered many of these questions. However, gaps remain in our knowledge and information base and 

in our ability to answer these questions with sufficient clarity and accuracy for policy decisions. 

BOX 6.1: Gaps in Our Understanding and Continued Information Needs
The U.S. Climate Change Science Program

164

Chapter 6

It has now become clear that it is critical to understand the linkages between stratospheric ozone depletion and 

climate change, because climate variations and change will alter the ozone recovery path and even the ozone 

abundance and distribution after ozone-depleting substances have returned to natural levels. This is owing to the 

fact that as the atmosphere moves toward a pre-1980 ODS abundance, other atmospheric conditions will not 

revert similarly toward their pre-1980 state. Understanding the implications of these different evolutions will be a 

key focus area of atmospheric ozone research. 

The climate protection afforded by the Montreal Protocol regulation of ODSs has been significant over the last 

two decades. This protection derives from the fact that the principal ODSs are effective greenhouse gases. The 

findings of the IPCC Fourth Assessment Report, released in 2007, enhanced the global focus on climate protection 

and also increased interest in questions that lie at the nexus of these two global environmental issues. Decision 

support information demands an evaluation of the implications for ozone and climate protection for scenarios of 

future regulation under the Montreal Protocol. Decision makers need to know in detail how the ozone layer and 

the climate system are connected and what aspects of this linkage are likely to be most important in this evolving 

Earth system. 

Based on our synthesis and assessment of the current state of knowledge and the above set of broad research 

imperatives, we have identified some key knowledge gaps. For simplicity they are listed in four different categories and 

are equally important. We believe that this description of the knowledge gaps will aid United States and international 

agencies in establishing research priorities and directions and in establishing reporting requirements.

Atmospheric Observations

Ozone observations: Precise and accurate ozone observations in the troposphere and stratosphere anchor our 

understanding of the present and future ozone layer. Furthermore, ozone observations must be geographically 

comprehensive and have extended duration. The recovery of the ozone layer is likely to manifest itself differently at 

different altitudes and regions in the stratosphere. Changes in the ozone abundance profiles in turn impact climate 

change. Therefore, more precise continued, uninterrupted, observations of temporal changes in distributions of 

column ozone as well as local ozone abundances over altitude and latitude are essential to identifying the path to 

recovery and to better predicting the future state of the ozone layer.

Observations and derivation of surface UV and associated factors: Predicting the surface UV changes, 

especially those due to stratospheric ozone changes, requires not only the measurements of the UV radiation 

but also of many associated factors. These include ozone (noted above), atmospheric fine particle abundances and 

properties, surface albedo, and transmission and reflection of radiation by clouds. Continuity in time, accuracy in 

value, and global coverage are necessary. Currently, most of the conclusions about the surface UV trends come 

from calculations that use observed ozone distributions. Facilitation and enhancement of this ability will better fill 

the need on recording and predicting UV changes.

ODS observations and EESC: The accuracy of our current predictions for the recovery of the ozone layer depends 

directly on the accuracy of our predictions of the return of ozone-depleting substances and EESC to their pre-ozonedepletion values. We lack adequate knowledge for many of the factors that influence how these quantities return 

to pre-1980 values. For example, we may not have sufficiently accurate values for the atmospheric lifetimes of many 

ODSs, especially in a changing climate. Another concern includes the accurate quantification and the eventual release 

of ODS from banks and stockpiles, an emission that is likely large enough to delay the ozone recovery by many years. 

Also, uncertainties in bank emissions can hinder the identification of potential ODS emissions in violation of Montreal 

Protocol regulations. Verification of these bank emissions and emissions from other unregulated activities, such as 

methyl bromide from QPS use, will require more extensive atmospheric monitoring on global and regional scales. 

BOX 6.1: Gaps in Our Understanding and Continued Information Needs cont’d
165

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

In addition, we cannot precisely quantify how much and in what chemical form the short-lived chemicals (especially 

those containing halogens and many of natural origin) are transported to the stratosphere, transformed to reactive 

compounds, and contribute to EESC and ozone depletion. Detailed knowledge of these factors will become more 

critical as overall ODS emissions and abundances decrease in the future, thereby increasing the relative contributions 

of the short-lived substances and bank and fugitive emissions to ozone depletion. Accuracy of emission information 

needs to be established via verification of emissions on global and regional scales.

Process Understanding

There are many specific inputs required to account for the past and predict the future ozone levels as well as climate 

change. They include accurate rates of chemical and photochemical processes, timescales and rates of dynamical 

processes including variations, and identification and quantification of many microphysical processes involved in 

formation, persistence, and characteristics of polar stratospheric clouds and stratospheric sulfate aerosols. The 

rates of many of these known processes are not sufficiently accurate and there may be some unrecognized processes 

that are not quantified. Examples include the recently highlighted uncertainties in the photolysis rate of dichlorine 

peroxide (Cl2O2), a molecule that plays a critical role in polar ozone depletion, and uncertainties in the destruction 

rates and pathways of existing ODSs (especially HCFCs) and of not yet released, but planned, substitutes for ODSs. 

Therefore, a continuing effort to understand and determine rates and mechanisms of such processes is essential.

Global Models

We are in the early stages of developing climate models and Earth System models that include the known 

interconnecting processes that link climate and ozone. Projections from three-dimensional chemistry climate models 

(that did not include explicit land surface and ocean interactions) were used extensively for the first time only during 

the WMO/UNEP ozone assessment of 2007. Such models will be essential for future evaluations. These models are 

highly complex because they include all known important chemical, physical, and dynamical processes that influence 

ozone and other atmospheric constituents. The identification and parameterization of contributing processes and 

the completion and validation of these maturing climate models together represent important improvements in 

our ability to project future ozone abundances. The models have demonstrated skill in predicting observed ozone 

changes and attributing the cause of global ozone decreases to ODS emissions. However, additional improvements 

are needed due to the demand for more precise and accurate projections of future ozone abundances that include 

the relevant climate feedbacks. For example, climate benefits related to reductions in ODS emissions are influenced 

by our understanding of the climate forcing associated with changes in stratospheric ozone, but this influence 

currently has large uncertainties. The ozone-climate models need to be sufficiently accurate to identify regulatory 

options that would optimize the dual ozone-climate benefit. 

Reporting and Documentation

The accumulation of global emissions of ODSs and other greenhouse gases has led to ozone depletion and climate 

change. The United States’ contributions in both emission categories have been significant. The ability to quantify 

the United States’ contributions is limited by gaps in our knowledge of country-specific emissions. Detail is lacking 

for historical emissions for ODSs as well as for other greenhouse gases. Efforts to fill these historical gaps will 

improve the statements of attribution and benefit concerning potential future United States actions. Efforts to 

avoid similar gaps in the future will add credibility to and confidence in documenting United States accountability 

for ozone depletion and climate change and in providing guidance for United States national regulations or United 

States participation in new international policy discussions.

BOX 6.1: Gaps in Our Understanding and Continued Information Needs cont’d
The U.S. Climate Change Science Program

166

Chapter 6
APPENDIX A

167

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

Appendix A is reprinted here with permission from:

Fahey, D.W. (Lead Author), Twenty Questions and Answers About the Ozone Layer: 2006 Update, 

50 pp., World Meteorological Organization, Geneva, Switzerland, 2007. [Reprinted from 

Scientific Assessment of Ozone Depletion: 2006, Global Ozone Research and Monitoring 

Project—Report no. 50, 572 pp., World Meteorological Organization, Geneva, Switzerland, 

2007.]

Repaginated for this report

Twenty Questions and Answers About the Ozone Layer: 

2006 Update

Coordinating Lead Author: David W. Fahey, NOAA
168

TWENTY QUESTIONS AND ANSWERSABOUT THE OZONE LAYER: 2006 UPDATE

Contents

page

INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 169

I. OZONE IN OUR ATMOSPHERE

Q1. What is ozone and where is it in the atmosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 170

Q2. How is ozone formed in the atmosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 171

Q3. Why do we care about atmospheric ozone? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 172

Q4. Is total ozone uniform over the globe? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 173

Q5. How is ozone measured in the atmosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 174

II. THE OZONE DEPLETION PROCESS

Q6. What are the principal steps in stratospheric ozone depletion caused by

human activities? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 176

Q7. What emissions from human activities lead to ozone depletion? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 178

Q8. What are the reactive halogen gases that destroy stratospheric ozone? . . . . . . . . . . . . . . . . . . . . . . . . . . 181

Q9. What are the chlorine and bromine reactions that destroy stratospheric ozone? . . . . . . . . . . . . . . . . . . . 184

Q10. Why has an “ozone hole” appeared over Antarctica when ozone-depleting gases are present

throughout the stratosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 186

III. STRATOSPHERIC OZONE DEPLETION

Q11. How severe is the depletion of the Antarctic ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 189

Q12. Is there depletion of the Arctic ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 193

Q13. How large is the depletion of the global ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 195

Q14. Do changes in the Sun and volcanic eruptions affect the ozone layer? . . . . . . . . . . . . . . . . . . . . . . . . . . 197

IV. CONTROLLING OZONE-DEPLETING GASES

Q15. Are there regulations on the production of ozone-depleting gases? . . . . . . . . . . . . . . . . . . . . . . . . . . . . 199

Q16. Has the Montreal Protocol been successful in reducing ozone-depleting gases

in the atmosphere? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 201

V. IMPLICATIONS OF OZONE DEPLETION

Q17. Does depletion of the ozone layer increase ground-level ultraviolet radiation? . . . . . . . . . . . . . . . . . . . 204

Q18. Is depletion of the ozone layer the principal cause of climate change? . . . . . . . . . . . . . . . . . . . . . . . . . . 207

VI. STRATOSPHERIC OZONE IN THE FUTURE

Q19. How will recovery of the ozone layer be identified? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 210

Q20. When is the ozone layer expected to recover? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 212

ADDITIONAL TOPICS

• Global Ozone Dobson Network . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.8

• Understanding Stratospheric Ozone Depletion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.10

• Heavier-Than-Air CFCs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.13

• Replacing the Loss of Ozone in the Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.16

• The Discovery of the Antarctic Ozone Hole . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.21

• The Anomalous 2002 Antarctic Ozone Hole . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Q.25
TWENTY QUESTIONS: 2006 UPDATE

169

Ozone is a very small part of our atmosphere, but its

presence is nevertheless vital to human well-being.

Most ozone resides in the upper part of the atmosphere. This region, called the stratosphere, is more than

10 kilometers (6 miles) above Earth’s surface. There,

about 90% of atmospheric ozone is contained in the

“ozone layer,” which shields us from harmful ultraviolet

radiation from the Sun.

However, it was discovered in the mid-1970s that

some human-produced chemicals could destroy ozone and

deplete the ozone layer. The resulting increase in ultraviolet radiation at Earth’s surface may increase the incidences of skin cancer and eye cataracts.

Following the discovery of this environmental issue,

researchers focused on a better understanding of this threat

to the ozone layer. Monitoring stations showed that the

abundances of the ozone-depleting chemicals were

steadily increasing in the atmosphere. These trends were

linked to growing production and use of chemicals like

chlorofluorocarbons (CFCs) for refrigeration and air

conditioning, foam blowing, and industrial cleaning.

Measurements in the laboratory and the atmosphere characterized the chemical reactions that were involved in

ozone destruction. Computer models employing this

information could predict how much ozone depletion was

occurring and how much more could occur in the future.

Observations of the ozone layer showed that depletion was indeed occurring. The most severe and most surprising ozone loss was discovered to be recurring in

springtime over Antarctica. The loss in this region is commonly called the “ozone hole” because the ozone depletion is so large and localized. A thinning of the ozone

layer also has been observed over other regions of the

globe, such as the Arctic and northern middle latitudes.

The work of many scientists throughout the world

has provided a basis for building a broad and solid scientific understanding of the ozone depletion process. With

this understanding, we know that ozone depletion is occurring and why. And, most important, we know that if

ozone-depleting gases were to continue to accumulate in

the atmosphere, the result would be more depletion of the

ozone layer.

In response to the prospect of increasing ozone depletion, the governments of the world crafted the 1987 United

Nations Montreal Protocol as a global means to address

this global issue. As a result of the broad compliance with

the Protocol and its Amendments and Adjustments and,

of great significance, industry’s development of “ozonefriendly” substitutes for the now-controlled chemicals,

the total global accumulation of ozone-depleting gases

has slowed and begun to decrease. This has reduced the

risk of further ozone depletion. Now, with continued compliance, we expect recovery of the ozone layer by the late

21st century. The International Day for the Preservation

of the Ozone Layer, 16 September, is now celebrated on

the day the Montreal Protocol was agreed upon.

This is a story of notable achievements: discovery,

understanding, decisions, actions, and verification. It is a

story written by many: scientists, technologists, economists, legal experts, and policymakers. And, dialogue has

been a key ingredient.

To help foster continued interaction, this component

of the Scientific Assessment of Ozone Depletion: 2006

presents 20 questions and answers about the often-complex science of ozone depletion. The answers are updates

of those first presented in the previous ozone Assessment,

Scientific Assessment of Ozone Depletion: 2002. The

questions address the nature of atmospheric ozone, the

chemicals that cause ozone depletion, how global and

polar ozone depletion occur, and what could lie ahead for

the ozone layer. A brief answer to each question is first

given in italics; an expanded answer then follows. The

answers are based on the information presented in the

2006 and earlier Assessment reports. These reports and

the answers provided here were all prepared and reviewed

by a large international group of scientists. 1

INTRODUCTION

1 The update of this component of the Assessment was discussed by the 77 scientists who attended the Panel Review Meeting for the

2006 Ozone Assessment (Les Diablerets, Switzerland, 19-23 June 2006). In addition, subsequent contributions, reviews, or comments were provided by the following individuals: S.A. Montzka (special recognition), R.J. Salawitch (special recognition), D.L.

Albritton, S.O. Andersen, P.J. Aucamp, M.P. Baldwin, A.F. Bias, G. Bodeker, J.F. Bornman, G.O. Braathen, J.P. Burrows, M.-L.

Chanin, C. Clerbaux, M. Dameris, J.S. Daniel, S.B. Diaz, E.G. Dutton, C.A. Ennis, V. Eyring, V.E. Fioletov, N.P. Gillet, N.R.P.

Harris, M.K.W. Ko, L. Kuijpers, G.L. Manney, R.L. McKenzie, R. Müller, E.R. Nash, P.A. Newman, T. Peter, A.R. Ravishankara, A.

Robock, M.L. Santee, U. Schmidt, G. Seckmeyer, T.G. Shepherd, R.S. Stolarski, W.T. Sturges, J.C. van der Leun, G.J.M. Velders,

D.W. Waugh, C.S. Zerefos.
TWENTY QUESTIONS: 2006 UPDATE

170

Ozone is a gas that is naturally present in our atmosphere. Ozone has the chemical formula O3 because an

ozone molecule contains three oxygen atoms (see Figure

Q1-1). Ozone was discovered in laboratory experiments

in the mid-1800s. Ozone’s presence in the atmosphere

was later discovered using chemical and optical measurement methods. The word ozone is derived from the Greek

word óζειν (ozein in Latin), meaning “to smell.” Ozone

has a pungent odor that allows it to be detected even at

very low amounts. Ozone will rapidly react with many

chemical compounds and is explosive in concentrated

amounts. Electrical discharges are generally used to produce ozone for industrial processes such as air and water

purification and bleaching textiles and food products.

Ozone location. Most ozone (about 90%) is found

in the stratosphere, a region that begins about 10-16 kilometers (6-10 miles) above Earth’s surface and extends up

to about 50 kilometers (31 miles) altitude (see Figure Q1-

2). Most ozone resides in the lower stratosphere in what

is commonly known as the “ozone layer.” The remaining

ozone, about 10%, is found in the troposphere, which is

the lowest region of the atmosphere, between Earth’s surface and the stratosphere.

Ozone abundance. Ozone molecules have a relatively low abundance in the atmosphere. In the stratosphere near the peak of the ozone layer, there are up to

12,000 ozone molecules for every billion air molecules (1

billion = 1,000 million). Most air molecules are either

oxygen (O2) or nitrogen (N2) molecules. In the troposphere near Earth’s surface, ozone is even less abundant,

with a typical range of 20 to 100 ozone molecules for

each billion air molecules. The highest surface values

are a result of ozone formed in air polluted by human

activities.

As an illustration of the low relative abundance of

ozone in our atmosphere, one can imagine bringing all the

ozone molecules in the troposphere and stratosphere down

to Earth’s surface and uniformly distributing these molecules into a gas layer over the globe. The resulting layer

of pure ozone would have a thickness of less than onehalf centimeter (about one-quarter inch).

Figure Q1-2. Atmospheric ozone. Ozone is present

throughout the lower atmosphere (troposphere and

stratosphere). Most ozone resides in the stratospheric

“ozone layer” above Earthʼs surface. Increases in ozone

occur near the surface as a result of pollution from

human activities.

Figure Q1-1. Ozone and oxygen. A molecule of ozone

(O3) contains three oxygen (O) atoms bound together.

Oxygen molecules (O2), which constitute 21% of Earthʼs

atmosphere, contain two oxygen atoms bound together.

I. OZONE IN OUR ATMOSPHERE

Q1: What is ozone and where is it in the atmosphere?

Ozone is a gas that is naturally present in our atmosphere. Each ozone molecule contains three atoms of oxygen and is

denoted chemically as O3. Ozone is found primarily in two regions of the atmosphere. About 10% of atmospheric

ozone is in the troposphere, the region closest to Earth (from the surface to about 10-16 kilometers (6-10 miles)). The

remaining ozone (90%) resides in the stratosphere, primarily between the top of the troposphere and about 50 kilometers (31 miles) altitude. The large amount of ozone in the stratosphere is often referred to as the “ozone layer.”

Ozone and Oxygen

Ozone in the Atmosphere
TWENTY QUESTIONS: 2006 UPDATE

171

Stratospheric ozone. Stratospheric ozone is naturally formed in chemical reactions involving ultraviolet

sunlight and oxygen molecules, which make up 21% of

the atmosphere. In the first step, sunlight breaks apart

one oxygen molecule (O2) to produce two oxygen atoms

(2 O) (see Figure Q2-1). In the second step, each atom

combines with an oxygen molecule to produce an ozone

molecule (O3). These reactions occur continually wherever ultraviolet sunlight is present in the stratosphere. As

a result, the greatest ozone production occurs in the tropical stratosphere.

The production of stratospheric ozone is balanced by

its destruction in chemical reactions. Ozone reacts continually with a wide variety of natural and humanproduced chemicals in the stratosphere. In each reaction,

an ozone molecule is lost and other chemical compounds

are produced. Important reactive gases that destroy ozone

are those containing chlorine and bromine (see Q8).

Some stratospheric ozone is transported down into

the troposphere and can influence ozone amounts at

Earth’s surface, particularly in remote, unpolluted regions

of the globe.

Tropospheric ozone. Near Earth’s surface, ozone is

produced in chemical reactions involving naturally occurring gases and gases from pollution sources. Production

reactions primarily involve hydrocarbon and nitrogen

oxide gases and require sunlight. Fossil fuel combustion

is a primary pollution source for tropospheric ozone production. The surface production of ozone does not significantly contribute to the abundance of stratospheric ozone.

The amount of surface ozone is too small in comparison,

and the transport of surface air to the stratosphere is not

effective enough. As in the stratosphere, ozone in the troposphere is destroyed in naturally occurring chemical

reactions and in reactions involving human-produced

chemicals. Tropospheric ozone can also be destroyed

when ozone reacts with a variety of surfaces, such as those

of soils and plants.

Balance of chemical processes. Ozone abundances

in the stratosphere and troposphere are determined by the

balance between chemical processes that produce ozone

and processes that destroy ozone. The balance is determined by the amounts of reacting gases and how the rate

or effectiveness of the various reactions varies with sunlight intensity, location in the atmosphere, temperature,

and other factors. As atmospheric conditions change to

favor ozone-producing reactions in a certain location,

ozone abundances will increase. Similarly, if conditions

change to favor reactions that destroy ozone, abundances

will decrease. The balance of production and loss reactions combined with atmospheric air motions determines

the global distribution of ozone on time scales of days to

many months. Global ozone has decreased in the last

decades because the amounts of reactive gases containing

chlorine and bromine have increased in the stratosphere

(see Q13).

Figure Q2-1. Stratospheric ozone production. Ozone

is naturally produced in the stratosphere in a two-step

process. In the first step, ultraviolet sunlight breaks apart

an oxygen molecule to form two separate oxygen atoms.

In the second step, each atom then undergoes a binding

collision with another oxygen molecule to form an ozone

molecule. In the overall process, three oxygen molecules plus sunlight react to form two ozone molecules.

Q2: How is ozone formed in the atmosphere?

Ozone is formed throughout the atmosphere in multistep chemical processes that require sunlight. In the stratosphere, the process begins with an oxygen molecule (O2 ) being broken apart by ultraviolet radiation from

the Sun. In the lower atmosphere (troposphere), ozone is formed in a different set of chemical reactions

involving hydrocarbons and nitrogen-containing gases.

Stratospheric Ozone Production
TWENTY QUESTIONS: 2006 UPDATE

172

Natural ozone. In the absence of human activities

on Earth’s surface, ozone would still be present near the

surface and throughout the troposphere and stratosphere

because ozone is a natural component of the clean atmosphere. All ozone molecules are chemically identical, with

each containing three oxygen atoms. However, ozone in

the stratosphere (good ozone) has very different environmental consequences for humans and other life forms than

excess ozone in the troposphere near Earth’s surface (bad

ozone). Natural ozone in the troposphere is also considered “good” because it initiates the chemical removal of

many pollutants, such as carbon monoxide and nitrogen

oxides, as well as greenhouse gases such as methane.

Good ozone. Stratospheric ozone is considered good

for humans and other life forms because it absorbs ultraviolet (UV)-B radiation from the Sun (see Figure Q3-1).

If not absorbed, UV-B would reach Earth’s surface in

amounts that are harmful to a variety of life forms. In

humans, increased exposure to UV-B increases the risk of

skin cancer (see Q17), cataracts, and a suppressed immune

system. UV-B exposure before adulthood and cumulative

exposure are both important factors in the risk. Excessive

UV-B exposure also can damage terrestrial plant life,

single-cell organisms, and aquatic ecosystems. Other UV

radiation, UV-A, which is not absorbed significantly by

ozone, causes premature aging of the skin.

The absorption of UV-B radiation by ozone is a source

of heat in the stratosphere. This helps to maintain the

stratosphere as a stable region of the atmosphere, with

temperatures increasing with altitude. As a result, ozone

plays a key role in controlling the temperature structure

of Earth’s atmosphere.

Protecting good ozone. In the mid-1970s, it was discovered that halogen source gases released in human

activities could cause stratospheric ozone depletion (see

Q6). Ozone depletion increases harmful UV-B amounts

at Earth’s surface. Global efforts have been undertaken

to protect the ozone layer through regulation of ozonedepleting gases (see Q15 and Q16).

Bad ozone. Excess ozone formed near Earth’s surface in reactions caused by the presence of human-made

pollutant gases is considered bad ozone. Increased ozone

amounts are harmful to humans, plants, and other living

systems because ozone reacts strongly to destroy or alter

many other molecules. Excessive ozone exposure reduces

crop yields and forest growth. In humans, exposure to

high levels of ozone can reduce lung capacity; cause chest

pains, throat irritation, and coughing; and worsen preexisting health conditions related to the heart and lungs.

In addition, increases in tropospheric ozone lead to a

warming of Earth’s surface (see Q18). The negative

effects of increasing tropospheric ozone contrast sharply

with the positive effects of stratospheric ozone as an

absorber of harmful UV-B radiation from the Sun.

Reducing bad ozone. Reducing the emission of pollutants can reduce bad ozone in the air surrounding

humans, plants, and animals. Major sources of pollutants

include large cities where fossil fuel consumption and

industrial activities are greatest. Many programs around

the globe have already been successful in reducing the

emission of pollutants that cause excess ozone production

near Earth’s surface.

Figure Q3-1. UV-B protection by the ozone layer.

The ozone layer resides in the stratosphere and surrounds the entire Earth. UV-B radiation (280- to 315-

nanometer (nm) wavelength) from the Sun is partially

absorbed in this layer. As a result, the amount of UV-B

reaching Earthʼs surface is greatly reduced. UV-A (315-

to 400-nm wavelength) and other solar radiation are

not strongly absorbed by the ozone layer. Human exposure to UV-B increases the risk of skin cancer,

cataracts, and a suppressed immune system. UV-B

exposure can also damage terrestrial plant life, singlecell organisms, and aquatic ecosystems.

Q3: Why do we care about atmospheric ozone?

Ozone in the stratosphere absorbs some of the Sun’s biologically harmful ultraviolet radiation. Because of this

beneficial role, stratospheric ozone is considered “good” ozone. In contrast, excess ozone at Earth’s surface that

is formed from pollutants is considered “bad” ozone because it can be harmful to humans, plants, and animals.

The ozone that occurs naturally near the surface and in the lower atmosphere is also beneficial because ozone

helps remove pollutants from the atmosphere.

UV Protection by the Ozone Layer
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Total ozone. Total ozone at any location on the globe

is found by measuring all the ozone in the atmosphere

directly above that location. Total ozone includes that

present in the stratospheric ozone layer and that present

throughout the troposphere (see Figure Q1-2). The contribution from the troposphere is generally only about 10%

of total ozone. Total ozone values are often reported in

Dobson units, denoted “DU.” Typical values vary

between 200 and 500 DU over the globe (see Figure Q4-

1). A total ozone value of 500 DU, for example, is equivalent to a layer of pure ozone gas on Earth’s surface having

a thickness of only 0.5 centimeters (0.2 inches).

Global distribution. Total ozone varies strongly

with latitude over the globe, with the largest values occurring at middle and high latitudes (see Figure Q4-1). This

is a result of winds that circulate air in the stratosphere,

moving tropical air rich in ozone toward the poles. Since

about 1980, regions of low total ozone have occurred at

polar latitudes in winter and spring as a result of the chemical destruction of ozone by chlorine and bromine gases

(see Q11 and Q12). The smallest values of total ozone

(other than in the Antarctic in spring) occur in the tropics

in all seasons, in part because the troposphere extends to

a higher altitude in the tropics, and consequently, the

thickness of the ozone layer is smallest there.

Natural variations. The variations of total ozone

with latitude and longitude come about for two reasons.

First, natural air motions mix air between regions of the

stratosphere that have high ozone values and those that

have low ozone values. Air motions also increase the vertical thickness of the ozone layer near the poles, which

increases the value of total ozone in those regions.

Tropospheric weather systems can temporarily reduce the

thickness of the stratospheric ozone layer in a region, lowering total ozone at the same time. Second, variations

occur as a result of changes in the balance of chemical

production and loss processes as air moves to different

locations over the globe. Reductions in ultraviolet radiation from the sun in its 11-year cycle, for example, reduce

the production of ozone.

Scientists have a good understanding of how chemistry and air motions work together to cause the observed

large-scale features in total ozone, such as those seen in

Figure Q4-1. Ozone changes are carefully monitored by

a large group of investigators using satellite, airborne, and

ground-based instruments. The analysis of these observations helps scientists to estimate the contribution of

human activities to ozone depletion.

Q4: Is total ozone uniform over the globe?

No, the total amount of ozone above the surface of Earth varies with location on time scales that range from

daily to seasonal and longer. The variations are caused by stratospheric winds and the chemical production

and destruction of ozone. Total ozone is generally lowest at the equator and highest near the poles because

of the seasonal wind patterns in the stratosphere.

Figure Q4-1. Total ozone. A total ozone value is

obtained by measuring all the ozone that resides in the

atmosphere over a given location on Earthʼs surface.

Total ozone values shown here are reported in

“Dobson units” as measured by a satellite instrument

from space. Total ozone varies with latitude, longitude,

and season, with the largest values at high latitudes

and the lowest values in tropical regions. Total ozone

at most locations varies with time on a daily to seasonal basis as ozone-rich air is moved about the globe

by stratospheric winds. Low total ozone values over

Antarctica in the 22 December image represent the

remainder of the “ozone hole” from the 1999 Antarctic

winter/spring season (see Q11).

Global Satellite Maps of Total Ozone
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The abundance of ozone in the atmosphere is measured by a variety of techniques (see Figure Q5-1). The

techniques make use of ozone’s unique optical and chemical properties. There are two principal categories of

measurement techniques: local and remote. Ozone

measurements by these techniques have been essential in

monitoring changes in the ozone layer and in developing

our understanding of the processes that control ozone

abundances.

Local measurements. Local measurements of

atmospheric ozone abundance are those that require air to

be drawn directly into an instrument. Once inside an

instrument, ozone can be measured by its absorption of

ultraviolet (UV) light or by the electrical current produced

in an ozone chemical reaction. The latter approach is used

in the construction of “ozonesondes,” which are lightweight, ozone-measuring modules suitable for launching

on small balloons. The balloons ascend far enough in the

atmosphere to measure ozone in the stratospheric ozone

layer. Ozonesondes are launched regularly at many locations around the world. Local ozone-measuring instruments using optical or chemical detection schemes are

also used routinely on board research aircraft to measure

the distribution of ozone in the troposphere and lower

stratosphere. High-altitude research aircraft can reach the

ozone layer at most locations over the globe and can reach

farthest into the layer at high latitudes in polar regions.

Ozone measurements are also being made on some commercial aircraft.

Remote measurements. Remote measurements of

ozone abundance are obtained by detecting the presence

of ozone at large distances away from the instrument.

Most remote measurements of ozone rely on its unique

absorption of UV radiation. Sources of UV radiation that

can be used are the Sun and lasers. For example, satellites use the absorption of UV sunlight by the atmosphere

or the absorption of sunlight scattered from the surface of

Earth to measure ozone over nearly the entire globe on a

daily basis. A network of ground-based detectors measures ozone by the amount of the Sun’s UV light that

reaches Earth’s surface. Other instruments measure ozone

using its absorption of infrared or visible radiation or its

emission of microwave or infrared radiation. Total ozone

amounts and the altitude distribution of ozone can be

obtained with remote measurement techniques. Lasers

are routinely deployed at ground sites or on board aircraft

to detect ozone over a distance of many kilometers along

the laser light path.

Figure Q5-1. Ozone measurements. Ozone is measured throughout the atmosphere with instruments on the

ground and on board aircraft, high-altitude balloons, and

satellites. Some instruments measure ozone locally in

sampled air and others measure ozone remotely some

distance away from the instrument. Instruments use

optical techniques, with the Sun and lasers as light

sources, or use chemical reactions that are unique to

ozone. Measurements at many locations over the globe

are made regularly to monitor total ozone amounts.

Q5: How is ozone measured in the atmosphere?

The amount of ozone in the atmosphere is measured by instruments on the ground and carried aloft on

balloons, aircraft, and satellites. Some measurements involve drawing air into an instrument that contains a system for detecting ozone. Other measurements are based on ozone’s unique absorption of light

in the atmosphere. In that case, sunlight or laser light is carefully measured after passing through a portion of the atmosphere containing ozone.

Measuring Ozone in the Atmosphere
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175

Global Ozone Dobson Network

The first instrument for routine monitoring of total ozone was developed by Gordon M. B. Dobson in

the 1920s. The instrument, now called a Dobson spectrophotometer, measures the intensity of sunlight at

two ultraviolet wavelengths: one that is strongly absorbed by ozone and one that is weakly absorbed. The

difference in light intensity at the two wavelengths is used to provide a measurement of total ozone above

the location of the instrument.

A global network of ground-based, total-ozone observing stations was established in 1957 as part of

the International Geophysical Year. Today, there are about 100 sites distributed throughout the world (from

South Pole, Antarctica (90°S), to Ellesmere Island, Canada (83°N)), many of which routinely measure total

ozone with Dobson instruments. The accuracy of these observations is maintained by regular calibrations

and intercomparisons. Data from the network have been essential for understanding the effects of chlorofluorocarbons (CFCs) and other ozone-depleting gases on the global ozone layer, starting before the launch

of space-based ozone-measuring instruments and continuing to the present day. Because of their stability

and accuracy, the Dobson instruments are now routinely used to help calibrate space-based observations

of total ozone.

Pioneering scientists have traditionally been honored by having units of measure named after them.

Accordingly, the unit of measure for total ozone is called the “Dobson unit” (see Q4).
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176

6Emission, accumulation, and transport. The

principal steps in stratospheric ozone depletion caused by

human activities are shown in Figure Q6-1. The process

begins with the emission, at Earth’s surface, of source

gases containing the halogens chlorine and bromine (see

Q7). The halogen source gases include manufactured

chemicals released to the atmosphere by a variety of human

activities. Chlorofluorocarbons (CFCs) are an important

example of chlorine-containing gases. Emitted source

gases accumulate in the lower atmosphere (troposphere)

and are eventually transported to the stratosphere. The

accumulation occurs because most source gases are unreactive in the lower atmosphere. However, small amounts

of these gases dissolve or are taken up in ocean waters.

Some emissions of halogen gases come from natural

sources (see Q7). These emissions also accumulate in the

troposphere and are transported to the stratosphere.

Conversion, reaction, and removal. Halogen

source gases do not react directly with ozone. Once in the

stratosphere, halogen source gases are chemically converted to reactive halogen gases by ultraviolet radiation

from the Sun (see Q8). The rate of conversion is related

to the atmospheric lifetime of a gas (see Q7). Source

gases with lifetimes greater than a few years may circulate between the troposphere and stratosphere multiple

times before full conversion occurs.

The reactive gases formed in the eventual conversion

of the halogen source gases react chemically to destroy

ozone in the stratosphere (see Q9). The average depletion

II. THE OZONE DEPLETION PROCESS

Q6: What are the principal steps in stratospheric ozone depletion caused by human activities?

The initial step in the depletion of stratospheric ozone by human activities is the emission, at Earth’s surface,

of ozone-depleting gases containing chlorine and bromine. Most of these gases accumulate in the lower

atmosphere because they are unreactive and do not dissolve readily in rain or snow. Eventually, these emitted source gases are transported to the stratosphere, where they are converted to more reactive gases containing chlorine and bromine. These more reactive gases then participate in reactions that destroy ozone. Finally,

when air returns to the lower atmosphere, these reactive chlorine and bromine gases are removed from Earth’s

atmosphere by rain and snow.

Figure Q6-1. Principal steps in stratospheric ozone

depletion. The stratospheric ozone depletion process

begins with the emission of halogen source gases at

Earthʼs surface and ends when reactive halogen gases

are removed by rain and snow in the troposphere and

deposited on Earthʼs surface. In the stratosphere, the

reactive halogen gases, namely chlorine monoxide

(ClO) and bromine monoxide (BrO), destroy ozone.

Principal Steps in the Depletion

of Stratospheric Ozone
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177

of total ozone attributed to reactive gases is smallest in

the tropics and largest at high latitudes (see Q13). In

polar regions, the presence of polar stratospheric clouds

greatly increases the abundance of the most reactive halogen gases (see Q10). This results in substantial ozone

destruction in polar regions in winter and spring (see Q11

and Q12).

After a few years, air in the stratosphere returns to the

troposphere, bringing along reactive halogen gases.

These gases are then removed from the atmosphere by

rain and other precipitation and deposited on Earth’s surface. This removal brings to an end the destruction of

ozone by chlorine and bromine atoms that were first

released to the atmosphere as components of halogen

source gas molecules.

Tropospheric conversion. Halogen source gases

with short lifetimes (see Q7) undergo significant chemical conversion in the troposphere, producing reactive

halogen gases and other compounds. Source gas molecules that are not converted accumulate in the troposphere

and are transported to the stratosphere. Because of

removal by precipitation, only small portions of the reactive halogen gases produced in the troposphere are also

transported to the stratosphere. Important examples of

gases that undergo some tropospheric removal are the

HCFCs, which are used as substitute gases for other halogen source gases (see Q15 and Q16), bromoform, and

gases containing iodine (see Q7).

Understanding Stratospheric Ozone Depletion

Scientists learn about ozone destruction through a combination of laboratory studies, computer

models, and stratospheric observations. In laboratory studies scientists are able to discover and evaluate

individual chemical reactions that also occur in the stratosphere. Chemical reactions between two gases

follow well-defined physical rules. Some of these reactions occur on the surfaces of particles formed in the

stratosphere. Reactions have been studied that involve a wide variety of molecules containing chlorine,

bromine, fluorine, and iodine and other atmospheric constituents such as oxygen, nitrogen, and hydrogen.

These studies show that there exist several reactions involving chlorine and bromine that can directly or

indirectly cause ozone destruction in the atmosphere.

With computer models, scientists can examine the overall effect of a large group of known reactions

under the chemical and physical conditions found in the stratosphere. These models include winds, air

temperatures, and the daily and seasonal changes in sunlight. With such analyses, scientists have shown

that chlorine and bromine can react in catalytic cycles in which one chlorine or bromine atom can destroy

many ozone molecules. Scientists use model results to compare with past observations as a test of our

understanding of the atmosphere and to evaluate the importance of new reactions found in the laboratory.

Computer models also enable scientists to explore the future by changing atmospheric conditions and other

model parameters.

Scientists have conducted observations to find out which gases are present in various regions of the

stratosphere and at what concentrations. They have monitored the change in these abundances over time

periods spanning a daily cycle to decades. Observations have shown that halogen source gases and reactive halogen gases are present in the stratosphere at expected amounts. Ozone and chlorine monoxide

(ClO), for example, have been observed extensively with a variety of instruments. Instruments on the

ground and on board satellites, balloons, and aircraft detect ozone and ClO at a distance (remotely) using

optical and microwave signals. High-altitude aircraft and balloon instruments detect both gases locally in

the stratosphere (see Q5). For example, these observations show that ClO is present at elevated amounts

in the Antarctic and Arctic stratospheres in the late winter/early spring season, when the most severe ozone

depletion occurs (see Q8).
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178

Principal human-produced chlorine and bromine

gases. Human activities cause the emission of halogen

source gases that contain chlorine and bromine atoms.

These emissions into the atmosphere ultimately lead to

stratospheric ozone depletion. The source gases that contain only carbon, chlorine, and fluorine are called “chlorofluorocarbons,” usually abbreviated as CFCs. CFCs,

along with carbon tetrachloride (CCl4) and methyl chloroform (CH3CCl3), historically have been the most important chlorine-containing gases that are emitted by human

activities and destroy stratospheric ozone (see Figure Q7-

1). These and other chlorine-containing gases have been

used in many applications, including refrigeration, air conditioning, foam blowing, aerosol propellants, and cleaning

of metals and electronic components. These activities

have typically caused the emission of halogen-containing

gases to the atmosphere.

Another category of halogen source gases contains

bromine. The most important of these are the “halons”

and methyl bromide (CH3Br). Halons are halogenated

hydrocarbon gases originally developed to extinguish

fires. Halons are widely used to protect large computers,

military hardware, and commercial aircraft engines.

Because of these uses, halons are often directly released

into the atmosphere. Halon-1211 and halon-1301 are the

most abundant halons emitted by human activities (see

Figure Q7-1). Methyl bromide, used primarily as an agricultural fumigant, is also a significant source of bromine

to the atmosphere.

Human emissions of the principal chlorine- and

bromine-containing gases have increased substantially

since the middle of the 20th century (see Q16). The result

has been global ozone depletion, with the greatest losses

occurring in polar regions (see Q11 to Q13).

Other human sources of chlorine and bromine.

Other chlorine- and bromine-containing gases are released

regularly in human activities. Common examples are the

use of chlorine gases to disinfect swimming pools and

wastewater, fossil fuel burning, and various industrial

processes. These activities do not contribute significantly

to stratospheric amounts of chlorine and bromine because

either the global source is small or the emitted gases are

short-lived (very reactive or highly soluble) and, therefore, are removed from the atmosphere before they reach

the stratosphere.

Natural sources of chlorine and bromine. There

are a few halogen source gases present in the stratosphere

that have large natural sources. These include methyl

chloride (CH3Cl) and methyl bromide (CH3Br), both of

which are emitted by oceanic and terrestrial ecosystems.

Natural sources of these two gases contribute about 17%

of the chlorine currently in the stratosphere and about 30%

of the bromine (see Figure Q7-1). Very short-lived source

gases containing bromine, such as bromoform (CHBr3),

are also released to the atmosphere primarily from the

oceans. Only a small fraction of these emissions reaches

the stratosphere, because these gases are rapidly removed

in the lower atmosphere. The contribution of these very

short-lived gases to stratospheric bromine is estimated to

be about 24%, but this has a large uncertainty. The contribution to stratospheric chlorine of short-lived chlorinated gases from natural and human sources is much

smaller (< 3%) and is included in the “Other gases” category in Figure Q7-1. Changes in the natural sources of

chlorine and bromine since the middle of the 20th century

are not the cause of observed ozone depletion.

Lifetimes and emissions. After emission, halogen

source gases are either naturally removed from the atmosphere or undergo chemical conversion. The time to

remove or convert about 60% of a gas is often called its

atmospheric “lifetime.” Lifetimes vary from less than 1

year to 100 years for the principal chlorine- and brominecontaining gases (see Table Q7-1). Gases with the shortest

lifetimes (e.g., the HCFCs, methyl bromide, methyl chloride, and the very short-lived gases) are substantially

destroyed in the troposphere, and therefore only a fraction of each emitted gas contributes to ozone depletion in

the stratosphere.

The amount of a halogen source gas present in the

atmosphere depends on the lifetime of the gas and the

Q7: What emissions from human activities lead to ozone depletion?

Certain industrial processes and consumer products result in the emission of “halogen source gases” to the

atmosphere. These gases bring chlorine and bromine to the stratosphere, which cause depletion of the ozone

layer. For example, chlorofluorocarbons (CFCs), once used in almost all refrigeration and air conditioning

systems, eventually reach the stratosphere, where they are broken apart to release ozone-depleting chlorine

atoms. Other examples of human-produced ozone-depleting gases are the “halons,” which are used in fire

extinguishers and contain ozone-depleting bromine atoms. The production and consumption of all principal

halogen source gases by human activities are regulated worldwide under the Montreal Protocol.
TWENTY QUESTIONS: 2006 UPDATE

179

amount emitted to the atmosphere. Emissions vary greatly

for the principal source gases, as indicated in Table Q7-1.

Emissions of most gases regulated by the Montreal

Protocol have decreased since 1990, and emissions from

all regulated gases are expected to decrease in the coming

decades (see Q16).

Ozone Depletion Potential. The halogen source

gases in Figure Q7-1 are also known as “ozone-depleting

substances” because they are converted in the stratosphere

to reactive gases containing chlorine and bromine (see

Q8). Some of these reactive gases participate in reactions

that destroy ozone (see Q9). Ozone-depleting substances

are compared in their effectiveness to destroy stratospheric ozone using the “Ozone Depletion Potential”

(ODP), as listed in Table Q7-1 (see Q18). A gas with a

larger ODP has a greater potential to destroy ozone over

its lifetime in the atmosphere. The ODP is calculated on

a “per mass” basis for each gas relative to CFC-11, which

has an ODP defined to be 1. Halon-1211 and halon-1301

have ODPs significantly larger than CFC-11 and most

other emitted gases, because bromine is much more effective overall (about 60 times) on a per-atom basis than chlorine in chemical reactions that destroy ozone in the stratosphere. The gases with small ODP values generally have

short atmospheric lifetimes or fewer chlorine and bromine

atoms. The production and consumption of all principal

halogen source gases by humans are regulated under the

provisions of the Montreal Protocol (see Q15).

Fluorine and iodine. Fluorine and iodine are also

halogen atoms. Many of the source gases in Figure Q7-1

also contain fluorine atoms in addition to chlorine or

bromine. After the source gases undergo conversion in

the stratosphere (see Q6), the fluorine content of these

gases is left in chemical forms that do not cause ozone

depletion. Iodine is a component of several gases that are

naturally emitted from the oceans. Although iodine can

Figure Q7-1. Stratospheric source gases. A variety of gases transport chlorine and bromine into the stratosphere.

These gases, called halogen source gases, are emitted from natural sources and by human activities. These partitioned

columns show how the principal chlorine and bromine source gases contribute to the respective total amounts of chlorine

and bromine as measured in 2004. Note the large difference in the vertical scales: total chlorine in the stratosphere is

160 times more abundant than total bromine. For chlorine, human activities account for most that reaches the stratosphere. The CFCs are the most abundant of the chlorine-containing gases released in human activities. Methyl chloride

is the most important natural source of chlorine. HCFCs, which are substitute gases for CFCs and also are regulated

under the Montreal Protocol, are a small but growing fraction of chlorine-containing gases. The “Other gases” category

includes minor CFCs and short-lived gases. For bromine that reaches the stratosphere, halons and methyl bromide are

the largest sources. Both gases are released in human activities. Methyl bromide has an additional natural source.

Natural sources are a larger fraction of total bromine than of total chlorine. (The unit “parts per trillion” is used here as a

measure of the relative abundance of a gas in air: 1 part per trillion indicates the presence of one molecule of a gas per

trillion other air molecules.)

Primary Sources of Chlorine and Bromine for the Stratosphere in 2004
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180

participate in ozone destruction reactions, these iodinecontaining source gases generally have very short lifetimes and, as a result, most are removed in the troposphere

before they reach the stratosphere.

Other gases. Other gases that influence stratospheric ozone abundances also have increased in the

stratosphere as a result of human activities. Important

examples are methane (CH4) and nitrous oxide (N2O),

which react in the stratosphere to form water vapor and

reactive hydrogen, and nitrogen oxides, respectively.

These reactive products also participate in the production and loss balance of stratospheric ozone (see Q2).

The overall effect of increases in these other gases on

ozone is much smaller than that caused by increases in

chlorine- and bromine-containing gases from human

activities (see Q18).

Heavier-Than-Air CFCs

CFCs and other halogen source gases reach the stratosphere despite the fact that they are “heavier

than air.” All the principal source gases are emitted and accumulate in the lower atmosphere (troposphere).

The distributions of gases in the troposphere and stratosphere are not controlled by the molecular weight of

the gases because air is in continual motion in these regions as a result of winds and convection. Air

motions ensure that most source gases become horizontally and vertically well mixed throughout the troposphere in a matter of months. It is this well-mixed air that enters the lower stratosphere from upward air

motions in tropical regions, bringing with it source gas molecules emitted from a wide variety of locations on

Earthʼs surface.

Atmospheric measurements confirm that halogen source gases with long atmospheric lifetimes are

well mixed in the troposphere and are present in the stratosphere (see Figure Q8-2). The amounts found in

these regions are consistent with the emissions estimates reported by industries and governments.

Measurements also show that gases that are “lighter than air,” such as hydrogen (H2) and methane (CH4),

are also well mixed in the troposphere, as expected. Only at altitudes well above the troposphere and stratosphere (above 85 kilometers (53 miles)), where much less air is present, does the influence of winds and

convection diminish to the point where heavy gases begin to separate from lighter gases as a result of

gravity.

Table Q7-1. Atmospheric lifetimes, emissions,

and Ozone Depletion Potentials of halogen

source gases. a

Halogen Atmospheric Global Ozone

Source Lifetime Emissions Depletion

Gas (years) in 2003 b Potential

(ODP) d

Chlorine

CFC-12 100 101-144 1

CFC-113 85 1-15 1

CFC-11 45 60-126 1

Carbon tetrachloride

(CCl4) 26 58-131 0.73

HCFCs 1-26 312-403 0.02-0.12

Methyl chloroform

(CH3CCl3) 5 ~20 0.12

Methyl chloride 1.0 1700-13600 0.02

Bromine

Halon-1301 65 ~3 16

Halon-1211 16 7-10 7.1

Methyl bromide

(CH3Br) 0.7 160-200 0.51

Very short-lived gases

(e.g., CHBr3) < 0.5 c c

a Includes both human activities and natural sources.

b Emission in gigagrams per year (1 gigagram = 109 grams =

1000 metric tons). c Estimates are uncertain for most species. d Values are calculated for emissions of equal mass for each gas.
TWENTY QUESTIONS: 2006 UPDATE

181

Reactive gases containing the halogens chlorine and

bromine lead to the chemical destruction of stratospheric

ozone. Halogen-containing gases present in the stratosphere can be divided into two groups: halogen source

gases and reactive halogen gases. The source gases are

emitted at Earth’s surface by natural processes and by

human activities (see Q7). Once in the stratosphere, the

halogen source gases chemically convert at different rates

to form the reactive halogen gases. The conversion occurs

in the stratosphere instead of the troposphere because solar

UV radiation is more intense in the stratosphere.

Reactive halogen gases. The chemical conversion

of halogen source gases, which involves ultraviolet sunlight and other chemical reactions, produces a number of

reactive halogen gases. These reactive gases contain all

of the chlorine and bromine atoms originally present in

the source gases.

The most important reactive chlorine- and brominecontaining gases that form in the stratosphere are shown

in Figure Q8-1. Away from polar regions, the most abundant are hydrogen chloride (HCl) and chlorine nitrate

(ClONO2). These two gases are considered reservoir

gases because they do not react directly with ozone but

can be converted to the most reactive forms that do chemically destroy ozone. The most reactive forms are chlorine monoxide (ClO) and bromine monoxide (BrO), and

chlorine and bromine atoms (Cl and Br). A large fraction

of available stratospheric bromine is generally in the form

of BrO, whereas usually only a small fraction of stratospheric chlorine is in the form of ClO. In polar regions,

the reservoirs ClONO2 and HCl undergo a further conversion on polar stratospheric clouds to form ClO (see Q10).

In that case, ClO becomes a large fraction of available

reactive chlorine.

Reactive chlorine observations. Reactive chlorine

gases have been observed extensively in the stratosphere

with both local and remote measurement techniques. The

measurements from space at middle latitudes displayed in

Q8: What are the reactive halogen gases that destroy stratospheric ozone?

Emissions from human activities and natural processes include large sources of chlorine- and bromine-containing gases that eventually reach the stratosphere. When exposed to ultraviolet radiation from the Sun, these

halogen source gases are converted to more reactive gases also containing chlorine and bromine. Important

examples of the reactive gases that destroy stratospheric ozone are chlorine monoxide (ClO) and bromine

monoxide (BrO). These reactive gases participate in “catalytic” reaction cycles that efficiently destroy ozone.

Volcanoes can emit some chlorine-containing gases, but these gases are ones that readily dissolve in rainwater

and ice and are usually “washed out” of the atmosphere before they can reach the stratosphere.

Figure Q8-1. Conversion of halogen source gases. Halogen source gases (also known as ozone-depleting substances)

are chemically converted to reactive halogen gases primarily in the stratosphere. The conversion requires ultraviolet sunlight

and a few other chemical reactions. The short-lived gases undergo some conversion in the troposphere. The reactive halogen gases contain all the chlorine and bromine originally present in the source gases. The reactive gases separate into reservoir gases, which do not destroy ozone, and reactive gases, which participate in ozone destruction cycles (see Q9).
TWENTY QUESTIONS: 2006 UPDATE

182

Figure Q8-2 are representative of how the amounts of

chlorine-containing gases change between the surface and

the upper stratosphere. Available chlorine (see red line in

Figure Q8-2) is the sum of chlorine contained in halogen

source gases and the reactive gases HCl, ClONO2, ClO, and

other minor gases. Available chlorine is constant within a

few percent from the surface to 47 kilometers (31 miles)

altitude. In the troposphere, available chlorine is contained

almost entirely in the source gases described in Figure Q7-

1. At higher altitudes, the source gases become a smaller

fraction of available chlorine as they are converted to reactive chlorine gases. At the highest altitudes, available chlorine is all in the form of reactive chlorine gases.

In the altitude range of the ozone layer at midlatitudes, as shown in Figure Q8-2, the reactive chlorine gases

HCl and ClONO2 account for most of available chlorine.

ClO, the most reactive gas in ozone depletion, is a small

fraction of available chlorine. This small value limits the

amount of ozone destruction that occurs outside of polar

regions.

Reactive chlorine in polar regions. Reactive chlorine gases in polar regions in summer look similar to the

altitude profiles shown in Figure Q8-2. In winter, however, the presence of polar stratospheric clouds (PSCs)

causes further chemical changes (see Q10). PSCs convert HCl and ClONO2 to ClO when temperatures are near

minimum values in the winter Arctic and Antarctic stratosphere. In that case, ClO becomes the principal reactive

chlorine species in sunlit regions and ozone loss becomes

very rapid. An example of the late-winter ClO and ozone

distributions is shown in Figure Q8-3 for the Antarctic

stratosphere. These space-based measurements show that

ClO abundances are high in the lower stratosphere over a

region that exceeds the size of the Antarctic continent

(greater than 13 million square kilometers or 5 million

square miles). The peak abundance of ClO exceeds 1500

parts per trillion, which is much larger than typical midlatitude values shown in Figure Q8-2 and represents a

large fraction of reactive chlorine in that altitude region.

Because high ClO amounts cause rapid ozone loss (see

Q9), ozone depletion is found in regions of elevated ClO

(see Figure Q8-3).

Reactive bromine observations. Fewer measurements are available for reactive bromine gases in the lower

stratosphere than for reactive chlorine, in part because of

the lower abundance of bromine. The most widely

observed bromine gas is bromine monoxide (BrO).

Recent observations have shown that measured BrO abundances in the stratosphere are larger than expected from

the conversion of the halons and methyl bromide to BrO,

suggesting a significant contribution from the very shortlived bromine-containing gases.

Other sources. Some reactive halogen gases are also

produced at Earth’s surface by natural processes and by

human activities. However, because reactive halogen

gases are soluble in water, almost all become trapped in

the lower atmosphere by dissolving in rainwater and ice,

and ultimately are returned to Earth’s surface before they

can reach the stratosphere. For example, reactive chlorine is present in the atmosphere as sea salt (sodium chloride) produced by evaporation of ocean spray. Because

sea salt dissolves in water, this chlorine is removed and

does not reach the stratosphere in appreciable quantities.

Another ground-level source is emission of chlorine gases

from swimming pools, household bleach, and other uses.

Figure Q8-2. Reactive chlorine gas observations.

The abundances of chlorine source gases and reactive chlorine gases as measured from space are displayed with altitude for a midlatitude location. In the

troposphere (below about 10 kilometers), all chlorine is

contained in the source gases. In the stratosphere,

reactive chlorine gases increase with altitude as chlorine source gases decrease. This is a consequence of

chemical reactions involving ultraviolet sunlight (see

Figure Q8-1). The principal reactive gases formed are

HCl, ClONO2, and ClO. Summing the source gases

with the reactive gases gives total available chlorine,

which is nearly constant with altitude up to 47 km. In

the ozone layer, HCl and ClONO2 are the most abundant reactive chlorine gases. (The unit “parts per trillion” is defined in the caption of Figure Q7-1.)

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When released to the atmosphere, this chlorine is rapidly

converted to forms that are soluble in water and removed.

The Space Shuttle and other rocket motors release reactive chlorine gases directly in the stratosphere: in this case,

the quantities are very small in comparison with other tropospheric sources.

Volcanoes. Volcanic plumes generally contain large

quantities of chlorine in the form of hydrogen chloride

(HCl). Because the plumes also contain a considerable

amount of water vapor, the HCl is efficiently scavenged

by rainwater and ice and removed from the atmosphere.

As a result, most of the HCl in the plume does not enter

the stratosphere. After large recent eruptions, the

increase in HCl in the stratosphere has been small compared with the total amount of chlorine in the stratosphere from other sources.

Figure Q8-3. Antarctic chlorine monoxide and ozone.

Satellite instruments monitor ozone and reactive chlorine gases in the global stratosphere. Results are shown

here for Antarctic winter for a narrow altitude region within

the ozone layer. In winter, chlorine monoxide (ClO)

reaches high values (1500 parts per trillion) in the ozone

layer, much higher than observed anywhere else in the

stratosphere because ClO is produced by reactions on

polar stratospheric clouds (see Q10). These high ClO

values in the lower stratosphere last for 1 to 2 months,

cover an area that at times exceeds that of the Antarctic

continent, and efficiently destroy ozone in sunlit regions

in late winter/early spring. Ozone values measured

simultaneously within the ozone layer show very

depleted values.

Replacing the Loss of Ozone in the Stratosphere

The idea is sometimes put forth that humans could replace the loss of global stratospheric ozone by

making ozone and transporting it to the stratosphere. Ozone amounts in the stratosphere reflect a balance

between continual production and destruction by mostly naturally occurring reactions (see Q2). The addition of chlorine and bromine to the stratosphere from human activities has increased ozone destruction and

lowered stratospheric ozone amounts. Adding manufactured ozone to the stratosphere would upset the

existing balance. As a consequence, most added ozone would be destroyed in chemical reactions within

weeks to months as the balance was restored. So, it is not practical to consider replacing the loss of global

stratospheric ozone because the replacement effort would need to continue indefinitely, or as long as

increased chlorine and bromine amounts remained.

Other practical difficulties in replacing stratospheric ozone are the large amounts of ozone required

and the delivery method. The total amount of atmospheric ozone is approximately 3,000 megatons (1

megaton = 1 billion kilograms) with most residing in the stratosphere. The replacement of the average

global ozone loss of about 4% would require 120 megatons of stratospheric ozone to be distributed

throughout the layer located many kilometers above Earthʼs surface. The energy required to produce this

amount of ozone would be a significant fraction of the electrical power generated in the United States,

which is now approximately 5 trillion kilowatt hours. Processing and storing requirements for ozone,

which is explosive and toxic in large quantities, would increase the energy requirement. In addition,

methods suitable to deliver and distribute large amounts of ozone to the stratosphere have not been

demonstrated. Concerns for a global delivery system would include further significant energy use and

unforeseen environmental consequences.

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Stratospheric ozone is destroyed by reactions involving

reactive halogen gases, which are produced in the chemical conversion of halogen source gases (see Figure Q8-1).

The most reactive of these gases are chlorine monoxide

(ClO), bromine monoxide (BrO), and chlorine and bromine

atoms (Cl and Br). These gases participate in three principal reaction cycles that destroy ozone.

Cycle 1. Ozone destruction Cycle 1 is illustrated

in Figure Q9-1. The cycle is made up of two basic reactions: ClO + O and Cl + O3. The net result of Cycle 1 is

to convert one ozone molecule and one oxygen atom

into two oxygen molecules. In each cycle, chlorine acts

as a catalyst because ClO and Cl react and are reformed.

In this way, one Cl atom participates in many cycles,

destroying many ozone molecules. For typical stratospheric conditions at middle or low latitudes, a single

chlorine atom can destroy hundreds of ozone molecules

before it happens to react with another gas, breaking the

catalytic cycle.

Polar Cycles 2 and 3. The abundance of ClO is

greatly increased in polar regions during winter as a result

of reactions on the surfaces of polar stratospheric cloud

(PSC) particles (see Q10). Cycles 2 and 3 (see Figure

Q9-2) become the dominant reaction mechanisms for

polar ozone loss because of the high abundances of ClO

and the relatively low abundance of atomic oxygen

(which limits the rate of ozone loss by Cycle 1). Cycle 2

begins with the self-reaction of ClO. Cycle 3, which

begins with the reaction of ClO with BrO, has two reaction pathways to produce either Cl and Br or BrCl. The

net result of both cycles is to destroy two ozone molecules and create three oxygen molecules. Cycles 2 and 3

account for most of the ozone loss observed in the Arctic

and Antarctic stratospheres in the late winter/early spring

Figure Q9-1. Ozone destruction

Cycle 1. The destruction of ozone in

Cycle 1 involves two separate chemical

reactions. The net or overall reaction is

that of atomic oxygen with ozone, forming two oxygen molecules. The cycle

can be considered to begin with either

ClO or Cl. When starting with ClO, the

first reaction is ClO with O to form Cl.

Cl then reacts with (and thereby

destroys) ozone and reforms ClO. The

cycle then begins again with another

reaction of ClO with O. Because Cl or

ClO is reformed each time an ozone

molecule is destroyed, chlorine is considered a catalyst for ozone destruction. Atomic oxygen (O) is formed

when ultraviolet sunlight reacts with

ozone and oxygen molecules. Cycle 1

is most important in the stratosphere at

tropical and middle latitudes, where

ultraviolet sunlight is most intense.

Q9: What are the chlorine and bromine reactions that destroy stratospheric ozone?

Reactive gases containing chlorine and bromine destroy stratospheric ozone in “catalytic” cycles made up of

two or more separate reactions. As a result, a single chlorine or bromine atom can destroy many hundreds of

ozone molecules before it reacts with another gas, breaking the cycle. In this way, a small amount of reactive

chlorine or bromine has a large impact on the ozone layer. Certain ozone destruction reactions become most

effective in polar regions because the reactive gas chlorine monoxide reaches very high levels there in the late

winter/early spring season.

Ozone Destruction Cycle 1
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185

season (see Q11 and Q12). At high ClO abundances, the

rate of ozone destruction can reach 2 to 3% per day in

late winter/early spring.

Sunlight requirement. Sunlight is required to complete and maintain Cycles 1 through 3. Cycle 1 requires

sunlight because atomic oxygen is formed only with ultraviolet sunlight. Cycle 1 is most important in the stratosphere at tropical and middle latitudes, where ultraviolet

sunlight is most intense.

Cycles 2 and 3 require visible sunlight to complete

the reaction cycles and to maintain ClO abundances. In

the continuous darkness of winter in the polar stratospheres, reaction Cycles 2 and 3 cannot occur. It is only in

late winter/early spring when sunlight returns to the polar

regions that these cycles can occur. Therefore, the greatest

destruction of ozone occurs in the partially to fully sunlit

periods after midwinter in the polar stratospheres. The

visible sunlight needed in Cycles 2 and 3 is not sufficient

to form ozone because this process requires ultraviolet

sunlight. In the stratosphere in the late winter/early spring

period, ultraviolet sunlight is weak because Sun angles

are low. As a result, ozone is destroyed by Cycles 2 and 3

in the sunlit winter stratosphere but is not produced in significant amounts.

Other reactions. Global ozone abundances are

controlled by many reactions that both produce and

destroy ozone (see Q2). Chlorine and bromine catalytic

reactions are but one group of ozone destruction reactions. Reactive hydrogen and reactive nitrogen gases,

for example, are involved in other catalytic ozonedestruction cycles that also occur in the stratosphere.

These reactions occur naturally in the stratosphere and

their importance has not been as strongly influenced by

human activities as have reactions involving halogens.

Figure Q9-2. Polar ozone destruction Cycles 2 and 3. Significant destruction of ozone occurs in polar regions because

ClO abundances reach large values. In this case, the cycles initiated by the reaction of ClO with another ClO (Cycle 2) or the

reaction of ClO with BrO (Cycle 3) efficiently destroy ozone. The net reaction in both cases is two ozone molecules forming

three oxygen molecules. The reaction of ClO with BrO has two pathways to form the Cl and Br product gases. Ozone

destruction Cycles 2 and 3 are catalytic, as illustrated for Cycle 1 in Figure Q9-1, because chlorine and bromine gases react

and are reformed in each cycle. Sunlight is required to complete each cycle and to help form and maintain ClO abundances.

Ozone Destruction Cycles
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186

The severe depletion of stratospheric ozone in

Antarctic winter is known as the “ozone hole” (see Q11).

Severe depletion first appeared over Antarctica because

atmospheric conditions there increase the effectiveness of

ozone destruction by reactive halogen gases (see Q8). The

formation of the Antarctic ozone hole requires abundant

reactive halogen gases, temperatures low enough to form

polar stratospheric clouds (PSCs), isolation of air from

other stratospheric regions, and sunlight.

Distributing halogen gases. Halogen source gases

emitted at Earth’s surface are present in comparable abundances throughout the stratosphere in both hemispheres

even though most of the emissions occur in the Northern

Hemisphere. The abundances are comparable because

most source gases have no important natural removal

processes in the lower atmosphere and because winds and

warm-air convection redistribute and mix air efficiently

throughout the troposphere. Halogen gases (in the form

of source gases and some reactive products) enter the stratosphere primarily from the tropical upper troposphere.

Atmospheric air motions then transport them upward and

toward the poles in both hemispheres.

Low temperatures. The severe ozone destruction

represented by the ozone hole requires that low temperaQ10: Why has an “ozone hole” appeared over Antarctica when ozone-depleting gases are

present throughout the stratosphere?

Ozone-depleting gases are present throughout the stratospheric ozone layer because they are transported

great distances by atmospheric air motions. The severe depletion of the Antarctic ozone layer known as the

“ozone hole” occurs because of the special weather conditions that exist there and nowhere else on the globe.

The very low temperatures of the Antarctic stratosphere create ice clouds called polar stratospheric clouds

(PSCs). Special reactions that occur on PSCs and the relative isolation of polar stratospheric air allow chlorine and bromine reactions to produce the ozone hole in Antarctic springtime.

Figure Q10-1. Arctic and Antarctic temperatures. Stratospheric air temperatures

in both polar regions reach minimum values

in the lower stratosphere in the winter

season. Average minimum values over

Antarctica are as low as –90°C in July and

August in a typical year. Over the Arctic,

average minimum values are near –80°C in

January and February. Polar stratospheric

clouds (PSCs) are formed when winter minimum temperatures fall below the formation

temperature (about –78°C). This occurs on

average for 1 to 2 months over the Arctic and

5 to 6 months over Antarctica (see heavy red

and blue lines). Reactions on PSCs cause

the highly reactive chlorine gas ClO to be

formed, which increases the destruction of

ozone (see Q9). The range of winter minimum temperatures found in the Arctic is

much greater than in the Antarctic. In some

years, PSC formation temperatures are not

reached in the Arctic, and significant ozone

depletion does not occur. In the Antarctic,

PSCs are present for many months, and

severe ozone depletion now occurs in each

winter season.

Minimum Air Temperatures in the

Polar Lower Stratosphere
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187

tures be present over a range of stratospheric altitudes,

over large geographical regions, and for extended time

periods. Low temperatures are important because they

allow polar stratospheric clouds (PSCs) to form.

Reactions on the surfaces of the cloud particles initiate a

remarkable increase in the most reactive halogen gases

(see below and Q8). Temperatures are lowest in the stratosphere over both polar regions in winter. In the Antarctic

winter, minimum temperatures are generally lower and

less variable than in the Arctic winter (see Figure Q10-1).

Antarctic temperatures also remain below the PSC formation temperature for much longer periods during

winter. This occurs, in part, because there are significant

meteorological differences between the hemispheres,

resulting from the differences in the distributions of land,

ocean, and mountains at middle and high latitudes. The

winter temperatures are low enough for PSCs to form for

nearly the entire Antarctic winter but usually only for part

of every Arctic winter.

Isolated conditions. Air in the polar stratospheric

regions is relatively isolated from other stratospheric

regions for long periods in the winter months. The isolation comes about because of strong winds that encircle

the poles, preventing substantial motion of air in or out of

the polar stratospheres. The isolation is much more effective in the Antarctic than the Arctic. Once chemical

changes occur in the cold air as a result of the presence of

PSCs, the changes remain for many weeks to months.

Polar stratospheric clouds (PSCs). Polar stratospheric clouds cause changes in the relative abundances

of reactive chlorine gases. Reactions occur on the surfaces of PSC particles that convert the reservoir forms of

reactive chlorine gases, ClONO2 and HCl, to the most

reactive form, ClO (see Figure Q8-1). ClO increases from

a small fraction of available reactive chlorine gases to

nearly all that is available (see Q8). With increased ClO,

additional catalytic cycles involving ClO and BrO become

active in the chemical destruction of ozone when sunlight

is available (see Q9).

PSCs form when stratospheric temperatures fall

below about –78°C (–108°F) in polar regions (see Figure

Q10-1). As a result, PSCs are often found over large areas

of the winter polar regions and over a significant altitude

range. At low polar temperatures, nitric acid (HNO3) and

water condense on preexisting sulfur-containing particles

to form solid and liquid PSC particles. At even lower temperatures, ice particles also form. PSC particles grow

large enough and are numerous enough that cloud-like

features can be observed from the ground under certain

conditions, particularly when the Sun is near the horizon

(see Figure Q10-2). PSCs are often found near mountain

ranges in polar regions because the motion of air over the

mountains can cause local cooling of stratospheric air.

When temperatures increase by early spring, PSCs

no longer form and the production of ClO ends. Without

continued ClO production, ClO amounts decrease as other

chemical reactions reform ClONO2 and HCl. As a result,

the intense period of ozone depletion ends.

PSC removal. Once formed, PSC particles move

downward because of gravity. The largest particles move

down several kilometers or more in the stratosphere

during the low-temperature winter/spring period. Because

most PSCs contain nitric acid, their downward motion

removes nitric acid from regions of the ozone layer. That

process is called denitrification. With less nitric acid, the

Figure Q10-2. Polar stratospheric clouds. This photograph of an Arctic polar stratospheric cloud (PSC) was

taken from the ground at Kiruna, Sweden (67°N), on 27

January 2000. PSCs form during winters in the Arctic and

Antarctic stratospheres. The particles grow from the condensation of water and nitric acid (HNO3). The clouds

often can be seen with the human eye when the Sun is

near the horizon. Reactions on PSCs cause the highly

reactive chlorine gas ClO to be formed, which is very effective in the chemical destruction of ozone (see Q9).

Arctic Polar Stratospheric Clouds
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188

highly reactive chlorine gas ClO remains chemically

active for a longer period, thereby increasing chemical

ozone destruction. Denitrification occurs each winter in

the Antarctic and in some, but not all, Arctic winters,

because PSC formation temperatures are required over an

extensive time period.

Discovering the role of PSCs. The formation of

PSCs has been recognized for many years from groundbased observations. However, the geographical and altitude extent of PSCs in both polar regions was not known

fully until PSCs were observed by a satellite instrument

in the late 1970s. The role of PSCs in converting reactive

chlorine gases to ClO was not understood until after the

discovery of the Antarctic ozone hole in 1985. Our understanding of the PSC role developed from laboratory

studies of their surface reactivity, computer modeling

studies of polar stratospheric chemistry, and sampling of

PSC particles and reactive chlorine gases, such as ClO, in

the polar stratospheric regions.

The Discovery of the Antarctic Ozone Hole

The first decreases in Antarctic total ozone were observed in the early 1980s over research stations

located on the Antarctic continent. The measurements were made with ground-based Dobson spectrophotometers (see box in Q5). The observations showed unusually low total overhead ozone during the late

winter/early spring months of September, October, and November. Total ozone was lower in these months

compared with previous observations made as early as 1957. The early published reports came from the

British Antarctic Survey and the Japan Meteorological Agency. The results became more widely known in

the international community after three scientists from the British Antarctic Survey published them in the

journal Nature in 1985. Soon after, satellite measurements confirmed the spring ozone depletion and further showed that in each late winter/early spring season starting in the early 1980s, the depletion extended

over a large region centered near the South Pole. The term “ozone hole” came about from satellite images

of total ozone that showed very low values encircling the Antarctic continent each spring (see Q11).

Currently, the formation and severity of the Antarctic “ozone hole” are documented each year by a combination of satellite, ground-based, and balloon observations of ozone.
TWENTY QUESTIONS: 2006 UPDATE

189

The severe depletion of Antarctic ozone, known as

the “ozone hole,” was first observed in the early 1980s.

The depletion is attributable to chemical destruction by

reactive halogen gases, which increased in the stratosphere in the latter half of the 20th century (see Q16).

Conditions in the Antarctic winter stratosphere are highly

suitable for ozone depletion because of (1) the long

periods of extremely low temperatures, which promote

polar stratospheric cloud (PSC) formation; (2) the abundance of reactive halogen gases, which chemically destroy

ozone; and (3) the isolation of stratospheric air during the

winter, which allows time for chemical destruction to

occur (see Q10). The severity of Antarctic ozone depletion can be seen using satellite observations of total ozone,

ozone altitude profiles, and long-term average values of

polar total ozone.

Antarctic ozone hole. The most widely used images

of Antarctic ozone depletion are those from space-based

measurements of total ozone. Satellite images made

during Antarctic winter and spring show a large region

centered near the South Pole in which total ozone is highly

depleted (see Figure Q11-1). This region has come to be

called the “ozone hole” because of the near-circular contours of low ozone values in the images. The area of the

ozone hole is defined here as the area contained within

the 220-Dobson unit (DU) contour in total ozone maps

(light blue color in Figure Q11-1). The maximum area

has reached 25 million square kilometers (about 10 million square miles) in recent years, which is nearly twice

the area of the Antarctic continent (see Figure Q11-2).

Minimum values of total ozone inside the ozone hole averaged in late September have reached below 100 DU,

which is well below normal springtime values of about

200 DU (see Figure Q11-2).

Altitude profiles of Antarctic ozone. Ozone within

the “ozone hole” is also measured using balloonborne

instruments (see Q5). Balloon measurements show

changes within the ozone layer, the vertical region that

contains the highest ozone abundances in the stratosphere.

At geographic locations where the lowest total ozone

values occur in ozone hole images, balloon measurements

show that the chemical destruction of ozone is complete

over a vertical region of several kilometers. Balloon

measurements shown in Figure Q11-3 give an example of

such depletion over South Pole, Antarctica, on 2 October

2001. The altitude region of total depletion (14-20 kilometers) in the profile corresponds to the region of lowest

III. STRATOSPHERIC OZONE DEPLETION

Q11: How severe is the depletion of the Antarctic ozone layer?

Figure Q11-1. Antarctic “ozone hole.” Total ozone

values are shown for high southern latitudes as measured by a satellite instrument. The dark blue and purple

regions over the Antarctic continent show the severe

ozone depletion or “ozone hole” now found during every

spring. Minimum values of total ozone inside the ozone

hole are close to 100 Dobson units (DU) compared with

normal springtime values of about 200 DU (see Q4). In

late spring or early summer (November-December) the

ozone hole disappears in satellite images as ozonedepleted air is displaced and mixed with ozone-rich air

transported poleward from outside the ozone hole.

Severe depletion of the Antarctic ozone layer was first observed in the early 1980s. Antarctic ozone depletion

is seasonal, occurring primarily in late winter and early spring (August-November). Peak depletion occurs in

early October when ozone is often completely destroyed over a range of altitudes, reducing overhead total

ozone by as much as two-thirds at some locations. This severe depletion creates the “ozone hole” in images

of Antarctic total ozone made from space. In most years the maximum area of the ozone hole far exceeds the

size of the Antarctic continent.

Antarctic Ozone Hole
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190

winter temperatures and highest chlorine monoxide (ClO)

abundances. The average South Pole ozone profiles for

the decades 1962-1971 and 1992-2001 (see Figure Q11-

3) show how reactive halogen gases have dramatically

altered the ozone layer. For the 1960s, the ozone layer is

clearly evident in the October average profile and has a

peak near 16 kilometers. For the 1990s, minimum average

values in the center of the layer have fallen by 90% from

the earlier values.

Long-term total ozone changes. Low winter temperatures and isolated conditions occur each year in the

Antarctic stratosphere, but significant spring ozone depletion has been observed every year only since the early

1980s. In prior years, the amounts of reactive halogen

gases in the stratosphere were insufficient to cause significant depletion. Satellite observations can be used to

examine how ozone depletion has changed with time in

both polar regions for the last three decades. Changes in

ozone hole areas and minimum Antarctic ozone amounts

are shown in Figure Q11-2. Depletion has increased since

1980 to become fairly stable in the 1990s and early 2000s,

with the exception of 2002 (see Q11-box). Total ozone

averaged over the Antarctic region in late winter/early

spring shows similar features (Figure Q12-1). Average

values decreased steadily through the 1980s and 1990s,

reaching minimum values that were 37% less than in preozone-hole years (1970-1982). The year-to-year changes

in the average values reflect variations in the meteorological conditions, which affect the extent of low polar temperatures and the transport of air into and out of the

Antarctic winter stratosphere (see Figure Q11-box).

However, essentially all of the ozone depletion in the

Antarctic in most years is attributable to chemical loss

from reactive halogen gases.

Restoring ozone in spring. The depletion of

Antarctic ozone occurs primarily in the late winter/early

spring season. In spring, temperatures in the lower polar

stratosphere eventually warm, thereby ending PSC formation as well as the most effective chemical cycles that

destroy ozone (see Q10). The transport of air between the

polar stratosphere and lower latitudes also increases

during this time, ending winter isolation. This allows

ozone-rich air to be transported to polar regions, displacing air in which ozone has been severely depleted.

This displaced air is mixed at lower latitudes with more

abundant ozone-rich air. As a result, the ozone hole disappears by December and Antarctic ozone amounts

remain near normal until the next winter season.

Figure Q11-2. Antarctic ozone hole

features. Values are shown for key

parameters of the Antarctic ozone hole:

the area enclosed by the 220-DU total

ozone contour and the minimum total

ozone amount, as determined from

space-based observations. The values

are averaged for each year near the

peak of ozone depletion, as defined by

the dates shown in each panel. The

ozone hole areas are contrasted to the

areas of continents in the upper panel.

The intensity of ozone depletion gradually increased beginning in 1980. In the

1990s, the depletion reached fairly

steady values, except for the anomalously low depletion in 2002 (see Figure

Q11-box). The intensity of Antarctic

ozone depletion will decrease as part of

the ozone recovery process (see Q19

and Q20).

Antarctic Ozone Depletion
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191

Figure Q11-3. Arctic and Antarctic ozone distribution. The stratospheric ozone layer resides between about 10

and 50 kilometers (6 to 31 miles) above Earthʼs surface over the globe. Long-term observations of the ozone layer

with balloonborne instruments allow the winter Antarctic and Arctic regions to be compared. In the Antarctic at the

South Pole, halogen gases have destroyed ozone in the ozone layer beginning in the 1980s. Before that period, the

ozone layer was clearly present, as shown here using average ozone values from balloon observations made between

1962 and 1971. In more recent years, as shown here for 2 October 2001, ozone is destroyed completely between 14

and 20 kilometers (8 to 12 miles) in the Antarctic in spring. Average October values in the ozone layer now are reduced

by 90% from pre-1980 values. The Arctic ozone layer is still present in spring as shown by the average March profile

obtained over Finland between 1988 and 1997. However, March Arctic ozone values in some years are often below

normal average values as shown here for 30 March 1996. In such years, winter minimum temperatures are generally

below PSC formation temperatures for long periods. Ozone abundances are shown here with the unit “milli-Pascals”

(mPa), which is a measure of absolute pressure (100 million mPa = atmospheric sea-level pressure).

Polar Ozone Depletion
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The Anomalous 2002 Antarctic Ozone Hole

The 2002 Antarctic ozone hole showed features that surprised scientists. They considered it anomalous at the time because the hole had much less area as viewed from space and much less ozone depletion as measured by minimum column ozone amounts when compared with values in several preceding

years (see Figure Q11-box). The 2002 ozone hole area and minimum ozone values stand out clearly in displays of the year-to-year changes in these quantities (see Figure Q11-2). The smaller area was unexpected because the conditions required to deplete ozone, namely low temperatures and available reactive

halogen gases, are not expected to have large year-to-year variations. Ozone was being depleted in

August and early September 2002, but the hole broke apart into two separate depleted regions during the

last week of September. The depletion in these two regions was significantly less than was observed inside

either the 2001 or 2003 ozone holes, but still substantially greater than was observed in the early 1980s.

The anomalous behavior in 2002 occurred because of specific atmospheric air motions that sometimes

occur in polar regions, not large decreases in reactive chlorine and bromine amounts in the Antarctic stratosphere. The Antarctic stratosphere was warmed by very strong, large-scale weather systems in 2002 that

originated in the lower atmosphere (troposphere) at midlatitudes in late September. In late September,

Antarctic temperatures are generally very low (see Q10) and ozone destruction rates are near their peak

values. These tropospheric systems traveled poleward and upward into the stratosphere, upsetting the circumpolar wind flow and warming the lower stratosphere where ozone depletion was ongoing. The higherthan-normal impact of these weather disturbances during the critical time period for ozone loss reduced the

total loss of ozone in 2002.

The warming in 2002 was unprecedented in Antarctic meteorological observations. Warming events

are difficult to predict because of their complex formation conditions.

Large Antarctic ozone depletion returned in 2003 through 2005, in a manner similar to that observed

from the mid-1990s to 2001 (see Figures Q11-box and Q11-2). The high ozone depletion found since the

mid-1990s, with the exception of 2002, is expected to be typical of coming years. A significant, sustained

reduction of Antarctic ozone depletion, defined as ozone recovery, requires the removal of halogen source

gases from the stratosphere (see Q19 and Q20).

Figure Q11-Box. Anomalous 2002 ozone hole. Views from space of the Antarctic ozone hole as observed on 24

September in each of three consecutive years. The hole split and elongated in 2002, reducing the total depletion of

ozone observed that year in comparison with 2001 and 2003. The anomalous depletion in 2002 is attributable to an

early warming of the polar stratosphere caused by air disturbances originating in midlatitudes, rather than to large

changes in the amounts of reactive chlorine and bromine in the Antarctic stratosphere.

Antarctic Ozone Hole
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193

Significant ozone depletion in the Arctic stratosphere

occurs in cold winters because of reactive halogen gases.

The depletion, however, is much less than the depletion

that now occurs in every Antarctic winter and spring.

Although Arctic depletion does not generally create persistent “ozone hole”-like features in Arctic total ozone

maps, depletion is observed in altitude profiles of ozone

and in long-term average values of polar ozone.

Altitude profiles of Arctic ozone. Arctic ozone is

measured using a variety of instruments (see Q5), as for the

Antarctic (see Q11). These measurements show changes

within the ozone layer, the vertical region that contains the

highest ozone abundances in the stratosphere. Figure Q11-2

shows an example of balloonborne measurements of a

depleted ozone profile in theArctic region on 30 March 1996,

and contrasts the depletion with that found in the Antarctic.

The 30 March spring profile shows much less depletion than

the 2 October spring profile in theAntarctic. In general, some

reduction in the Arctic ozone layer occurs each late

winter/early spring season. However, complete depletion

each year over a broad vertical layer, as is now common in

the Antarctic stratosphere, is not found in the Arctic.

Long-term total ozone changes. Satellite and

ground-based observations can be used to examine the

average total ozone abundances in the Arctic region for

the last three decades and to contrast them with Antarctic

abundances (see Figure Q12-1). Decreases from the preozone-hole average values (1970-1982) were observed in

the Arctic beginning in the 1980s, when similar changes

were occurring in the Antarctic. The decreases have

reached a maximum of about 30% but have remained

smaller than those found in the Antarctic since the mid1980s. The year-to-year changes in the Arctic and

Antarctic average ozone values reflect annual variations

in meteorological conditions that affect the extent of low

polar temperatures and the transport of air into and out of

the polar stratosphere. The effect of these variations is

generally greater for the Arctic than the Antarctic.

Q12: Is there depletion of the Arctic ozone layer?

Yes, significant depletion of the Arctic ozone layer now occurs in some years in the late winter/early spring

period (January-April). However, the maximum depletion is less severe than that observed in the Antarctic

and is more variable from year to year. A large and recurrent “ozone hole,” as found in the Antarctic stratosphere, does not occur in the Arctic.

Figure Q12-1. Average polar ozone.

Total ozone in polar regions is measured by

well-calibrated satellite instruments. Shown

here is a comparison of average springtime

total ozone values found between 1970 and

1982 (solid and dashed red lines) with those

in later years. Each point represents a

monthly average in October in the Antarctic

or in March in the Arctic. After 1982, significant ozone depletion is found in most years

in the Arctic and all years in the Antarctic.

The largest average depletions have

occurred in the Antarctic since 1990. The

ozone changes are the combination of

chemical destruction and natural variations.

Variations in meteorological conditions influence the year-to-year changes in depletion,

particularly in the Arctic. Essentially all of the decrease in the Antarctic and usually most of the decrease in the Arctic

each year are attributable to chemical destruction by reactive halogen gases. Average total ozone values over the Arctic

are naturally larger at the beginning of each winter season because more ozone is transported poleward each season in

the Northern Hemisphere than in the Southern Hemisphere.

Average Total Ozone in Polar Regions
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194

Arctic vs. Antarctic. The Arctic winter stratosphere

is generally warmer than its Antarctic counterpart (see

Figure Q10-1). Higher temperatures reduce polar stratospheric cloud (PSC) formation, which slows the conversion of reactive chlorine gases to form ClO and, as a consequence, reduces the amount of ozone depletion (see

Q10). Furthermore, the temperature and wind conditions

are much more variable in the Arctic from winter to winter

and within a winter season than in the Antarctic. Large

year-to-year differences occur in Arctic minimum temperatures and the duration of PSC-forming temperatures

into early spring. In a few Arctic winters, minimum temperatures are not low enough for PSCs to form. These

factors combine to cause ozone depletion to be variable

in the Arctic from year to year, with some years having

little to no ozone depletion.

As in the Antarctic, depletion of ozone in the Arctic

is confined to the late winter/early spring season. In

spring, temperatures in the lower stratosphere eventually

warm, thereby ending PSC formation as well as the most

effective chemical cycles that destroy ozone. The subsequent transport of ozone-rich air into the Arctic stratosphere displaces ozone-depleted air. As a result, ozone

layer abundances are restored to near-normal values until

the following winter.

High Arctic total ozone. A significant difference

exists between the Northern and Southern Hemispheres

in how ozone-rich stratospheric air is transported into the

polar regions from lower latitudes during fall and winter.

In the northern stratosphere, the poleward and downward

transport of ozone-rich air is stronger. As a result, total

ozone values in the Arctic are considerably higher than in

the Antarctic at the beginning of each winter season (see

Figure Q12-1).
TWENTY QUESTIONS: 2006 UPDATE

195

Stratospheric ozone has decreased over the globe

since the 1980s. The depletion, which in the period

1997-2005 averaged about 4% (see Figure Q13-1), is

larger than natural variations in ozone. The observations shown in Figure Q13-1 have been smoothed to

remove regular ozone changes that are due to seasonal

and solar effects (see Q14). The increase in reactive

halogen gases in the stratosphere is considered to be the

primary cause of the average depletion. The lowest

ozone values in recent years occurred following the

1991 eruption of Mt. Pinatubo, which increased the

number of sulfur-containing particles in the stratosphere. The particles remain in the stratosphere for several years, increasing the effectiveness of reactive

halogen gases in destroying ozone (see Q14).

Observed ozone depletion varies significantly with

latitude on the globe (see Figure Q13-1). The largest losses

occur at the highest southern latitudes as a result of the

severe ozone loss over Antarctica each late winter/early

spring period. The next largest losses are observed in the

Northern Hemisphere, caused in part by late winter/early

spring losses over the Arctic. Ozone-depleted air over both

polar regions is dispersed away from the poles during and

after each winter/spring period. Ozone depletion also

occurs directly at latitudes between the equator and polar

regions but is smaller because of the presence of lower

amounts of reactive halogen gases (see Q8).

Tropical regions. There has been little or no depletion of total ozone in the tropics (between about 20° latitude north and south of the equator in Figure Q13-1). In

Figure Q13-1. Global total ozone changes. Satellite

observations show a decrease in global total ozone

values over more than two decades. The top panel

compares global ozone values (annual averages) with

the average from the period 1964 to 1980. Seasonal

and solar effects have been removed from the data. On

average, global ozone decreased each year between

1980 and the early 1990s. The decrease worsened

during the few years when volcanic aerosol from the

Mt. Pinatubo eruption in 1991 remained in the stratosphere. Now global ozone is about 4% below the 1964-

to-1980 average. The bottom panel compares ozone

changes between 1980 and 2004 for different latitudes.

The largest decreases have occurred at the highest

latitudes in both hemispheres because of the large

winter/spring depletion in polar regions. The losses in

the Southern Hemisphere are greater than those in the

Northern Hemisphere because of the Antarctic ozone

hole. Long-term changes in the tropics are much

smaller because reactive halogen gases are less abundant in the tropical lower stratosphere.

Q13: How large is the depletion of the global ozone layer?

The ozone layer has been depleted gradually since 1980 and now is about an average of 4% lower over the

globe. The average depletion exceeds the natural variability of the ozone layer. The ozone loss is very small

near the equator and increases with latitude toward the poles. The larger polar depletion is primarily a result

of the late winter/early spring ozone destruction that occurs there each year.

Global Total Ozone Change
TWENTY QUESTIONS: 2006 UPDATE

196

this region of the lower stratosphere, air has only recently

(less than 18 months) been transported from the lower

atmosphere. As a result, the conversion of halogen source

gases to reactive halogen gases is very small. Because of

the low abundance of reactive gases, total ozone depletion in this region is very small. In contrast, stratospheric

air in polar regions has been in the stratosphere for an

average of 4 to 7 years; therefore, the abundance of reactive halogen gases is much larger.

Seasonal changes. The magnitude of global ozone

depletion also depends on the season of the year. In comparison with the 1964-1980 averages, total ozone averaged for 2002-2005 is about 3% lower in northern middle

latitudes (35°N-60°N) and about 6% lower at southern

middle latitudes (35°S-60°S). The seasonality of these

changes is also somewhat different in the two hemispheres. In the summer/autumn periods, the decline in

total ozone is about 2% in the Northern Hemisphere and

5% in the Southern Hemisphere. In winter/spring, the

decline is about 5-6% in both hemispheres.
TWENTY QUESTIONS: 2006 UPDATE

197

Changes in solar radiation and increases in stratospheric particles from volcanic eruptions both affect the

abundance of stratospheric ozone, but they have not

caused the long-term decreases observed in total ozone.

Solar changes. The formation of stratospheric ozone

is initiated by ultraviolet (UV) radiation coming from the

Sun (see Figure Q2-1). As a result, an increase in the Sun’s

radiation output increases the amount of ozone in Earth’s

atmosphere. The Sun’s radiation output and sunspot

number vary over the well-known 11-year solar cycle.

Observations over several solar cycles (since the 1960s)

show that global total ozone levels vary by 1 to 2%

between the maximum and minimum of a typical cycle.

Changes in solar output at a wavelength of 10.7 cm,

although much larger than changes in total solar output,

are often used to show when periods of maximum and

minimum total output occur (see Figure Q14-1). The

Sun’s output has gone through maximum values around

1969, 1980, 1991, and 2002. In 2006, the solar output

was decreasing towards a minimum.

Figure Q14-1. Solar changes and volcanoes. Total

ozone values have decreased beginning in the early

1980s (see middle panel). The ozone values shown are

3-month averages corrected for seasonal effects.

Incoming solar radiation, which produces ozone in the

stratosphere, changes on a well-recognized 11-year

cycle. The amount of solar radiation at a wavelength of

10.7-cm is often used to document the 11-year cycle (see

top panel). A comparison of the top and middle panels

indicates that the cyclic changes in solar output cannot

account for the long-term decreases in ozone. Volcanic

eruptions occurred frequently in the 1965 to 2005 period.

The largest recent eruptions are El Chichón (1982) and

Mt. Pinatubo (1991) (see red arrows in bottom panel).

Large volcanic eruptions are monitored by the decreases

in solar transmission to Earthʼs surface that occur

because new particles are formed in the stratosphere

from volcanic sulfur emissions (see bottom panel). These

particles increase ozone depletion only temporarily

because they do not remain in the stratosphere for more

than a few years. A comparison of the middle and bottom

panels indicates that large volcanic eruptions also cannot

account for the long-term decreases found in global total

ozone.

Q14: Do changes in the Sun and volcanic eruptions affect the ozone layer?

Yes, factors such as changes in solar radiation, as well as the formation of stratospheric particles after volcanic eruptions, do influence the ozone layer. However, neither factor can explain the average decreases

observed in global total ozone over the last two decades. If large volcanic eruptions occur in the coming

decades, ozone depletion will increase for several years after the eruption.

The Solar Cycle, Global Ozone,

and Volcanic Eruptions
TWENTY QUESTIONS: 2006 UPDATE

198

Over the last two decades, average total ozone has

decreased over the globe. Average values in recent years

show about 4% depletion from pre-1980 values (see

Figure Q14-1). The ozone values shown are 3-month

averages corrected for seasonal effects but not for solar

effects. Over the same period, changes in solar output

show the expected 11-year cycle but do not show a

decrease with time. For this reason, the long-term

decreases in global ozone cannot result from changes in

solar output alone. Most examinations of long-term ozone

changes presented in this and previous international scientific assessments quantitatively account for the influence of the 11-year solar cycle.

Past volcanoes. Large volcanic eruptions inject

sulfur gases directly into the stratosphere, causing new

sulfate particles to be formed. The particles initially form

in the stratosphere above and downwind of the volcano

location and then often spread throughout the hemisphere

or globally as air is transported by stratospheric winds.

The presence of volcanic particles in the stratosphere is

shown in observations of solar transmission through the

atmosphere. When large amounts of particles are present

in the stratosphere, transmission of solar radiation is significantly reduced. The large eruptions of El Chichón

(1982) and Mt. Pinatubo (1991) are recent examples of

events that temporarily reduced solar transmission (see

Figure Q14-1).

Laboratory measurements and stratospheric observations have shown that chemical reactions on the surface of volcanically produced particles increase ozone

destruction by increasing the amounts of the highly reactive chlorine gas, chlorine monoxide (ClO). The amount

of ClO produced is proportional to the total abundance

of reactive chlorine in the stratosphere (see Figure Q16-

1). Ozone depletion increases as a consequence of

increased ClO. The most recent large eruption was that

of Mt. Pinatubo, which resulted in up to a 10-fold increase

in the number of particles available for surface reactions.

Both El Chichón and Mt. Pinatubo increased global ozone

depletion for a few years (see Figure Q14-1). After a few

years, however, the effect of volcanic particles on ozone

is diminished by their gradual removal from the stratosphere by natural air circulation. Because of particle

removal, the two large volcanic eruptions of the last two

decades cannot account for the long-term decreases

observed in ozone over the same period.

Future volcanoes. Observations and atmospheric

models indicate that the record-low ozone levels observed

in 1992-1993 resulted from the large number of particles

produced by the Mt. Pinatubo eruption, combined with

the relatively large amounts of reactive halogen gases

present in the stratosphere in the 1990s. If the Mt.

Pinatubo eruption had occurred before 1980, changes to

global ozone would have been much smaller than

observed in 1992-1993 because the abundances of reactive halogen gases in the stratosphere were smaller. In

the early decades of the 21st century, the abundance of

halogen source gases will still be substantial in the global

atmosphere (see Figure Q16-1). If large volcanic eruptions occur in these early decades, ozone depletion will

increase for several years. If an eruption larger than Mt.

Pinatubo occurs, ozone losses could be larger than previously observed and persist longer. Only later in the 21st

century when halogen gas abundances have declined close

to pre-1980 values will the effect of volcanic eruptions on

ozone be lessened.
TWENTY QUESTIONS: 2006 UPDATE

199

Montreal Protocol. In 1985, a treaty called the

Vienna Convention for the Protection of the Ozone Layer

was signed by 20 nations in Vienna. The signing nations

agreed to take appropriate measures to protect the ozone

layer from human activities. The Vienna Convention supported research, exchange of information, and future protocols. In response to growing concern, the Montreal

Protocol on Substances that Deplete the Ozone Layer was

signed in 1987 and, following country ratification, entered

into force in 1989. The Protocol established legally

binding controls for developed and developing nations on

the production and consumption of halogen source gases

known to cause ozone depletion. National consumption

of a halogen gas is defined as the amount that production

and imports of a gas exceed its export to other nations.

Amendments and Adjustments. As the scientific

basis of ozone depletion became more certain after 1987

and substitutes and alternatives became available to

replace the principal halogen source gases, the Montreal

Protocol was strengthened with Amendments and

Adjustments. These revisions put additional substances

under regulation, accelerated existing control measures,

and prescribed phaseout dates for the production and conIV. CONTROLLING OZONE-DEPLETING GASES

Q15: Are there regulations on the production of ozone-depleting gases?

Yes, the production of ozone-depleting gases is regulated under a 1987 international agreement known as the

“Montreal Protocol on Substances that Deplete the Ozone Layer” and its subsequent Amendments and

Adjustments. The Protocol, now ratified by over 190 nations, establishes legally binding controls on the

national production and consumption of ozone-depleting gases. Production and consumption of all principal

halogen-containing gases by developed and developing nations will be significantly phased out before the

middle of the 21st century.

Figure Q15-1. Effect of the Montreal Protocol. The

purpose of the Montreal Protocol is to achieve reductions

in stratospheric abundances of chlorine and bromine.

The reductions follow from restrictions on the production

and consumption of manufactured halogen source

gases. Projections of the future abundance of effective

stratospheric chlorine (see Q16) are shown in the top

panel assuming (1) no Protocol regulations, (2) only the

regulations in the original 1987 Montreal Protocol, and

(3) additional regulations from the subsequent Amendments and Adjustments. The city names and years indicate where and when changes to the original 1987

Protocol provisions were agreed upon. Effective stratospheric chlorine as used here accounts for the combined

effect of chlorine and bromine gases. Without the

Protocol, stratospheric halogen gases are projected to

increase significantly in the 21st century. The “zero emissions” line shows a hypothetical case of stratospheric

abundances if all emissions were reduced to zero beginning in 2007. The lower panel shows how excess skin

cancer cases (see Q17) might increase with no regulation and how they might be reduced under the Protocol

provisions. (The unit “parts per trillion” is defined in the

caption of Figure Q7-1.)

Effect of the Montreal Protocol
TWENTY QUESTIONS: 2006 UPDATE

200

sumption of certain gases. The initial Protocol called for

only a slowing of chlorofluorocarbon (CFC) and halon

production. The 1990 London Amendments to the

Protocol called for a phaseout of the production and consumption of the most damaging ozone-depleting substances in developed nations by 2000 and in developing

nations by 2010. The 1992 Copenhagen Amendments

accelerated the date of the phaseout to 1996 in developed

nations. Further controls on ozone-depleting substances

were agreed upon in later meetings in Vienna (1995),

Montreal (1997), and Beijing (1999).

Montreal Protocol projections. Future stratospheric

abundances of effective stratospheric chlorine (see Q16)

can be calculated based on the provisions of the Montreal

Protocol. The concept of effective stratospheric chlorine

accounts for the combined effect on ozone of chlorineand bromine-containing gases. The results are shown in

Figure Q15-1 for the following cases:

• No Protocol and continued production increases of

3% per year (business-as-usual scenario).

• Continued production and consumption as allowed

by the Protocol’s original provisions agreed upon

in Montreal in 1987.

• Restricted production and consumption as outlined

in the subsequent Amendments and Adjustments

as decided in London in 1990, Copenhagen in 1992,

and Beijing in 1999.

• Zero emissions of ozone-depleting gases starting

in 2007.

In each case, production of a gas is assumed to result

in its eventual emission to the atmosphere. Without the

Montreal Protocol and with continued production and use

of CFCs and other ozone-depleting gases, effective stratospheric chlorine is projected to have increased tenfold by

the mid-2050s compared with the 1980 value. Such high

values likely would have increased global ozone depletion far beyond that currently observed. As a result,

harmful UV-B radiation would have also increased substantially at Earth’s surface, causing a rise in excess skin

cancer cases (see Q17 and lower panel of Figure Q15-1).

The 1987 provisions of the Montreal Protocol alone

would have only slowed the approach to high effective

chlorine values by one or more decades in the 21st century.

Not until the 1992 Copenhagen Amendments and

Adjustments did the Protocol projections show a decrease

in future effective stratospheric chlorine values. Now, with

full compliance to the Montreal Protocol and its

Amendments and Adjustments, use of the major humanproduced ozone-depleting gases will ultimately be phased

out and effective stratospheric chlorine will slowly decay,

reaching pre-1980 values in the mid-21st century (see Q16).

Zero emissions. Effective chlorine values in the

coming decades will be influenced by emissions of

halogen source gases produced in those decades, as well

as the emission of currently existing gases that are now

being used or stored in various ways. Examples of longterm storage are CFCs in refrigeration equipment and

foams, and halons in fire-fighting equipment. Some continued production and consumption of ozone-depleting

gases is allowed, particularly in developing nations, under

the agreements. As a measure of the contribution of these

continued emissions to the effective chlorine value, the

“zero emissions” case is included in Figure Q15-1. In this

hypothetical case, all emissions of ozone-depleting gases

are set to zero beginning in 2007. The reductions in effective stratospheric chlorine below the values expected with

the 1999 Beijing agreement would be relatively small.

HCFC substitute gases. The Montreal Protocol provides for the transitional use of hydrochlorofluorocarbons

(HCFCs) as substitute compounds for principal halogen

source gases such as CFC-12. HCFCs differ chemically

from most other halogen source gases in that they contain

hydrogen (H) atoms in addition to chlorine and fluorine

atoms. HCFCs are used for refrigeration, for blowing

foams, and as solvents, which were primary uses of CFCs.

HCFCs are 88 to 98% less effective than CFC-12 in

depleting stratospheric ozone because they are chemically

removed primarily in the troposphere (see Q18). This

removal partially protects stratospheric ozone from the

halogens contained in HCFCs. In contrast, CFCs and

many other halogen source gases are chemically inert in

the troposphere and, hence, reach the stratosphere without

being significantly removed. Because HCFCs still contribute to the chlorine abundance in the stratosphere, the

Montreal Protocol requires a gradual phaseout of HCFC

consumption in developed and developing nations that

will be complete in 2040.

HFC substitute gases. Hydrofluorocarbons (HFCs)

are also used as substitute compounds for CFCs and other

halogen source gases. HFCs contain only hydrogen, fluorine, and carbon atoms. Because HFCs contain no chlorine or bromine, they do not contribute to ozone depletion

(see Q18). As a consequence, the Montreal Protocol does

not regulate the HFCs. However, HFCs (as well as all

halogen source gases) are radiatively active gases that

contribute to human-induced climate change as they accumulate in the atmosphere (see Q18). HFCs are included

in the group of gases listed in the Kyoto Protocol of the

United Nations Framework Convention on Climate

Change (UNFCCC).
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201

Effective stratospheric chlorine. The Montreal

Protocol has been successful in slowing and reversing the

increase of ozone-depleting gases (halogen source gases)

in the atmosphere. An important measure of its success is

the change in the value of effective stratospheric chlorine.

Effective stratospheric chlorine values are a measure of

the potential for ozone depletion in the stratosphere,

obtained by summing over adjusted amounts of all chlorine and bromine gases. The adjustments account for the

different rates of decomposition of the gases and the

greater per-atom effectiveness of bromine in depleting

ozone (see Q7). Although chlorine is much more abundant in the stratosphere than bromine (160 times) (see

Figure Q7-1), bromine atoms are about 60 times more

effective than chlorine atoms in chemically destroying

ozone molecules. Increases in effective stratospheric

chlorine in the past decades have caused ozone depletion.

Accordingly, ozone is expected to recover in the future as

effective stratospheric chlorine values decrease.

Effective stratospheric chlorine changes. In the

latter half of the 20th century up until the 1990s, effective

stratospheric chlorine values steadily increased (see

Figure Q16-1). Values are derived from individual

halogen source gas abundances obtained from measurements, historical estimates of abundance, and projections

of future abundance. As a result of the Montreal Protocol

regulations, the long-term increase in effective stratospheric chlorine slowed, reached a peak, and began to

decrease in the 1990s. This initial decrease means that

the potential for stratospheric ozone depletion has begun

to lessen as a result of the Montreal Protocol. The decrease

in effective chlorine is projected to continue throughout

the 21st century if all nations continue to comply with the

provisions of the Protocol. The decrease will continue

because, as emissions become small, natural destruction

processes gradually remove halogen-containing gases

from the global atmosphere. Reduction of effective stratospheric chlorine amounts to 1980 values or lower will

require many decades because the lifetimes of halogen

source gas molecules in the atmosphere range up to 100

years (see Figure Q16-1 and Table Q7-1).

Individual halogen source gas reductions. The

reduction in the atmospheric abundance of a gas in

response to regulation depends on a number of factors that

include (1) how rapidly gas reserves are used and released

to the atmosphere, (2) the lifetime for the removal of the

gas from the atmosphere, and (3) the total amount of the

gas that has already accumulated in the atmosphere.

The regulation of human-produced halogen source

gases under the Montreal Protocol is considered separately for each class of one or more gases and is based on

several factors. The factors include (1) the effectiveness

of each class in depleting ozone in comparison with other

halogen source gases, (2) the availability of suitable substitute gases for domestic and industrial use, and (3) the

impact of regulation on developing nations.

Methyl chloroform and CFCs. The largest reduction in the abundance of a halogen source gas has

occurred for methyl chloroform (CH3CCl3) (see Figure

Q16-1). The implementation of the Montreal Protocol

caused global production of methyl chloroform to be

reduced to near zero. Atmospheric abundances subsequently dropped rapidly because methyl chloroform has

a short atmospheric lifetime (about 5 years). Methyl chloroform is used mainly as a solvent and has no significant

long-term storage following production. The reduction

in effective chlorine in the 1990s came primarily from

the reduction in methyl chloroform abundance in the

atmosphere. Significant emissions reductions have also

occurred for the chlorofluorocarbons CFC-11, CFC-12,

and CFC-113 starting in the 1990s. As a result, the atmospheric amounts of these gases have all peaked, and CFC11 and CFC-113 abundances have decreased slightly (see

Figure Q16-1). As emissions of CFCs are reduced, their

atmospheric abundances will decrease more slowly than

methyl chloroform because of longer CFC atmospheric

lifetimes (see Table Q7-1) and because CFCs escape very

slowly to the atmosphere from their use in refrigeration

Q16: Has the Montreal Protocol been successful in reducing ozone-depleting gases in the

atmosphere?

Yes, as a result of the Montreal Protocol, the total abundance of ozone-depleting gases in the atmosphere

has begun to decrease in recent years. If the nations of the world continue to follow the provisions of the

Montreal Protocol, the decrease will continue throughout the 21st century. Some individual gases, such as

halons and hydrochlorofluorocarbons (HCFCs), are still increasing in the atmosphere but will begin to

decrease in the next decades if compliance with the Protocol continues. Around midcentury, the effective

abundance of ozone-depleting gases should fall to values that were present before the Antarctic “ozone

hole” began to form in the early 1980s.
TWENTY QUESTIONS: 2006 UPDATE

202

Figure Q16-1. Halogen source gas changes. The rise in effective stratospheric chlorine values in the 20th century has

slowed and reversed in the last decade (top left panel). Effective stratospheric chlorine values are a measure of the

potential for ozone depletion in the stratosphere, obtained by summing over adjusted amounts of all chlorine and bromine

gases. Effective stratospheric chlorine levels as shown here for midlatitudes will return to 1980 values around 2050. The

return to 1980 values will occur around 2065 in polar regions. In 1980, ozone was not significantly depleted by the chlorine and bromine then present in the stratosphere. A decrease in effective stratospheric chlorine abundance follows

reductions in emissions of individual halogen source gases. Overall emissions and atmospheric concentrations have

decreased and will continue to decrease given international compliance with the Montreal Protocol provisions (see Q15).

The changes in the atmospheric abundance of individual gases at Earthʼs surface shown in the panels were obtained

using a combination of direct atmospheric measurements, estimates of historical abundance, and future projections of

abundance. The past increases of CFCs, along with those of CCl4 and CH3CCl3, have slowed significantly and most have

reversed in the last decade. HCFCs, which are used as CFC substitutes, will continue to increase in the coming decades.

Some halon abundances will also continue to grow in the future while current halon reserves are depleted. Smaller relative decreases are expected for CH3Br in response to production and use restrictions because it has substantial natural

sources. CH3Cl has large natural sources and is not regulated under the Montreal Protocol. (See Figure Q7-1 for chemical names and formulas. The unit “parts per trillion” is defined in the caption of Figure Q7-1.)

Past and Expected Future Abundances of Atmospheric Halogen Source Gases
TWENTY QUESTIONS: 2006 UPDATE

203

and foam products.

HCFC substitute gases. The Montreal Protocol

allows for the use of hydrochlorofluorocarbons (HCFCs)

as short-term substitutes for CFCs. As a result, the abundances of HCFC-22, HCFC-141b, and HCFC-142b continue to grow in the atmosphere (see Figure Q16-1).

HCFCs pose a lesser threat to the ozone layer than CFCs

because they are partially destroyed in the troposphere by

chemical processes, thus reducing the overall effectiveness of their emissions in destroying stratospheric ozone.

Under the Montreal Protocol, HCFC consumption will

reach zero in developed nations by 2030 and in developing nations by 2040 (see Q15). Thus, the future projections in Figure Q16-1 show HCFC abundances

reaching a peak in the first decades of the 21st century and

steadily decreasing thereafter.

Halons. The atmospheric abundances of halon-1211

and halon-1301 account for a significant fraction of

bromine from all source gases (see Figure Q7-1) and continue to grow despite the elimination of production in

developed nations in 1994 (see Figure Q16-1). The

growth in abundance continues because substantial

reserves are held in fire-extinguishing equipment and are

gradually being released, and production and consumption are still allowed in developing nations. Atmospheric

halon abundances can be expected to remain high well

into the 21st century because of their long lifetimes and

continued release.

Methyl chloride and methyl bromide. Both

methyl chloride (CH3Cl) and methyl bromide (CH3Br)

are distinct among principal halogen source gases

because a substantial fraction of their emissions is associated with natural processes (see Q7). The average

atmospheric abundance of methyl chloride, which is not

regulated under the Montreal Protocol, will remain fairly

constant throughout this century if natural sources remain

unchanged. At century’s end, methyl chloride is expected

to account for a large fraction of remaining effective

stratospheric chlorine because the abundances of other

gases, such as the CFCs, are expected to be greatly

reduced (see Figure Q16-1). The abundance of methyl

bromide, which is regulated under the Protocol, has

already decreased in recent years and is projected to

decrease further as a result of production phaseouts in

developed and developing countries. For the later decades

of the century, methyl bromide abundances are shown as

nearly constant in Figure Q16-1. However, these abundances are uncertain because the amounts of exempted

uses of methyl bromide under the Montreal Protocol are

not known for future years.
TWENTY QUESTIONS: 2006 UPDATE

204

The depletion of stratospheric ozone leads to an

increase in surface ultraviolet radiation. The increase

occurs primarily in the ultraviolet-B (UV-B) component

of the Sun’s radiation. UV-B is defined as radiation in the

wavelength range of 280 to 315 nanometers. Changes in

UV-B at the surface have been observed directly and can

be estimated from ozone changes.

Surface UV-B radiation. The amount of ultraviolet

radiation reaching Earth’s surface depends in large part

on the amount of ozone in the atmosphere. Ozone molecules in the stratosphere absorb UV-B radiation, thereby

significantly reducing the amount of this radiation that

reaches Earth’s surface (see Q3). If total ozone amounts

are reduced in the stratosphere, then the amount of UV

radiation reaching Earth’s surface generally increases.

This relationship between total ozone and surface UV

radiation has been studied at a variety of locations with

direct measurements of both ozone and UV. The actual

amount of UV reaching a location depends on a large

number of additional factors, including the position of

the Sun in the sky, cloudiness, and air pollution. In general, surface UV at a particular location on Earth changes

throughout the day and with season as the Sun’s position

in the sky changes.

Long-term surface UV changes. Satellite observations of long-term global ozone changes can be used

to estimate changes in global surface UV that have

occurred over the past two decades. These changes are

of interest because UV radiation can cause harm to

humans, other life forms, and materials (see Q3). The

amount of UV that produces an “erythemal” or sunburning

response in humans is often separately evaluated. Longterm changes in sunburning UV at a particular location

have been estimated from the changes in total ozone at

that location. The results show that average erythemal

UV has increased due to ozone reduction by up to a few

percent per decade between 1979 and 1998 over a wide

range of latitudes (see Figure Q17-1). The largest

increases are found at high polar latitudes in both hemispheres. As expected, the increases occur where

decreases in total ozone are observed to be the largest

(see Figure Q13-1). The smallest changes in erythemal

UV are in the tropics, where long-term total ozone

changes are smallest.

UV Index changes. The “UV Index” is a measure

of daily surface UV levels that is relevant to the effects

of UV on human skin. The UV Index is used internationally to increase public awareness about the detrimental

effects of UV on human health and to guide the need for

protective measures. The UV Index is essentially the

amount of erythemal irradiance as measured on a horizontal surface. The daily maximum UV Index varies with

location and season, as shown for three locations in

Figure Q17-1. Changes in surface UV radiation.

Ultraviolet (UV) radiation at Earthʼs surface has increased

over much of the globe since 1979. Also known as “erythemal radiation,” sunburning UV is harmful to humans

and other life forms. The increases shown here for 1979-

1998 are estimated from observed decreases in ozone

and the relationship between ozone and surface UV

established at some surface locations. The estimates

are based on the assumption that all other factors that

influence the amount of UV radiation reaching the Earthʼs

surface, such as aerosol abundances and cloudiness,

are unchanged. The estimated changes in ultraviolet

radiation in the tropics are the smallest because observed

ozone changes are the smallest there.

V. IMPLICATIONS OF OZONE DEPLETION

Q17: Does depletion of the ozone layer increase ground-level ultraviolet radiation?

Yes, ultraviolet radiation at Earth’s surface increases as the amount of overhead total ozone decreases,

because ozone absorbs ultraviolet radiation from the Sun. Measurements by ground-based instruments and

estimates made using satellite data have confirmed that surface ultraviolet radiation has increased in regions

where ozone depletion is observed.

Changes in Surface Ultraviolet Radiation
TWENTY QUESTIONS: 2006 UPDATE

205

Figure Q17-2. The highest daily values generally occur

at the lowest latitudes (tropics) and in summer when the

midday Sun is closest to overhead. Values in San Diego,

California, for example, normally are larger year round

than those found in Barrow, Alaska, which is at higher

latitude. At a given latitude, UV Index values increase in

mountainous regions. The UV Index becomes zero in

periods of continuous darkness found during winter at

high-latitude locations.

An illustrative example of how polar ozone depletion increases the maximum daily UV Index is shown

in Figure Q17-2. Normal UV Index values for Palmer,

Antarctica, in spring were estimated from satellite

measurements made during the period 1978-1983,

before the appearance of the “ozone hole” over

Antarctica (see red dotted line). In the last decade

(1991-2001), severe and persistent ozone depletion in

spring has increased the UV Index well above normal

values for several months (see thick red line). Now,

spring UV Index values in Palmer, Antarctica (64°S),

sometimes equal or exceed even the peak summer

values measured in San Diego, California (32°N).

Other causes of long-term UV changes. The surface UV values may also change as a result of other

human activities or climate change. Long-term changes

in cloudiness, aerosols, pollution, and snow or ice cover

will cause long-term changes in surface UV. At some

ground sites, measurements indicate that long-term

changes in UV have resulted from changes in one or more

of these factors. The impact of some of the changes can

be complex. For example, an increase in cloud cover

usually results in a reduction of UV radiation below the

clouds, but can increase radiation above the clouds (in

mountainous regions).

UV changes and skin cancer. Skin cancer cases in

humans are expected to increase with the amount of UV

reaching Earth’s surface. Atmospheric scientists working

together with health professionals can estimate how skin

cancer cases will change in the future. The estimates are

based on knowing how UV increases as total ozone is

depleted and how total ozone depletion changes with

effective stratospheric chlorine (see Q16). Estimates of

future excess skin cancer cases are shown in Figure Q15-

1 using future estimates of effective stratospheric chlorine based on the 1992 and earlier Montreal Protocol

provisions and assuming that other factors (besides

Figure Q17-2. Changes in UV Index. The

maximum daily UV Index is a measure of

peak sunburning UV that occurs during

the day at a particular location. UV-B,

which is absorbed by ozone, is an important component of sunburning UV. The

UV Index varies with latitude and season,

and with the Sunʼs elevation in the local

sky. The highest values of the maximum

daily UV Index occur in the tropics, where

the midday Sun is highest throughout the

year and where total ozone values are

lowest. The lowest average UV Index

values occur at high latitudes. As an

example, the figure compares the seasonal UV Index at three locations. The

UV Index is higher throughout the year

in San Diego, a low-latitude location, than

in Barrow, a high-latitude location. Index

values are zero at high latitudes in winter

when darkness is continuous. The effect

of Antarctic ozone depletion is demonstrated by comparing the Palmer and San Diego data in the figure. Normal

values estimated for Palmer are shown for the 1978-1983 period before the “ozone hole” occurred each season (see

red dotted line). In the decade 1991-2001, Antarctic ozone depletion has increased the maximum UV Index value at

Palmer throughout spring (see yellow shaded region). Values at Palmer now sometimes equal or exceed those measured in spring and even in the summer in San Diego, which is located at much lower latitude.

Seasonal Changes in the UV Index
TWENTY QUESTIONS: 2006 UPDATE

206

ozone) affecting surface UV are unchanged in the future.

The cases are those that would occur in a population

with the UV sensitivity and age distribution such as that

of the United States. The cases counted are those in

excess of the number that occurred in 1980 before ozone

depletion was observed (about 2000 per million population), with the assumption that the population’s sun exposure remains unchanged. The case estimates include the

fact that skin cancer in humans occurs long after the

exposure to sunburning UV. The results illustrate that,

with current Protocol provisions, excess skin cancer

cases are predicted to increase in the early to middle

decades of the 21st century. By century’s end, with the

expected decreases in halogen source gas emissions, the

number of excess cases is predicted to return close to

1980 values. Without the provisions of the Protocol,

excess skin cancer cases would have been expected to

increase substantially throughout the century.
TWENTY QUESTIONS: 2006 UPDATE

207

Radiative forcing of climate change. Human activities and natural processes have led to the accumulation

in the atmosphere of several long-lived and radiatively

active gases known as “greenhouse gases.” Ozone is a

greenhouse gas, along with carbon dioxide (CO2),

methane (CH4), nitrous oxide (N2O), and halogen source

gases. The accumulation of these gases in Earth’s atmosphere changes the balance between incoming solar radiation and outgoing infrared radiation. Greenhouse gases

generally change the balance by absorbing outgoing radiation, leading to a warming at Earth’s surface. This

change in Earth’s radiative balance is called a radiative

forcing of climate change.

A summary of radiative forcings resulting from the

increases in long-lived greenhouse gases in the industrial

era is shown in Figure Q18-1. All forcings shown relate

to human activities. Positive forcings generally lead to

warming and negative forcings lead to cooling of Earth’s

surface. The accumulation of carbon dioxide represents

the largest forcing term. Carbon dioxide concentrations

are increasing in the atmosphere primarily as the result of

burning coal, oil, and natural gas for energy and transportation; and from cement manufacturing. The atmospheric abundance of carbon dioxide is currently about

35% above what it was 250 years ago, in preindustrial

times. In other international assessments, much of the

observed surface warming over the last 50 years has been

linked to increases in carbon dioxide and other greenhouse

gas concentrations caused by human activities.

Stratospheric and tropospheric ozone. Stratospheric and tropospheric ozone both absorb infrared radiation emitted by Earth’s surface, effectively trapping heat

in the atmosphere. Stratospheric ozone also significantly

absorbs solar radiation. As a result, increases or decreases

Q18: Is depletion of the ozone layer the principal cause of climate change?

No, ozone depletion itself is not the principal cause of climate change. However, because ozone absorbs solar

radiation and is a greenhouse gas, ozone changes and climate change are linked in important ways. Stratospheric ozone depletion and increases in global tropospheric ozone that have occurred in recent decades both

contribute to climate change. These contributions to climate change are significant but small compared with

the total contribution from all other greenhouse gases. Ozone and climate change are indirectly linked

because both ozone-depleting gases and substitute gases contribute to climate change.

Figure Q18-1. Radiative forcing of climate

change from atmospheric gas changes.

Human activities since the start of the

Industrial Era (around 1750) have caused

increases in the abundances of several longlived gases, changing the radiative balance

of Earthʼs atmosphere. These gases, known

as “greenhouse gases,” result in radiative

forcings, which can lead to climate change.

Other international assessments have

shown that the largest radiative forcings

come from carbon dioxide, followed by

methane, tropospheric ozone, the halogencontaining gases (see Figure Q7-1), and

nitrous oxide. Ozone increases in the troposphere result from pollution associated with

human activities. All these forcings are positive, which leads to a warming of Earthʼs surface. In contrast, stratospheric ozone depletion represents a small negative forcing, which leads to cooling of Earthʼs

surface. In the coming decades, halogen gas abundances and stratospheric ozone depletion are expected to be

reduced along with their associated radiative forcings. The link between these two forcing terms is an important aspect

of the radiative forcing of climate change.

Radiative Forcing of Climate Change from

Atmospheric Gas Changes (1750-2000)
TWENTY QUESTIONS: 2006 UPDATE

208

in stratospheric or tropospheric ozone cause radiative forcings and represent direct links of ozone to climate change.

In recent decades, stratospheric ozone has decreased due

to rising chlorine and bromine amounts in the atmosphere,

while troposphere ozone in the industrial era has increased

due to pollution from human activities (see Q3).

Stratospheric ozone depletion causes a negative radiative

forcing, while increases in tropospheric ozone cause a

positive radiative forcing (see Figure Q18-1). The radiative forcing due to tropospheric ozone increases is currently larger than that associated with stratospheric ozone

depletion. The negative forcing from ozone depletion represents an offset to the positive forcing from the halogen

source gases, which cause ozone depletion.

Halogen source gases and HFCs. An important link

between ozone depletion and climate change is the radiative forcing from halogen source gases and hydrofluorocarbons (HFCs). Halogen source gases are the cause of

ozone depletion (see Q7) and HFCs are substitute gases

(see Q15). Both groups of gases cause radiative forcing

in the atmosphere, but with a wide range of effectiveness.

The principal gases in each group are intercompared in

Figure Q18-2 (top panel) using their “ozone depletion

potentials” (ODPs) and “global warming potentials”

(GWPs), which indicate the effectiveness of each gas in

causing ozone depletion and climate change, respectively.

The ODPs of CFC-11 and CFC-12, and the GWP of CO2

all are assigned a value of 1.0. For ozone depletion, the

halons are the most effective gases (for equal mass

amounts) and HFCs cause no ozone depletion (see Q7).

For climate change, all gases make a contribution, with

CFC-12 and HFC-23 having the largest effect (for equal

mass amounts). Montreal Protocol actions (see Q15) that

have led to reductions in CFC concentrations and

increases in HCFC and HFC concentrations have also

reduced the total radiative forcing from these gases. It is

important to note that, despite a GWP that is small in comparison to many other greenhouse gases, CO2 is the most

important greenhouse gas related to human activities

because its atmospheric abundance is so much greater than

the abundance of other gases.

The relative importance of total emissions of halogen

source gases and HFCs to ozone depletion and climate

change is illustrated for a single year (2004) of emissions

in the bottom panel of Figure Q18-2. The values displayed

are proportional to the product of 2004 annual global

emissions and the ODP or GWP. The results in the lower

panel are shown relative to CFC-11, because it is often

used as a reference gas. The comparison shows that the

importance of CFC emissions in 2004 to future ozone

depletion exceeds that of the halons, despite the higher

halon ODP values, because CFC emissions are larger.

Similarly, the contributions of CFC and HCFC-22 emissions in 2004 to climate change are currently larger than

the halon or HFC contributions. These 2004 results represent only incremental contributions of these gases to

either ozone depletion or climate change. The overall

contribution of a gas depends on its total accumulation

in the atmosphere, which in turn depends on its longterm emission history and atmospheric lifetime (see Q7

and Q16). In the case of ozone depletion, the relative

contributions of the halogen source gases can be compared through their respective contributions to effective

stratospheric chlorine (see Q16).

As a group, the principal halogen source gases represent a positive direct radiative forcing in the Industrial

Era that is comparable to the forcing from methane, the

second most important greenhouse gas. In the coming

decades, the abundances of these ozone-depleting gases

and their associated positive radiative forcings are

expected to decrease (see Q16). Future growth in HFC

emissions, while uncertain, will contribute a positive

forcing that will counter the decrease from ozonedepleting gases. Finally, reductions in ozone-depleting

gases will be followed by reductions in stratospheric

ozone depletion and its associated negative radiative

forcing.

Impact of climate change on ozone. Certain

changes in Earth’s climate could affect the future of the

ozone layer. Stratospheric ozone is influenced by changes

in temperatures and winds in the stratosphere. For

example, lower temperatures and stronger polar winds

could both affect the extent and severity of winter polar

ozone depletion. While the Earth’s surface is expected to

warm in response to the net positive radiative forcing from

greenhouse gas increases, the stratosphere is expected to

cool. A cooler stratosphere would extend the time period

over which polar stratospheric clouds (PSCs) are present

in polar regions and, as a result, might increase winter

ozone depletion. In the upper stratosphere at altitudes

above PSC formation regions, a cooler stratosphere is

expected to increase ozone amounts and, hence, hasten

recovery, because lower temperatures favor ozone production over loss (see Q2). Similarly, changes in atmospheric composition that lead to a warmer climate may also

alter ozone amounts (see Q20).
TWENTY QUESTIONS: 2006 UPDATE

209

Figure Q18-2. Evaluation of ozone-depleting gases and their substitutes. Ozone-depleting gases (halogen source

gases) and their substitutes can be compared via their ozone depletion potentials (ODPs) and global warming potentials

(GWPs). The GWPs are evaluated for a 100-yr time interval after emission. The CFCs, halons, and HCFCs are ozonedepleting gases (see Q7) and HFCs, used as substitute or replacement gases, do not destroy ozone. The ODPs of CFC11 and CFC-12, and the GWP of CO2 have values of 1.0 by definition. Larger ODPs or GWPs indicate greater potential

for ozone depletion or climate change, respectively. The top panel compares ODPs and GWPs for emissions of equal

mass amounts of each gas. The ODPs of the halons far exceed those of the CFCs. HFCs have zero ODPs. All gases

have non-zero GWPs that span a wide range of values. The bottom panel compares the contributions of the 2004

emissions of each gas, using CFC-11 as the reference gas. Each bar represents the product of a global emission

value and the respective ODP or GWP factor. The comparison shows that 2004 emissions of ozone-depleting gases

currently contribute more than substitute gas emissions to both ozone depletion and climate change. Future projections guided by Montreal Protocol provisions suggest that the contributions of ozone-depleting gases to climate change

will decrease, while those of the substitute gases will increase.

Evaluation of Selected Ozone-Depleting Substances and Substitute Gases
TWENTY QUESTIONS: 2006 UPDATE

210

Recovery process. Identifying the recovery of the

ozone layer from depletion associated with halogen gases

will rely on comparisons of the latest ozone values with

values measured in the past. Because of its importance,

ozone will likely be measured continuously in the future

using a variety of techniques and measurement platforms

(see Q5). Atmospheric computer models will be used to

predict future abundances of ozone and attribute observed

changes to ozone-depleting gases and other factors.

The recovery process is schematically shown for

global ozone in Figure Q19-1. Ozone has declined from

pre-1980 amounts due to past increases in halogen gases

in the stratosphere (see Q16). In the future, as the overall

decline in these gases continues in response to Montreal

Protocol provisions, global ozone is expected to recover,

approaching or exceeding pre-1980 values (see Q20).

Ozone recovery attributable to decreases in ozonedepleting gases can be described, in general, as a process

VI. STRATOSPHERIC OZONE IN THE FUTURE

Q19: How will recovery of the ozone layer be identified?

Scientists expect to identify the recovery of the ozone layer with detailed ozone measurements in the atmosphere and with global models of ozone amounts. Increases in global ozone and reductions in the extent and

severity of the Antarctic “ozone hole” will be important factors in gauging ozone recovery. Natural variations

in ozone amounts will limit how soon recovery can be detected with future ozone measurements.

Figure Q19-1. Recovery stages of global ozone. Significant ozone depletion from the release of ozone-depleting

gases in human activities first became recognized in the 1980s. The Montreal Protocol provisions are expected to further

reduce and eliminate these gases in the atmosphere in the coming decades, thereby leading to the return of ozone

amounts to near pre-1980 values. The timeline of the recovery process is schematically illustrated with three stages

identified. The large uncertainty range illustrates natural ozone variability in the past and potential uncertainties in global

model projections of future ozone amounts. When ozone reaches the full recovery stage, global ozone values may be

above or below pre-1980 values, depending on other changes in the atmosphere (see Q20).

Recovery Stages of Global Ozone
TWENTY QUESTIONS: 2006 UPDATE

211

involving three stages:

(1) The initial slowing of ozone decline, identified as

the occurrence of a statistically significant reduction in the rate of decline in ozone.

(2) The onset of ozone increases (turnaround), identified as the occurrence of statistically significant

increases in ozone above previous minimum

values.

(3) The full recovery of ozone from ozone-depleting

gases, identified as when ozone is no longer significantly affected by ozone-depleting gases from

human activities.

Each recovery stage is noted in Figure Q19-1. The red

line and shaded region in the figure indicate the expected

average value and the uncertainty range, respectively, in

global ozone amounts. The large uncertainty range illustrates natural ozone variability in the past and potential

uncertainties in global model projections of future ozone

amounts.

In the full recovery of global ozone, the milestone of

the return of ozone to pre-1980 levels is considered important because prior to 1980 ozone was not significantly

affected by human activities. As a consequence, this milestone is useful, for example, to gauge when the adverse

impacts of enhanced surface ultraviolet (UV) radiation on

human health and ecosystems caused by ozone-depleting

substances are likely to become negligible. The uncertainty range in model results indicates that ozone amounts

may be below or above pre-1980 values when ozone has

fully recovered from the effects of ozone-depleting gases

from human activities (see Q20). The wide range of

uncertainty for global ozone in the final stage of recovery

represents, in part, the difficulty in accurately forecasting

the effects of future changes in climate and atmospheric

composition on the abundance of ozone (see Q20).

Natural factors. Stratospheric ozone is influenced

by two important natural factors, namely, changes in the

output of the Sun and volcanic eruptions (see Q14).

Evaluations of ozone recovery include the effects of these

natural factors. The solar effect on ozone is expected to

be predictable based on the well-established 11-year cycle

of solar output. The uncertainty range in Figure Q19-1

includes solar changes. Volcanic eruptions are particularly important because they enhance ozone depletion

caused by reactive halogen gases, but cannot be predicted.

The occurrence of a large volcanic eruption in the next

decades when effective stratospheric chlorine levels are

still high (see Figure Q16-1) may obscure progress in

overall ozone recovery by temporarily increasing ozone

depletion. The natural variation of ozone amounts also

limits how easily small improvements in ozone abundances can be detected.
TWENTY QUESTIONS: 2006 UPDATE

212

Q20: When is the ozone layer expected to recover?

Halogen source gas reductions. Ozone depletion

caused by human-produced chlorine and bromine gases is

expected to gradually disappear by about the middle of

the 21st century as the abundances of these gases decline

in the stratosphere. The decline in effective stratospheric

chlorine will follow the reductions in emissions that are

expected to continue under the provisions of the Montreal

Protocol and its Adjustments and Amendments (see Figure

Q16-1). The emission reductions are based on the

assumption of full compliance by the developed and

developing nations of the world. The slowing of increases

in atmospheric abundances and the initial decline of several halogen gases have already been observed (see Figure

Q16-1). One gas, methyl chloroform, has already

decreased by about 90% from its peak value. Natural

chemical and transport processes limit the rate at which

halogen gases are removed from the stratosphere. The

atmospheric lifetimes of the halogen source gases range

up to 100 years (see Table Q7-1). Chlorofluorocarbon-12

(CFC-12), with its 100-year lifetime, will require about

200 to 300 years before it is removed (less than 5%

remaining) from the atmosphere (see Figure Q16-1). At

midlatitudes, effective stratospheric chlorine is not

expected to reach pre-1980 values until about 2050.

Ozone projections. Computer models of the atmosphere are used to assess past changes in the global ozone

distribution and to project future changes. Two important

measures of ozone considered by scientists are global total

ozone averaged between 60°N and 60°S latitudes, and

minimum ozone values in the Antarctic “ozone hole.”

Both measures show ongoing ozone depletion that began

in the 1980s (see Figure Q20-1). The model projections

indicate that for 60°N-60°S total ozone, the first two

stages of recovery (slowing of the decline and turnaround

Figure Q20-1. Global ozone recovery predictions.

Observed values of midlatitude total ozone (top panel) and

September-October minimum total ozone values, over

Antarctica (bottom panel) have decreased beginning in the

early 1980s. As halogen source gas emissions decrease in

the 21st century, ozone values are expected to recover by

increasing toward pre-1980 values. Atmospheric computer

models that account for changes in halogen gases and other

atmospheric parameters are used to predict how ozone

amounts will increase. These model results show that full

recovery is expected in midlatitudes by 2050, or perhaps

earlier. Recovery in the Antarctic will occur somewhat later.

The range of model projections comes from the use of several different models of the future atmosphere.

Substantial recovery of the ozone layer is expected near the middle of the 21st century, assuming global compliance with the Montreal Protocol. Recovery will occur as chlorine- and bromine-containing gases that

cause ozone depletion decrease in the coming decades under the provisions of the Protocol. However, the

influence of changes in climate and other atmospheric parameters could accelerate or delay ozone recovery,

and volcanic eruptions in the next decades could temporarily reduce ozone amounts for several years.

Recovery of Global Ozone
TWENTY QUESTIONS: 2006 UPDATE

213

(see Q19)) will be reached before 2020. Full recovery,

with ozone reaching or exceeding pre-1980 values, is

expected to occur by the middle of the 21st century. The

range of projections comes from several computer models

of the atmosphere. Some of these models indicate that

recovery of 60°N-60°S total ozone may come well before

midcentury.

Models predict that Antarctic ozone depletion will

also reach the first two stages of recovery by 2020, but

somewhat more slowly than 60°N-60°S total ozone.

Full recovery could occur by mid-century but some

models show later recovery, between 2060 and 2070.

Declines in effective stratospheric chlorine amounts will

occur later over the Antarctic than at lower latitudes

because air in the Antarctic stratosphere is older than

air found at lower latitudes. As a result, reductions in

halogen loading to pre-1980 values will occur 10-15

years later in the Antarctic stratosphere than in the midlatitude stratosphere.

A different atmosphere in 2050. By the middle of

the 21st century, halogen amounts in the stratosphere are

expected to be similar to those present in 1980 before the

onset of significant ozone depletion (see Figure Q16-1).

However, climate and other atmospheric factors will not

be the same in 2050 as in 1980, and this could cause ozone

abundances in 2050 to be somewhat different from those

observed in 1980. Stratospheric ozone abundances are

affected by a number of natural and human-caused factors in addition to the atmospheric abundance of halogen

gases. Important examples are stratospheric temperatures

and air motions, volcanic eruptions, solar activity, and

changes in atmospheric composition. Separating the

effects of these factors is challenging because of the complexity of atmospheric processes affecting ozone.

The ozone recovery projections in Figures Q19-1 and

Q20-1 attempt to take these various factors into account.

For example, since 1980 human activities have increased

the atmospheric abundance of important greenhouse

gases, including carbon dioxide, methane, and nitrous

oxide. Other international assessments have shown that

the accumulation of these gases is linked to the warmer

surface temperatures and lower stratospheric temperatures observed within recent decades. Warmer surface

temperatures could change the emission rates of naturally

occurring halogen source gases. Lower temperatures in

the upper stratosphere (at about 40 kilometers (25 miles)

altitude) accelerate ozone recovery because ozone

destruction reactions proceed at a slower rate. In contrast,

reduced temperatures in the polar lower stratosphere

during winter might increase the occurrence of polar

stratospheric clouds (PSCs) and, therefore, enhance chemical ozone destruction (see Q10). Further increases of

stratospheric water vapor, such as those that have occurred

over the last two decades, could also increase PSC occurrences and associated ozone destruction. Therefore, a

cooler, wetter polar stratosphere could delay polar ozone

recovery beyond what would be predicted for the 1980

atmosphere. Increased abundances of methane and

nitrous oxide due to human activities also cause some

change in the overall balance of the chemical production

and destruction of global stratospheric ozone. Finally,

one outcome that cannot be included precisely in models

is the occurrence of one or more large volcanic eruptions

in the coming decades. Large eruptions would increase

stratospheric sulfate particles for several years, temporarily reducing global ozone amounts (see Q14).

As a consequence of these potential changes, the

return of effective stratospheric chlorine and ozone to pre1980 levels may not occur at the same time. In some

regions of the stratosphere, ozone may remain below pre1980 values after effective chlorine has declined to pre1980 levels.
214

The U.S. Climate Change Science Program Appendices
215

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

GLOSSARY AND ACRONYMS Glossary

aerosol

A suspension of very fine particles (solid or liquid) 

in air.

anthropogenic

Human-caused.

bank

The amount of a chemical that has been produced but 

has not yet been emitted or chemically altered. They 

exist either in reserve storage or in current applications 

(i.e., refrigerators, air conditioners, fire extinguishers, 

etc.). 

catalytic reaction

Acceleration (increase in rate) of a chemical reaction 

by means of a substance, called a catalyst. Chlorine 

acts as a catalyst in the destruction of ozone in the 

stratosphere.

chlorofluorocarbons (CFCs)

Halocarbons containing only chlorine, fluorine, and 

carbon atoms. CFCs are ozone-depleting substances 

(ODSs).

climate forcing

Changes that affect the energy balance of the planet and 

that consequently “force” the climate to change (see 

also radiative forcing). Examples of climate forcing 

include changes in atmospheric carbon dioxide, or 

suspended particulates (see aerosols), or energy from 

the sun.

CO2-equivalent

The amount of carbon dioxide (CO2) that would cause 

the same amount of radiative forcing as a given amount 

of another greenhouse gas.

column ozone

The total amount of ozone in a vertical column above 

the Earth’s surface. Column ozone is measured in 

Dobson units (DU).

consumption

Used here as defined by the Montreal Protocol: The 

magnitude of an ozone-depleting substance produced 

or imported, minus the amount of that substance that 

is exported.

Equivalent Effective Chlorine (EECl)

An index used to approximately quantify overall changes 

in reactive halogen trends based on the measured mix 

of ODSs in the troposphere. It accounts for the number 

of halogen atoms in the ODS molecule, their relative 

efficiency at destroying ozone, and the rate at which 

the ODSs decompose in the stratosphere to release the 

halogen atoms.

Equivalent Effective Stratospheric Chlorine (EESC)

An index related to EECl that considers the time lags 

associated with transporting air from the troposphere 

to the stratosphere. EESC roughly estimates the effect 

of ODSs on stratospheric ozone. The EESC index is 

used to estimate the time evolution of ozone-depleting 

halogen (chlorine, bromine) in the atmosphere. It is often 

used to estimate when the cumulative effect of all ODSs 

on ozone will return to a level attained at some earlier 

time, or to evaluate the relative effects of potential policy 

scenarios.

Global Warming Potential

An index comparing the cumulative radiative forcing 

(i.e., climate impact) of a unit mass of a greenhouse gas 

relative to the same quantity of a unit mass of a reference 

gas (usually CO2) over some time period (usually 100 

years).

100-year GWP

The Global Warming Potential of a chemical integrated 

over a 100-year time horizon relative to CO2. When 

applied as a weighting factor to emissions or production 

of other chemicals, the resulting quantity provides a CO2-

equivalent emission or production. 

greenhouse gases

Gases including water vapor, carbon dioxide, methane, 

nitrous oxide, and halocarbons that trap infrared heat, 

warming the air near the surface and in the lower levels 

of the atmosphere.

halocarbons

Chemical compounds containing carbon atoms, and 

one or more atoms of the halogens chlorine (Cl), 

bromine (Br), or iodine (I). Halocarbons include 

chlorofluorocarbons (CFCs), hydrochlorofluorocarbons 

(HCFCs), hydrofluorocarbons (HFCs), perfluorocarbons 

(PFCs), and halons.
The U.S. Climate Change Science Program Glossary and Acronyms

216

halogen

A member of the family of elements that includes fluorine, 

chlorine, bromine, and iodine. In the context of ozone 

depletion, chlorine (Cl), bromine (Br), and iodine (I) are the 

halogens of greatest relevance. “Halogenated” compounds 

are chemicals that contain one or more halogen atoms in 

their molecular structure. Fully halogenated halocarbons 

contain only carbon and halogen atoms, whereas partially 

halogenated halocarbons also contain hydrogen (H) 

atoms.

halon

Fully halogenated halocarbons that contain bromine 

and f luorine. Halons are used in fire-extinguishing 

applications.

mixing ratio (by mass)

The ratio of the mass of a specific substance in a volume of 

air to the total mass of that volume of air.

ozone-depleting substance (ODS)

A substance known to deplete the stratospheric ozone layer. 

This includes CFCs, halons, HCFCs, and other halocarbons. 

Most ODSs are regulated by the Montreal Protocol, though 

some with very short lifetimes (e.g., CHBr3) or small 

anthropogenic sources (e.g., CH3Cl) are not.

ozone depletion

Chemical destruction of the stratospheric ozone layer by 

substances produced by human activities.

Ozone Depletion Potential

An index comparing the amount of global ozone destroyed 

by a particular ozone-depleting substance per unit mass 

compared to the amount destroyed by a reference gas 

(CFC-11) per unit mass.

production

The magnitude of ODS or substitute chemical produced by 

industry.

radiative forcing

Broadly defined as the difference between the incoming 

radiation energy and the outgoing radiation energy in the 

climate system. If more energy is incoming than outgoing, 

it tends to warm the climate (and is a planetary energy 

imbalance). A source of radiative forcing might be more 

solar energy, or more greenhouse gases for example. (This 

term is used in a more specific manner in IPCC).

stratosphere

The highly stratified region of the atmosphere above the 

troposphere extending from about 10 km (ranging from 

9 km in high latitudes to 16 km in the tropics on average) 

to about 50 km. In the stratosphere, temperatures increase 

with height.

substitutes for ozone-depleting substances

Chemicals that are used in place of the ozone-depleting 

substances that are regulated by the Montreal Protocol. 

These include hydrochlorofluorocarbons (HCFCs), which 

are ozone-depleting substances but not as potent as those 

they replace, and perfluorocarbons (PFCs), which do not 

deplete ozone but are greenhouse gases.

troposphere

The lowest part of the atmosphere from the surface to about 

10 km in altitude in mid-latitudes (ranging from 9 km in 

high latitudes to 16 km in the tropics on average) where 

clouds and “weather” phenomena occur. In the troposphere, 

temperatures generally decrease with height. 
217

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

Acronyms and abbreviations

A1 baseline halocarbon scenario

ACE Atmospheric Chemistry Experiment

AFEAS Alternative Fluorocarbons 

 Environmental Acceptability Study

AGAGE Advanced Global Atmospheric Gases 

 Experiment

AMTRAC Atmospheric Model with Transport and

 Chemistry

ATMOS Atmospheric Trace Molecule 

 Spectroscopy

Br atomic bromine

BrO bromine monoxide

BrOx reactive bromine

BrONO2 bromine nitrate

Bry inorganic bromine

C carbon

°C degree celsius

CALIPSO Cloud-Aerosol Lidar and Infrared 

 Pathfinder Satellite Observation

CCl4 carbon tetrachloride

C2Cl4 tetrachloroethene

CCM Chemistry Climate Model

CCSP Climate Change Science Program

CFC chlorofluorocarbon

CFC-11 trichlorofluoromethane

CFC-12 dichlorodifluoromethane

CH4 methane

CHBr3 tribromomethane

CHCl3 trichloromethane (chloroform)

CH2Br2 dibromomethane

CH2Cl2 dichloromethane or methylene chloride

CH3Br methyl bromide

CH3Cl methyl chloride

CH3CCl3 methyl chloroform

Cl chlorine

ClO chlorine monoxide

ClOx reactive chlorine

ClONO2 chlorine nitrate

ClOOCl chlorine monoxide dimer

CMDL Climate Monitoring and Diagnostics 

 Laboratory (NOAA)

CO2 carbon dioxide

CTM chemical transport model

CUE Critical Use Exemption

DU Dobson units

EECl Equivalent Effective Chlorine

EESC Equivalent Effective Stratospheric 

 Chlorine

EP Earth-Probe TOMS

EPA U.S. Environmental Protection Agency

ESRL Earth System Research Laboratory 

 (NOAA)

Gg gigagram, or one billion grams

GHG greenhouse gas

GMD Global Monitoring Division (NOAA/

 ESRL)

GOMOS Global Ozone Monitoring by Occultation 

 of Stars

Gt gigaton, or one billion metric tons

GtCO2 gigatons of carbon dioxide

GWP Global Warming Potential

H hydrogen

HALOE Halogen Occultation Experiment

HCFC hydrochlorofluorocarbon

HCl hydrogen chloride (or hydrochloric acid)

HFC hydrofluorocarbon

H2 hydrogen gas

H2O water

HOx reactive hydrogen

hPa hectopascal

ILAS Improved Limb Atmospheric 

 Spectrometer

IPCC Intergovernmental Panel on Climate 

 Change

IR infrared

K kelvin (unit of temperature)

kJ/m2 kilojoules per meter squared

km kilometer, or one thousand meters 

Kt kiloton, or one thousand metric tons

LOSU level of scientific understanding

MAM March-April-May

MBTOC Methyl Bromide Technical Options 

 Committee

MLS Microwave Limb Sounder

MSU Microwave Sounding Unit

Mt megaton, or one million metric tons

N nitrogen

NASA National Aeronautics and Space 

 Administration

NIH National Institutes of Health

NH Northern Hemisphere

NRL Naval Research Laboratory

NO nitric oxide

NOx reactive nitrogen

NO2 nitrogen dioxide

N2O nitrous oxide

NOAA National Oceanic and Atmospheric 

 Administration

NOCAR NOAA and NCAR model

O oxygen atom

O2 molecular oxygen

O3 ozone

Ox odd oxygen

ODP Ozone Depletion Potential

ODS ozone-depleting substance
The U.S. Climate Change Science Program Glossary and Acronyms

218

OH hydroxyl radical

OMI Ozone Monitoring Instrument

ppb parts per billion

ppmv parts per million by volume

ppt parts per trillion

PSC polar stratospheric cloud

POAM Polar Ozone and Aerosol Measurement

QBO quasi-biennial oscillation

QPS quarantine and pre-shipment

RAF radiation amplification factor

RF radiative forcing

SAGE Stratospheric Aerosol and Gas 

 Experiment

SAM Stratospheric Aerosol Monitor

SBUV Solar Backscatter Ultraviolet

SH Southern Hemisphere

SLIMCAT Single-Layer Isentropic Model of 

 Chemistry and Transport

SO2 sulfur dioxide

SPARC Stratospheric Processes and their Role 

 in Climate (WCRP)

SRES Special Report on Emissions Scenarios 

 (IPCC)

SROC Special Report on Ozone and Climate 

 (IPCC)

2-D two-dimensional

3-D three-dimensional

TEAP Technology and Economic Assessment 

Panel (UNEP)

TOMS Total Ozone Mapping Spectrometer

UNEP United Nations Environment 

 Programme

UNFCCC United Nations Framework Convention 

 on Climate Change

UV ultraviolet radiation

UV-vis ultraviolet/visible camera

UVA ultraviolet–A radiation

UVB ultraviolet–B radiation

UVC ultraviolet–C radiation

VPSC volume of polar stratospheric clouds

VSL very short-lived

VSLS very short-lived substances

WCRP World Climate Research Programme

WG1-AR4 Working Group I–Fourth Assessment

 Report (IPCC)

W per m2 watts per meter squared

WMGHG well-mixed greenhouse gases

WMO World Meteorological Organization
219

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure

REFERENCES

EXECUTIVE SUMMARY 

REFERENCES

Fahey, D.W., 2007: Twenty questions and answers about 

the ozone layer: 2006 update. In: Scientific Assessment 

of Ozone Depletion: 2006. Global Ozone Research and 

Monitoring Project report no. 50. World Meteorological 

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WMO (World Meteorological Organization), 2003: Scientific Assessment of Ozone Depletion: 2002. Global 

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World Meteorological Organization, Geneva, Switzerland, 498 pp.

WMO (World Meteorological Organization), 2007: Scientific Assessment of Ozone Depletion: 2006. Global 

Ozone Research and Monitoring Project report no. 50. 

World Meteorological Organization, Geneva, Switzerland, 572 pp. <http://www.esrl.noaa.gov/csd/assessments/2006/>

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Photography Credits

Cover/Title Page/Table of Contents

Image for Chapter 1, page 23, (Ozone cartoon), Used with permission from D. W. Fahey and the World Meteorological 

Organization, Geneva, Switzerland, 2007.

Image for Chapter 2, page 29, (Visible light spectrum), Used 

with permission from the Scientific Assessment of Ozone 

Depletion, 2006.

Image for Chapter 3, page 79, (Ozone hole), Paul Newman, 

NASA’s Goddard Space Flight Center.

Image for Chapter 4, page 111, (Polar stratospheric clouds), 

David J. Hofmann, NOAA.

Chapter 2

Page 72, (Polar stratospheric clouds), Lamont Poole, NASA’s 

Langley Research Center.
The U.S. Climate Change Science Program References

232
233

Trends in Emissions of Ozone-Depleting Substances, Ozone 

Layer Recovery, and Implications for Ultraviolet Radiation Exposure
The U.S. Climate Change Science Program References
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