Future Climates of the World: A Modelling Perspective

World Survey of Climatology Volume 16 FUTURE CLIMATES OF THE WORLD" A MODELLING PERSPECTIVE

World Survey of Climatology Volume 16 FUTURE CLIMATES OF THE WORLD" A MODELLING PERSPECTIVE

World Survey of Climatology

Editor in Chief"

H. E. LANDSBERG, College Park, Md. (U.S.A.)

Editors:

H. ARAKAWA,Tokyo (Japan) R. A. BRYSON, Madison, Wisc. (U.S.A.) O. M. ESSENWANGER,Huntsville, A1. (U.S.A.) H. FLOHN, Bonn (Germany) J. GENTILLI,Nedlands, W.A. (Australia) J. F. GRIFFITHS, College Station, Texas (U.S.A.) F. K. HARE, Ottawa, Ont. (Canada) H. E. LANDSBERG,College Park, Md. (U.S.A.) P. E. LYDOLPH,Milwaukee, Wisc. (U.S.A.) S. ORVIG, Montreal, Que. (Canada) D. F. REX, Boulder, Colo. (U.S.A.) W. SCHWERDTFEGER,Madison, Wisc. (U.S.A.) K. TAKAHASHI,Tokyo, Japan H. VAN LOON, Boulder, Colo. (U.S.A.) C. C. WALLI~N,Geneva (Switzerland)

World Survey of Climatology Volume 16

Future climates of the world: a modelling perspective

edited by A. H E N D E R S O N - S E L L E R S Macquarie University, Sydney (Australia)

1995 ELSEVIER Amsterdam- Lausanne - New Y o r k - Oxford- Shannon- Tokyo

ELSEVIER SCIENCE B.V. Sara Burgerhartstraat 25 P.O. Box 211, 1000 AE Amsterdam, The Netherlands

ISBN 0-444-89322-9

9 1995 Elsevier Science B.V. All rights reserved. No part of this publication may be reproduced, stored in a retrieval system, or transmitted in any form or by any means, electronic, mechanical, photocopying, recording, or otherwise, without the prior written permission of the publisher, Elsevier Science B.V., Copyrights & Permissions Department, P.O. Box 521, 1000 AM Amsterdam, The Netherlands. Special regulations for readers in the U.S.A. - This publication has been registerd with the Copyright Clearance Center, Inc. (CCC), 222 Rosewood Drive, Danvers, MA 01923. Information can be obtained from the CCC about conditions under which photocopies of parts of this publication may be made in the U.S.A. All other copyright questions, including photocopying outside of the U.S.A., should be referred to the copyright owner, Elsevier Science B.V., unless otherwise specified. No responsibility is assumed by the publisher for any injury and/or damage to persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the material herein.

This book is printed on acid-free paper. Printed in The Netherlands.

World Survey of Climatology

Editor in Chief: H.E. LANDSBERG Volume I

General Climatology, 1 Editor: H. FLOHN

Volume 2

General Climatology, 2 Editor: H. FLOHN

Volume 3

General Climatology, 3 Editor: H. FLOHN

Volume 4

Climate of the Free Atmosphere Editor: D. F. REX

Volume 5

Climates of Northern and Western Europe Editor: C. C. WALLI~N

Volume 6

Climates of Central and Southern Europe Editor. C. C. WALLI~N

Volume 7

Climates of the Soviet Union Editor: P. E. LYDOLPH

Volume 8

Climates of Northern and Eastern Asia Editor: H. ARAKAWA

Volume 9

Climates of Southern and Western Asia Editor: H. ARAKAWA

Volume 10 Climates of Africa

Editor: J. F GRIFFITHS Volume 11 Climates of North America

Editor: R. A. BRYSON Volume 12 Climates of Central and South America

Editor: W. SCHWERDTFEGER Volume 13 Climates of Australia and New Zealand

Editor: J. GENTILLI Volume 14 Climates of the Polar Regions

Editor: S. ORVIG Volume 15 Climates of the Oceans

Editor: H. VAN LOON Volume 16 Future Climates of the World: A Modelling Perspective

Editor: A. HENDERSON-SELLERS

This Page Intentionally Left Blank

Dedication for: The Holly Grange Household: Laura, Patrick, Philip, Stephen, Nicholas, Laurence, Alice and Eve

This Page Intentionally Left Blank

List of Contributors to this Volume

M. O. ANDREAE Max Planck Institut for Chemie Mainz (Germany)

J. C. GILLE National Center for Atmospheric Research Boulder, Colorado (U.S.A.)

E. J. BARRON

Earth System Science Center University Park, Pennsylvania (U.S.A.) A. BERGER

Universit6 Catholique de Louvain Institut d'Astronomie et de Geophysique Louvain-la-Neuve (Belgium)

A. HENDERSON-SELLERS

Climatic Impacts Centre Macquarie University North Ryde (Australia) G. J. HOLLAND BMRC Melbourne Victoria (Australia)

G. BRASSEUR

National Center for Atmospheric Research Boulder, Colorado (U.S.A.) H. CLEUGH

Centre for Environmental Mechanics CSIRO Canberra (Australia) H. DIAZ Climate Research Division, ARL Environmental Research Laboratories Boulder, Colorado (U.S.A.) R. E. DICKINSON Institute of Atmospheric Physics University of Arizona Tucson, Arizona (U.S.A.) M. P. DUDEK Atmospheric Sciences Research Center University at Albany Albany, New York (U.S.A.)

P. D. JONES Climatic Research Unit University of East Anglia Norwich (U.K.) England G. N. KILADIS Cooperative Institute for Research in Environmental Sciences Boulder, Colorado (U.S.A.) L. KUMP Earth System Science Center University Park, Pennsylvania (U.S.A.) X.-Z. LIANG Atmospheric Sciences Research Center University at Albany Albany, New York (U.S.A.) J. E. LOVELOCK Cornwall (U.K.)

S. MADRONICH National Center for Atmospheric Research Boulder, Colorado (U.S.A.) B. MCAVANEY BMRC Melbourne Victoria (Australia) T. H. PENG Ocean Chemistry Division NOAA/AOML/OCD Miami, Florida (U.S.A.)

M. RAMPINO

New York University New York, (U.S.A.) W.-C. WANG Atmospheric Sciences Research Center University at Albany Albany, New York (U.S.A.)

Preface

Future Climates of The WorM: A Modelling Perspective is Volume 16 of the highly prestigious series World Survey of Climatology (editor-in-chief Professor H.E. Landsberg) published by Elsevier Science. As such, Future Climates of the Worm is part of a major, highly respected series of climatology reference books. All the chapter authors and, indeed, their students have, on numerous occasions, used the earlier volumes of the World Survey of Climatology. Since these vary in depth, breadth and insight and, to a relatively minor extent, in contemporary relevance, no single phrase can truly describe them. Nevertheless, they almost invariably provide a highly useful, often magisterial, reference source for comparative analyses, lectures and papers. Certainly for an introduction to the current climate of most regions of the world, and to climatology in general, they are very helpful. Most graduates with climatological interests have used them profitably. Although the series never seems to have had an explicit audience in mind, it seems to have been directed at and presented material at a level comprehensible to an interested non-climatologist with a scientific background. The present volume is intended to be aimed at a similar audience (the generally interested reader, climatology undergraduates and graduates in any discipline) and to offer a state-ofthe-art overview of our understanding of future climates. However, as its name suggests, this book clearly differs from all the earlier volumes in The Worm Survey of Climatology. All the chapter authors accepted a very difficult challenge when we agreed to contribute to

Future Climates of the Worm for we all know that we cannot yet predict future climates. This book does not and could not give tables of values of future climatic variables. Instead, each topic we discuss is described so that the way in which it may impact projections of the Earth's future climate can be more fully understood. Discussing Future Climates of the Worm is difficult for a number of reasons. The space and time scales employed by those describing future climates are capricious. Here most chapters take their space scale to be at least continental, and often global, while the time scale is decadal, and the time span the next couple of centuries. This is obviously not the complete suite of future climates of the world in time or space and, not coincidentally, they are the scales that are best addressed by Global Climate Models (GCMs) and their associated needs for observational data for evaluation, tuning and improvement. This same, understandable, GCM bias has produced a focus on the climatic elements of energy and temperature at the expense of many other climatic parameters. This book comprises a state-of-the-art review of our current understanding of climatic prediction and the way in which this depends crucially upon improvements in, and improved understanding of, climatic models. Of necessity, this Volume contains material that is much more ephemeral than the climatic facts and figures in Volumes 1-15. Nevertheless, the study of climate has moved from data

XI

Preface collection "climatology" to the model and experimentally based predictions of "climatic science". A wag once questioned a weather forecaster friend of mine as follows: "Long ago the science of astronomy developed from the classification and mythology of astrology; when will meteorology become quantitative enough to be termed meteoronomy?" Perhaps, Future

Climates, building on all the previous books in this series on climatology, is a first step towards "climatonomy". The book comprises four main themes which follow an introductory chapter. The themes encompass the geologic perspective (I) and present-day observations (II) as they pertain to future climates. Theme III focusses on human factors affecting future climates and the final theme (IV) tackles planetary geophysiology and future climates. Within these sections there are a total of 14 chapters. The Syst~me International (SI) of units is employed throughout. Most of the material for this book was collected during 1992 and 1993 and thus both benefitted from, and was held up by, the revision of the Intergovernmental Panel on Climate Change (IPCC) Reports. We are very grateful to all our colleagues around the world who took time out of their busy schedules to review and comment upon earlier drafts. I am particularly grateful to Dr Graham Cogley, Dr Michael Manton and Dr Peter Robinson, who acted as the publisher's reviewers for the whole book, completing this arduous task in a timely and supportive fashion. I am also, personally, most grateful to my own research group at Macquarie University who reviewed my chapters and many of the others. Much of the research that underpins my own chapters in this book and the reference research which I undertook as editor were completed at the National Cenl~er for Atmospheric Research. I am grateful to Dr Bob Serafin, the Director, and Dr Warren Washington, the Director of the Climate and Global Dynamics Division, for their continuing hospitality to Affiliate Scientists and other visitors. I wish to express my thanks to the Vice-Chancellor of Macquarie University, Emeritus Professor Di Yerbury, and the Senior Executives of the University who have supported my research and, especially, the activities of the Climatic Impacts Centre so wholeheartedly over the last six years. Finally, I wish to thank Brian and Kendal for supporting me in everything I do - the mad and ludicrous ventures as well as the r e s t - thank you both. April 1994

Professor A. Henderson-Sellers Director Climatic Impacts Centre Macquarie University Sydney Australia

XII

Contents

Chapter 1. CLIMATES OF THE F U T U R E by A. HENDERSON-SELLERS Why consider the climate of the future? .........................................................................................

1

Focus of this book ..........................................................................................................................

1

Relationship to other reviews of the future climate ........................................................................

2

Earth's climate and its future ..........................................................................................................

4

Modelling the Earth's climate ........................................................................................................

7

Climates of the future .....................................................................................................................

14

References ......................................................................................................................................

17

I. GEOLOGICAL PERSPECTIVE ON FUTURE CLIMATE

Chapter 2. M O D E L L I N G THE RESPONSE OF THE CLIMATE SYSTEM TO A S T R O N O M I C A L FORCING by A. BERGER Introduction ....................................................................................................................................

21

Climates of the past ........................................................................................................................

22

Astronomical theory of paleoclimates ............................................................................................

25

Perturbations of Earth's orbital and rotational parameters ...............................................

25

Impact on solar radiation ..................................................................................................

29

Milankovitch theory .........................................................................................................

30

"Snapshot" simulations of past climate ..........................................................................................

32

Equilibrium models ..........................................................................................................

32

The Last Glacial Maximum (LGM) .................................................................................

34

Climate since the LGM ....................................................................................................

36

The Eemian interglacial ...................................................................................................

38

Initiation of the last glaciation ..........................................................................................

39

Conclusions ......................................................................................................................

39

Simulations of temporal evolution of climate ................................................................................

39

Transient response of a 2-D climate model to astronomical forcing ..............................................

44

Description of the 2-D LLN model ..................................................................................

44

Simulation of the last 200 ka ............................................................................................

46

The feedback mechanisms ...............................................................................................

50

Entering the glaciation, 5 1 - - T h e deglaciation process, 51

XIII

Contents The C O 2 - w a t e r vapour-albedo feedback mechanisms at the L G M ................................

52

Other feedbacks and future work .....................................................................................

54

Future climates under astronomical forcing ...................................................................................

55

References ......................................................................................................................................

59

Chapter 3. W A R M E R WORLDS: GLOBAL C H A N G E LESSONS F R O M E A R T H HISTORY by E. J. B ARRON Introduction ....................................................................................................................................

71

Cretaceous warmth .........................................................................................................................

72

The climate record ...........................................................................................................

72

Cretaceous climatic forcing factors ..................................................................................

75

Model predictions ............................................................................................................

75

Comparison of predictions with observations ..................................................................

79

Implications for the prediction of future climate .............................................................

80

An alternative warm climate: an Eocene example .........................................................................

81

The climate record ...........................................................................................................

81

Eocene climatic forcing factors ........................................................................................

84

Model studies ...................................................................................................................

84

Comparison with observations .........................................................................................

86

Implications for the prediction of future climates ............................................................

88

Discussion of implications for future climate and conclusions ......................................................

89

References ......................................................................................................................................

90

Chapter 4. CATASTROPHE: IMPACT OF COMETS AND ASTEROIDS by M. R. RAMPINO Introduction ....................................................................................................................................

95

Theoretical estimates of impact-induced climate change ...............................................................

97

Global dust cloud .............................................................................................................

97

Water vapour in the atmosphere ......................................................................................

99

Formation of NO and acid rain effects .............................................................................

100

Global wildfires ................................................................................................................

101

Impact into carbonates or evaporites ................................................................................

102

Other possible effects .......................................................................................................

102

Climatic changes from perturbations of biogeochemical cycles in the aftermath of impact-induced extinctions ..............................................................................................

102

Ocean alkalinity crisis at the K/T boundary? 104mPlankton extinction, DMS reduction and possible climatic warming, 106 Evidence of environmental perturbations at the K/T boundary .....................................................

107

Carbon isotope and carbonate changes ............................................................................

107

Oxygen-isotopic studies and paleotemperatures ..............................................................

108

Oxygen-isotopic paleotemperature studies at the K/T boundary .....................................

109

XIV

Contents Paleofloral, paleosol and other analyses, 111 Climate change at the K/T boundary: summary ...............................................................

112

Other geological boundaries: climatic changes and possible role of impacts ................................

113

Climatic changes at times of extinction events ................................................................

117

Late Pliocene, 117--Eocene-Oligocene, 1 1 8 ~ L a t e Triassic, 119~Permian-Triassic, 120---Frasnian-Famennian, 124--Ordovician-Silurian, 125~Precambrian/Cambrian, 126 Impact-induced volcanism? ............................................................................................................

127

Climatic changes caused by flood-basalt eruptions ........................................................................

129

Cooling of the climate ......................................................................................................

129

Greenhouse effect from volcanic outgassing ...................................................................

132

The possible role of volcanism in mass extinctions .........................................................

134

Future climates? .............................................................................................................................

135

References ......................................................................................................................................

136

II. OBSERVATIONS OF CLIMATE AND FUTURE PROJECTION

Chapter 5. OBSERVATIONS FROM THE SURFACE: PROJECTION FROM TRADITIONAL METEOROLOGICAL OBSERVATIONS by P. D. JONES Introduction ....................................................................................................................................

151

Development of instrumentation and meteorological observational networks ..............................

151

Homogeneity of temperature records ...............................................................................

152

Changes in instrumentation, exposure and measuring techniques, 153--Changes in station locations, 154mChanges in observation time and the methods used to calculate monthly averages, 154---Changes in the station environment, 155--Precipitation and pressure homogeneity, 155mData homogenisation techniques, 156 What surface observations tell us about the last 140 years ............................................................. Surface temperature .........................................................................................................

156 157

Hemispheric averages, 157~Spatial patterns of change, 159~Diurnal temperature change, 163 Precipitation .....................................................................................................................

164

Pressure ............................................................................................................................

167

The last 140 years in a longer term context ....................................................................................

168

Upper air data during the last 40 years ...........................................................................................

169

Comparisons of lower tropospheric data sets ...................................................................

170

Comparisons of lower stratospheric data sets ..................................................................

171

Explanations of the instrumental temperature record .....................................................................

172

Internal forcing of the climate ..........................................................................................

173

External forcing of the climate .........................................................................................

175

Projections to the future .................................................................................................................

179

The near future ( 1904

1854

(a) T H E C A S T 100 Y E A R S

1

Recent warming

2.

Little Ice Age

3.

Cold interval (Younger Dryas)

4.

Present interglacial (Holocene)

5.

Previous interglacial (Eemian)

6.

Present glacial age

7.

Permo-Carboniferous glacial age

8

Ordovician glacial age

9

Late Precambrian glacial age(s)

10.

Earth's origin

warm

1900 1700 1500

< 19

)

1300

)

1100

)

900

t

I

..2oc

(b) THE LAST 1,000 YEARS

M~dlabtude A=r Temperature cold warm 0

-

Estimated Summer Sea-Surface Temperature (~ 5 10 15 I''

''1

'''

I '

10

_

20 < ~-

30

10

>-

40

o

50

w 15 13 r (1:1 ( ot -

~

20

_

60

J

70

\~

80 90 100 110 120 130

~ 10~ (C) THE LAST 10,000 YEAI:::i'S

(d) THE LAST 100,000 YEARS

0 100

8, 200 ,a: 300 19 >- 400 "6 500 1/)

,"

.Q

600

=-9 700

QUATERNARY _0 , -, ~.4-- Last major chmat~c cycle E]'P- Ice sheets appear in N. Hemisphere 5 ~LIOCEN _ _ Antarctic ice sheet expands 10 k

15-

~

~dlOCENEI

30-

cold

sheet forms, mountain Qlaciers occur in the N Her~isphere

I /

o

] OLIGO- ~ CENE /

Small glaciers are widespread in Antarctica

"= >-

2

35-

900 1000

40-

J~.~---

c

Waters around Antarctica cool, sea ice forms

._

I h

O

800

warm I

Antarctic ice

2025-

Global Mean Temperature

3

J

>

I

4 5 - EOCENE I

\ 50?

~

55 '

- ~

Antarctica- Australian passage opens Cenozoic decline begins

~50"C (e) THE LAST 1,000,000,000 YEARS

(f) THE LAST 4,500,000,000 YEARS

Fig. l. The climate of the Earth as represented by average air temperature records over periods (a) 100 years (after JONES, 1995, Chapter 5), (b) 1000 years (after HUGHESand DIAZ, 1994; reprinted with permission, Kluwer Academic Publishers), (c) 10,000years (after US NATIONAL ACADEMY of SCIENCES, 1975), (d) 100,000 years (after BERGER, 1995, Chapter 2), (e) 1,000,000,000 years (after BERGER, 1988, 9 American Geophysical Union and BARRON, 1995, Chapter 3) and (f) 4,500,000,000 years (estimated from many sources including KUMP and LOVELOCK, 1995, Chapter 15 and HENDERSON-SELLERS,1989). Dashed lines indicate that the proxy records are inadequate for the reconstruction portrayed.

Climates of the future which is usually interpreted as indicating that globally averaged climatic conditions, such as temperature and precipitation, were not too different from the present day (see Chapter 15 by LOVELOCK and KUMP). Figure If shows mean temperatures rising slightly over the whole of Earth's history because we know that solar luminosity has increased by about 30% since the Earth was formed. It should also be noted that Fig. If is an interpretation of the probable record of the Earth's global mean temperatures based partly on observational evidence and partly on results from climate models which indicate that negative feedback effects can reduce large swings while "recovery" of habitable conditions from either a total planetary freeze or a "runaway greenhouse" is virtually impossible (KASTING, 1989; HENDERSON-SELLERS et al., 1991). The format of this book is guided, in part, by our understanding of this record of past climates and, in part, by the increasing recognition that human activities have the capacity to be as significant an influence on future climate as the natural factors which have, so far, shaped climatic evolution. The first section of this book reviews the geological history of climatic change assessing astronomical factors, pal~eo-evidence and massively disruptive events. The second section reviews observations of climate including both the historical record and the exploitation of satellite data. In the third section, the possible impacts of human activities on future climates are considered including the trace gas greenhouse, increased aerosols, pollution, ozone depletion and land-use change. In this section the future climates of the Earth's cities are also reviewed. Finally, the last section of the book closes the loop back to the geological record by reviewing the potential range of future climates in the context of global geophysiology: the perspective which considers the Earth system holistically. A potentially useful way of predicting patterns of future climate is to search for periods in the past when the global mean temperatures were similar to those we expect in the future, and then use the past spatial patterns as analogues of those which will arise in the future. For a good analogue, it is also necessary for the forcing factors (for example, greenhouse gases, orbital variations) and other conditions (for example, ice cover, topography) to be similar; direct comparisons with climate situations for which these conditions do not apply cannot be easily interpreted. HOUGHTON et al. (1990, p xxv) state that "analogues of future greenhouse-gas-changed climates have not been found" and therefore go on to say that, in the context of the IPCC scenarios for the next century, "we cannot therefore advocate the use of pala~o-climates as predictions of regional climate change due to future increases in greenhouse gases. However, pal~eo-climatological information can provide useful insights into climate processes, and can assist in the validation of climate models." (HOUGHTON et al., 1990, p xxv). In this book, pala~o-climatological evidence will be used (i) to offer a more complete view of "natural variability" (Chapter 2 by BERGER) and the processes underlying this variability, (ii) to set the framework for the longer time-scales of interest here and to examine the validation of climate models (Chapter 3 by BARRON) and (iii) to permit a quantitative discussion of unlikely events (Chapter 4 by RAMPINO) affecting climate or non-linear responses causing rapid changes (Chapter 14 by PENG). However, all chapters draw on the results of numerical models of the Earth's climate.

Modelling the Earth's climate

Modelling the Earth's climate The simplest possible way of constructing a model of the Earth's climate with which to predict its future is to consider the radiative balance of the globe as a whole. This is a zero dimensional model often written in the form S(1 - A )

(la)

"- tYTe 4

T s = T e + Tgreenhous e

(lb)

Here, S, the solar input, has a value of about 1370 W m -2 divided by 4 to give about 342 W m -2, the amount of solar radiation instantaneously incident at the planet per unit area of its (spherical) surface. Taking the Earth's albedo, A, as about 0.3, with tr, Stefan-Boltzmann constant of 5.67 x 10-8 W m -2 K -a, the effective blackbody radiating temperature of the Earth, Te, is found to be around 255 K. This is lower than the current global mean surface temperature of 288 K, the difference, about 33~

largely being due to the greenhouse ef-

fect. In Fig. 2, the Earth's globally and annually averaged radiation balance can be seen graphically. The numbers in parentheses represent energy as a percentage of the average solar constant- about 342 W m - 2 - at the top of the atmosphere. Note that nearly half the incoming solar radiation penetrates the clouds and greenhouse gases to the Earth's surface. These gases and clouds re-radiate most (i.e. 88 units) of the absorbed energy back down towards the surface. This is the basis of the mechanism of the greenhouse effect. The magnitude of the greenhouse effect is commonly measured as the difference between the blackbody emission at the surface temperature (a global average of 288 K gives 390 W m -2) and the outgoing infrared radiation at the top of the atmosphere (here 70 units or 239 W m-2), i.e. 151 W m -2 (see Fig. 9.1 in Chapter 9 by WANG et al.). In this zero dimensional model, the climate can only be changed if the incoming solar radiation alters, if the albedo changes or if the greenhouse effect is modified. Thus, to first order, 30

(a)

(b)

]~'~\~"~ IncomingSolar

/('.,:'l'..,,.=,~ so,..

k~,~,,~,~\~'~Radiation (100)

/

I ou,oo,~

I

/~, ~' Radiation(30) / / InT.ra.reo I Tota,Ener=,Ab~o~o~ / /'Rao=atK)n (70) 1 2 o I~.\"~ ~ \ \ ~ \ ~]x~/ and Infrared Re-emitted / 9 /~l~l I J ~ , \ \ \ \ \ \ \ ~ Reflectedb .~"~ by Atmosphere (66)--~./" --JU/ I I \~,,,~'~"~ ~/~ ~ . i # / i i ! (4) I

I~,\~\\~

,~

Atmospher,

[~ ~ I I

debr

by

OSPHERE

~ ~ ~ ' ~ - ~ ~ ' ' ~-----~',.~,a:esles" ~OPOPAUSE

\\\\.~ -"--(25)---~, ~ b y ~-"~.'-~,,~\ Surface(7) t ~ t Thermals (5 | I-"1 |

H,'

10

~ '11nnJ /,' / ,=n~aredlrom/ 1. . . . . ::::1// / an~tcn~uSdP; (e~)l

(~

~'-~

TROPOSP=HERE

~ t su~.aceis translated into sensible, latent and radiative losses J 200

I 250 Temperature

300 (K)

Fig. 2. Schematic of the Earth's energy budget (a) (modified from SCHNEIDER,1992, reproduced with the permission of Cambridge University Press and the University Corporation for Atmospheric Research) and temperature profile (b) of the lower atmosphere. Units are percentages of the incident solar radiation, 342 W m-2 and temperature.

Climates of the future

(a) 90ON 50~ 50~ 40~ 30~ 20= 10~ 0o 10~ 20~ 30~ 40~ 5O, 60" 90~

(b) Radiation

PolarCell ~" "~ ~

... i-".....

FerrelCell (' "=V'~~rr~~

A P~

H(~CelIs"

~,,

(~ TerrestrialInfraredRadiation ......

~r

~'~ L~

:'ha

~,,,.

~"

ReflectedSunlight /incOming Sunlight

~ ~

Outgoingiongwave .I~,-:... ~

80

160 240 Wm-2

SunHgat

i 320

TerrestrialInfraredRadiation Fig. 3. Schematic of (a) the latitudinal energy balance of the Earth and (b) the dynamics of the atmosphere (modified from HENDERSON-SELLERSand MCGUFFIE,1987). the Earth's climate is controlled by incident solar radiation, by the amount of this which is absorbed by the planet and by the thermal absorptivity of the gases in the atmosphere (e.g. KIEHL, 1992). Additional features control the climate. Of these, the most important are the latitudinal distribution of absorbed solar radiation (large at low latitudes and much less near the poles) as compared to the emitted thermal infrared radiation which varies much less with latitude (Fig. 3a). This latitudinal imbalance of net radiation for the surface-plus-atmosphere system as a whole (positive in low latitudes and negative in higher latitudes) combined with the effect of the Earth's rotation on its axis produces the dynamical circulation system of the atmosphere (Fig. 3b). Figures 2 and 3 describe schematically the most important processes controlling the atmospheric circulation and characteristics. The latitudinal radiative imbalance tends to warm air which rises in equatorial regions and would sink in polar regions were it not for the rotation of the Earth. The westerly waves in the upper troposphere in mid-latitudes and the associated high and low pressure systems are the product of planetary rotation affecting the thermally driven atmospheric circulation. The overall circulation pattern comprises thermally direct cells in low latitudes, strong waves in the midlatitudes and weak direct cells in polar regions. This circulation combined with the vertical distribution of temperature (Fig. 2b) represent the major aspects of the atmospheric climate system (e.g. SCHNEIDER, 1992). It is possible that even these fundamental characteristics of atmospheric dynamics could change in the future (Chapter 8 by MCAVANEY and HOLLAND). Processes in the atmosphere are strongly coupled to the land surface (Fig. 2a), to the oceans and to those parts of the Earth covered with ice and snow (the cryosphere). There is also strong coupling to the biosphere (the vegetation and other living systems on the land and in the ocean). These five components, the atmosphere, land, ocean, ice and biosphere (Fig. 4a), are now often considered to comprise the climate system (GARP, 1975). However, their time-scales differ quite significantly:

Modelling the Earth's climate atmosphere

-minutes to days

oceans

-weeks to thousands of years

biosphere

-years to hundreds of years

ice

-thousands to millions (or even hundreds of millions) of years -millions to hundreds of millions of years

land surface

The result of these different time-scales and of the complex interactions between these components of the climate system is a rich spectrum of climatic change (Fig. 4b). The largest

Changes of solar radiation SPACE

i Ten'estdalradiaItion

ATMOSPHERE

(a) u

H20, N2,02, CO2, 03 etc Aerosol Precipitation Atmosphere - land coupling Atmosphere - ice coupling

ICESHEEwT~ t ~ " " II I :.-: . . : . ~ 1

!

z~.. /

/

BIOMASS

-~~'

LAND \ -

III.so c.a

atmospheric cordposition

~

~SEA-ICE ~'

-'~ta~'/////////////A

ii

i?I

~

I

I

I

Wind stress

~X

II

FA _

ll- u~176 n

X l c e _ ocean ~

4k

EARTH

Changes of land features, orography, vegetation, albedo etc

104

Evaporation

Heat exchange

i

I

Changes of ocean basin shape, salinity etc. I

I

I

I

i

i

I

I

_

_

(b)

Annual

Tectonic

revolution

i

10 3

.~ -E

Orbital variations

Diurnal rotation r lo ,,..-

o r

102

t,,.

> G) ._> n-

c O

10

(1:1 "10

E

____j v , v

"'6

0.1 1010

tO4

-

to

1

1

1

1

1

1

I

1

1

1

1

1

1

10 9

10 8

10 7

10 6

10 5

10 4

10 3

10 2

10

I

011

10 .2

10 .3

10 -4

Period in Years

Fig. 4. (a) The Earth's climate system (after GARP, 1975). (b). Spectrum of climatic changes on Earth (after BERGER, 1988, 9 American Geophysical Union).

Climates of the future peaks in this spectrum relate to astronomical forcings: the Earth's rotation, its revolution around the Sun, variations in this orbit and the formation of the Solar System (see Chapter 2 by BERGER). Coupling amongst processes within the climate system components, such as the atmosphere with the oceans, the oceans and ice masses, is probably responsible for other peaks in this climate change spectrum. In this book, the perspective is longer than that adopted by IPCC, viz. the next century. Whilst recognizing the potential importance of anthropogenic activities, particularly but not exclusively the release of greenhouse gases (see Chapter 9 by WANG et al.), for future climate these are placed in the context of astrophysical, geological and biological variations which may also modify future climates. The most highly developed tool which we have to predict future climate is known as a general circulation model or global climate model, abbreviated to GCM. These models are based on the laws of physics and use descriptions in simplified physical terms (called parameterizations) of the smaller-scale processes such as those due to clouds and deep mixing in the ocean. In a climate model, an atmospheric component, essentially the same as a weather prediction model, is coupled to a model of the ocean, which can be equally complex (e.g. HENDERSON-SELLERS and MCGUFFIE, 1987; SCHLESINGER, 1988). The term GCM is nowadays taken to mean fully three-dimensional models of the atmosphere and oceans coupled together. If only the atmospheric component is represented the term AGCM (atmospheric GCM) is used. Such a model is generally linked to a simple representation of ocean temperatures such as offered by a mixed-layer ocean model. However, because the time-scales of, for example, the ice masses and the carbon cycle are very long they have not yet been incorporated into GCMs. For very long time-scale simulations of future, and past, climates simpler models are often used: energy balance models (EBMs) and twodimensional statistical dynamical models (2-D SDs). EBMs are called "one-dimensional", the dimension in which they vary being latitude. Vertical variations are ignored and the models are used with surface temperature as the dependent variable. Since the energy balance is allowed to vary from latitude to latitude, a horizontal energy transfer term must be introduced, so that the basic equation for the energy balance at each latitude, 0 is S[ 1 - A(0)] = OrZe(0) 4 + p C

AT"s(0) AT

+ transport out of zone 0

(2)

where pC is the heat capacity of the system and can be thought of as the system's "thermal inertia", S[ 1 - A(0)] is the absorbed solar radiation, crTe(0)4 the top-of-the-atmosphere net longwave radiative loss and Ts is the surface temperature varying with time, t. The radiation fluxes at the Earth's surface must be parameterized with care since conditions in the vertical are not considered. To a large extent, the effects of vertical temperature changes are treated explicitly. One possible use of EBMs is to employ them to obtain plausible distributions of ocean surface temperatures which can then be included in an AGCM simulation of a different climatic regime, such as a previous or future glacial epoch. The general circulation of the atmosphere can be schematically represented as cellular circulations. These Hadley, Ferrel and polar cells are meridional features, i.e. they consist solely of latitudinally averaged movement between zones (Fig. 3b). Most two-dimensional statistical dynamical climate models are constructed to simulate these motions. The two dimen-

10

Modelling the Earth's climate sions they represent explicitly are height in the atmosphere and latitude. Atmospheric variations around latitude zones (i.e. longitudinal variations) are neither resolved nor described. These models, in common with GCMs, solve numerically the fundamental equations listed in Table 1 and produce simulations of the large scale two-dimensional flow. The fundamental difference between these models and full atmospheric GCMs is that all the variables of interest are zonally averaged values. This zonal averaging is identified below by angle brackets. The equations to be solved are:

Zonal momentum 0 (u) Ot

O(u' v') S(v)-l-~ =F Oy

(3)

Meridional momentum (geostrophic balance) 0

f (u}+ R(T)--~y(In(p))= 0

(4)

Hydrostatic balance (vertical component) 0 (ln(p))

g

OZ

R(T)

(5)

Thermodynamic balance

o-~ ~o(.' r') Ot

+

o(w' r')

Oy

Oz

g

O(T)]=

Q

+ <w> (p)c. + oz j , ./.

o

"3

'

8.2 ,

.~_ I:~
200 km diameter) candidate crater, the Chicxulub impact structure in northern Yucatan (HILDEBRANDet al., 1991) dates from 65.2 _+0.4 Ma (SHARPTONet al., 1992). The apparent coincidence of the impact (or impacts) with the abrupt extinctions, including that of the dinosaurs and the possibility that climatic and environmental changes caused by the impact led to the extinctions, revived interest in the physical effects of large impact events. Although much attention has been focused on the relatively short-term (immediate to several hundred thousand years) effects of impacts on the environment, the possibility also exists that impact perturbations can trigger or set in motion longer term environmental

96

Theoretical estimates of impact-induced climate change and geological changes, such as ice ages (e.g. KYTE et al., 1988) or pulses of volcanic and/or tectonic activity that might, in themselves, affect long-term global climate (UREY, 1973; NAPIER and CLUBE, 1979; RAMPINO and STOTHERS, 1984a,b, 1988; RAMPINO and CALDEIRA, 1993). The possible connection between climatic change and mass extinctions has been a muchdebated subject, with some arguing that extinctions are generally related to cooling of the climate (e.g. STANLEY, 1984, 1988), while others have implicated warming in massextinction scenarios (e.g. MCLEAN, 1978). Estimates of climatic and environmental changes that might be caused by the impact of large extraterrestrial bodies have been approached in two basic ways: theoretical studies that attempt to estimate the kinds and magnitudes of climatic and geologic changes that impacts of various sizes might induce and study of proxy environmental and climate indicators in the geologic record at times of documented or suspected large-body impacts and associated mass extinctions.

Theoretical estimates of impact-induced climate change Global dust cloud Cosmic objects greater than a few kilometres in diameter, travelling tens of km s-1, release 107-108 megatonnes of TNT equivalent energy on impact with the Earth and are capable of producing global-scale dust clouds (Fig. 2). Events such as these are estimated from the population of Earth-crossing asteroids and comets to occur on time-scales of tens of millions of years (SHOEMAKER et al., 1990; CHAPMAN and MORRISON, 1994). For example, the initial kinetic energy of a 10 km diameter asteroid travelling at --20 km s-1, such as has been proposed to have caused the K/T mass extinction, is estimated at between --1024 and 1025 J (>108 Mt of TNT equivalent). Comets travel at higher velocities (up to --70 km s-1) and hence the impact of a significantly smaller comet can produce an explosion equivalent to

l O 0 i-..

Very Cloudy Day

10 "2

Limit of Photosynthesis t-"

0

1

0-4

-

.~ E 1 0-6 t,/) t10.8 _

Full Moonlight

l d 1~ Limit of Human

101~ 10 0

Vision

I

I

I

101

10 2

10 3

I

10 4

l

10 s

I

10 6

i

10 7

10 8

Impact Energy, Megatons Fig. 2. Estimates of the reduction in the light transmission through the atmosphere as a result of submicrometre dust in clouds from impacts of various magnitudes (in megatons of TNT equivalent). Submicrometre dust is estimated to comprise about 0.1% of the pulverized rock produced by the impact (K. ZAHNLE,personal communication, 1993).

97

Catastrophe: impact of comets and asteroids that of a 10 km asteroid. In such an impact, > 1016 g of excavated rock and vaporized asteroid, crust (and possibly ocean water) can be launched ballistically to >100 km and ~ 10-20% of these ejecta might be distributed globally (O'KEEFE and AHRENS, 1982a,b; MCLAREN and GOODFELLOW, 1990). The large amount of fine ejecta that was apparently distributed worldwide by the K/T boundary impact could have created a dense dust cloud while in the atmosphere and ALVAREZ et al. (1980) suggested that the darkness and short-term surface cooling beneath such a cloud led to cessation of photosynthesis and frigid conditions (impact winter), resulting in the observed mass extinctions. Calculations suggest that the solar transmission reduction resulting from a dust injection of>1016 g (only 0.01 times the mass of a 10 km diameter object) uniformly over the surface of the Earth would reduce photosynthesis by a factor of

105 (GERSTL and ZARDECKI, 1982) and could lead to a collapse of marine and terrestrial food chains. Initially, the Alvarez group estimated that the duration of the global dust cloud would have been several years, based on atmospheric opacity increases after historical volcanic eruptions. However, such relatively long-lived volcanic effects are the result of volcanogenic sulphuric acid aerosols in the stratosphere (RAMPINO and SELF, 1984) - volcanic dust settles out largely within 3-6 months - and thus the K/T boundary dust cloud would most likely have dissipated in a few months (POLLACK et al., 1982; TOON et al., 1982). Calculations (MILNE and MACKAY, 1982) and experiments (GRIFFIS and CHAPMAN, 1988) suggest that even a few months of such darkness would have been sufficient to produce the degree of extinction seen among pelagic phytoplankton at the K/T boundary, breaking the food chain. Analysis of carbon and nitrogen in the K/T boundary clay suggests that the amounts of incorporated marine and terrestrial biomass are approximately equivalent to the entire global standing biomass that would have been swept out of the ocean and buried by the ejecta and/or burned in global wildfires (GILMOUR et al., 1990). Predicted climatic effects from impact produced dust clouds are always severe, but to some extent model dependent. TOON et al. (1982) used a 1-D model of aerosol physics and a onedimensional radiative/convective climate model to calculate the evolution and climatic effects of such a dense dust cloud with an atmospheric lifetime of 3-6 months. Model results suggested that light levels would have remained too low for visibility for up to 6 months and too low for photosynthesis for up to 1 year after the impact. Calculations of surface cooling showed reductions in continental temperatures to below freezing for up to 2 years, whereas ocean-surface temperatures fell only a few degrees (Fig. 3). Increased reflectivity from snow cover in continental interiors might provide a positive feedback process that could prolong cooling once the dust settled out of the atmosphere (TOON et al., 1982). Possible changes in atmospheric thermal structure were also investigated in a similar model study by POLLACK et al. (1982). In normal times, solar energy is absorbed by the ground and the infrared emission back to space comes largely from the upper troposphere. At times of heavy dust loading, however, both the solar energy deposition level and the infrared emission level of the atmosphere are controlled by the dust and both lie within the stratosphere. The surface no longer receives any solar radiation and has a net energy deficit. Under these conditions, drastic surface cooling is inevitable and POLLACK et al. (1982) predicted land-surface temperature decreases of-30~

in the first 1-2 months after impact of a

10 km diameter body. Furthermore, these modelling studies did not include the possible ef-

98

Theoretical estimates of impact-induced climate change 1 Day

.....

300

1 Week 1 Month

.....

;

;

1 Year

......

......

290

280 t_.

t_.

270

260

250

.

i0 ~

.

.

.

.

.

.

.

1

.

.

.

.

.

.

.

10 5

.

!

10 6

.

.

.

.

.

.

.

.

!

10 7

.

.

.

.

.

.

.

.

10 8

Time (Seconds) after Impact Fig. 3. Estimates of the reduction in global temperatures as a result of a global cloud of fine dust produced by a 10 km diameter asteroid collision, using a one-dimensional radiative-convective climate model (open squares) (after TOON et al., 1982), and a 3-D Global Climate Model (diamonds) (after COVEYet al., 1990). fects of soot produced by global fires, which should exacerbate the situation (CROWLEY and NORTH, 1991). More recently, COVEY et al. (1990) utilized a more sophisticated three-dimensional global climate model (GCM) to study the effects of an impact-induced global dust cloud on climate. For a dust loading of 5 • 1015g after a 10 km diameter impact, their results were qualitatively similar to the earlier one-dimensional model results, although in the GCM land-surface cooling was somewhat mitigated by the heat capacity of the oceans and by atmospheric circulation patterns set up by the dust cloud that transported additional heat from the oceans to land areas. In this simulation, however, fixed sea-surface temperatures exaggerated the oceanic heat reservoir, providing an unrealistic situation. Land-surface cooling was less than that of POLLACK et al. (1982) by more than a factor of two, but the landsurface temperatures fell below freezing in less than a week (decreases of about 15~ compared to nearly 2 months in the one-dimensional simulation (Fig. 3). GCM studies also suggest an almost complete collapse of the hydrologic cycle after the impact. The GCM model results indicated that impacts of smaller objects (~1 km diameter) could cause less severe, but still quite significant decreases in temperature (-5~ over the continents within 20 days of impact. Such rapid and severe decreases in temperatures could have devastated global vegetation (TINUS and RODDY, 1990).

Water vapour in the atmosphere In the case of an oceanic impact, water would probably comprise a significant fraction of the ejecta lofted to high altitudes. For a 10 km impactor hitting the ocean, it is estimated that the expansion of a ,-500 km 3 steam bubble would drive a plume of mixed impactor vapour, de-

99

Catastrophe: impact of comets and asteroids bris and water vapour more than 100 kilometres above the surface (CROFT, 1982). Condensation of the cloud of water vapour and meteoroid vapour could produce a combination of dust and ice particles in the upper atmosphere (MELOSH, 1982). MCKAY and THOMAS (1982) considered the consequences of enhanced water vapour concentrations on the middle atmosphere (50-100 km) chemistry and heat budget. They estimated that the increased mixing ratio of hydrogen would have caused a decrease of ozone concentration above 60 km. The ozone reduction would lead to a lowering of the average height of the mesopause and lowering of average temperatures. These conditions were predicted to produce a long-lived (possibly 105-106 years) global layer of mesospheric ice clouds. An increase in global albedo of up to several percent is predicted from such clouds, but the net climatic effect (cooling or warming) depends upon how much infrared radiation from below is also trapped by the clouds, which is a function of the mean size of the ice crystals in the clouds. KYTE et al. (1988) suggested that similar ice clouds in the upper stratosphere following a relatively small (0.5 km diameter) ocean impact in the Late Pliocene might have triggered Northern Hemisphere Pleistocene glaciation. The size of the dust and ice grains produced in the atmosphere is also critical to the mean lifetime of the dust and ice clouds if the grains are too large, rapid fallout would occur and long-term darkness and climatic effects would be less likely. A measure of the initial grain-size distribution of material in the K/T boundary fallout layer would help in improving the understanding of the likely outcomes. The water vapour should locally supersaturate the stratosphere and thus much of the vapour might rapidly recondense and precipitate out of the atmosphere (CROFT, 1982). If the coagulation time of the water vapour is similar to that of the dust as calculated by TOON et al. (1982), then most of the water would leave the atmosphere within weeks to months. However, during this time, an enhanced greenhouse effect and reduction of stratospheric ozone could result (O'KEEFE and AHRENS, 1982b). EMILIANI(1980) and EMILIANI et al. (1982) suggested that such a water vapour greenhouse effect would cause sudden global heating and that an average global rise of only a few degrees in surface temperatures would be intolerable to many species, including reptiles weighing more than 25 kg. On the other hand, some evidence suggests that dinosaurs may have been quite adaptable to changing environmental conditions and hence climatic change alone may not have been responsible for their demise (LEHMAN, 1987). Formation of NO and acid rain effects

The shock waves from the KfI' bolide and the transfer of energy from the fine ejecta would have heated the global atmosphere. The fraction of energy transferred to the atmosphere is estimated to be ~40% (O'KEEFE and AHRENS, 1982a,b). Calculations suggest that, for a 1024 J impact, this could result in an immediate heat pulse with a global average temperature increase of >_15~ and ocean-surface temperature increases of ~5~ The high-temperature shock waves produced by the passage of the impactor through the atmosphere (TURCO et al., 1981; LEWIS et al., 1982) and interaction of the high velocity plume of ejecta with the atmosphere (O'KEEFE and AHRENS, 1982) could create large amounts of NO; PRINN and FEGLEY (1987) estimated as much as 3 x 1018 g. They predicted that the global atmosphere could be loaded with 100 ppmv of NO2, about 1,000 times more

100

Theoretical estimates of impact-induced climate change than during the heaviest air-pollution episodes today. This amount of NO 2 is enough to poison plants and animals, but it could also produce destructive nitric acid rain with a pH of ~ 1. NOx in the stratosphere would also rapidly remove the ozone layer. The large amount of fixed nitrogen deposited on the earth's surface after an impact might be denitrified to produce N20 and NO. Smog reactions, involving the oxidation of CH 4 produced by anaerobically decaying organic matter, might then lead to significant amounts of ozone in the troposphere (CRUTZEN, 1987). Production of N20 (along with CO2, HNO3 and CH4) could result in an enhanced greenhouse effect over the ~ 150 year lifetime of N20 in the atmosphere (CRUTZEN, 1987) and the N20 could also interact with the stratospheric ozone layer, depleting it over about a decade (TuRco et al., 1981). More recently, ZAHNLE (1990) estimated a much smaller total NO production of ~1014 mol from the impact of a 10 km diameter asteroid travelling at 20 km s-1. In his model, most of the global NO production came not from the initial shock wave or plume, but from the dispersed ejecta re-entering the earth's atmosphere. MELOSH et al. (1990) performed calculations on the heat emitted from these ejecta as they re-entered the atmosphere and found values that might reach 50-150 times the solar output for periods up to several hours. According to Zahnle's calculations, larger yields of NO would accompany only relatively rare grazing impacts. MACDOUGAL'S (1988) report of a global Sr-isotope spike at the K/T boundary that might indicate enhanced continental weathering supports the idea of acid rain at the boundary. The bleached white limestone beds below the boundary in Italy could have had a similar acidleached origin (LOWRIE et al., 1990). Oblique impacts were specifically investigated by SCHULTZ and GAULT (1990), who found that a 10 km object with a velocity of 20 km s-1, impacting at 10 ~ to the horizontal, could cause ricocheting fragments of size 0.1-1 km at hypervelocity, producing a global swarm of Tunguska-scale events. Nitrate production from such a swarm could greatly exceed that of a single large impactor. Energy partitioned to the target apparently increases for impact angles between 45 ~ and 15 ~, increasing the ability of the impact to vaporize potentially volatile targets (e.g. ocean water, CO2-rich terrains, sulphates). Oblique impacts also provide the possibility for inserting substantial amount of material in orbit, possibly forming a temporary debris ring around the planet. The ring shadow on the Earth's surface could create additional and extended climate changes through seasonal effects on solar insolation. Global wildfires

The immediate creation of a large, heated mass of low density air with peak temperatures of -20,000 K at the impact site and the intense heat emitted globally by the re-entering impact ejecta (MELOSH et al., 1990), are calculated to ignite combustible material and to create widespread wildfires. Large amounts of dead vegetation, killed by the lack of sunlight and the abrupt cooling (TINUS and RODDY, 1990), would have provided abundant fuel for the wildfires. The soot produced by the fires could add to the opacity of the atmosphere, exacerbating the darkness and cooling following an impact. Large amounts of such soot were discovered at the K/T boundary (estimated at > 1017 g worldwide), which supports the burning of a significant fraction of the terrestrial biomass (ANDERSet al., 1986; WOLBACH et al., 1985, 1988; GILMOUR et al., 1989). It has been estimated that such forest fires could pro-

101

Catastrophe: impact of comets and asteroids duce more than 1019 g of CO 2, 1018 g of CO, 1017 g of CH 4, 1016 g of N20 together with reactive hydrocarbons and oxides of nitrogen (NO + NO2). This might have led to an intense photochemical smog and an increase in the greenhouse effect estimated to produce a heating of ~ 10~ (CRUTZEN, 1987). Poisoning by noxious chemicals, such as polynuclear aromatic hydrocarbons (PAH) and CO from the wildfires (VENKATESANand DAHL, 1989), trace metals released by the impactor, the fires, and/or by acid-enhanced weathering (ERICKSON and DICKSON, 1987; LEARY and RAMPINO, 1990), or possibly cyanide released in a cometary impact (HSu, 1980; Hsu et al., 1982) could have contributed to the extinctions on land and in the surface oceans (DAVENPORT et al., 1990).

Impact into carbonates or evaporites Several recent studies suggest that the target material of the impact might be important in terms of climatic impact. For example, impact of a 10 km diameter asteroid into a carbonate-rich terrain (like that at Chicxulub) might lead to a release of CO2 that could increase atmospheric carbon dioxide levels by factors of 2-10. This increase in CO2 has been estimated to cause a possible rise in global temperatures of 2-10~ for 104-105 years after the impact (O'KEEFE and AHRENS, 1989). SIGURDSSON et al. (1993) noted that the proposed K/T impact site at Chicxulub in the Yucatan was underlain by thick Cretaceous evaporites including CaSO4 and that some K/T impact glasses were rich in Ca and contained up to 1% SO 3. They predicted that the total sulphur degassing from the evaporites at the Chicxulub impact site could have led to up to 1019 g of sulphate aerosols, enough to cause significant cooling after the dissipation of the short-lived dust/soot cloud and to create acid fallout equal to that proposed for the nitric acid produced by bolide-atmosphere interactions.

Other possible effects Comet showers on the Earth might be triggered by the passage of the Solar System through interstellar clouds, so that effects of the cloud encounter might coincide with comet impacts (RAMPINO and STOTHERS, 1984a, 1986). YABUSHITAand ALLEN (1989) suggested that hydrogen and dust accreted from interstellar clouds would form dust grains coated with ice in the upper atmosphere leading to global cooling and proposed that this may have happened at the end of the Cretaceous. Dust from large comets can also be swept up by the Earth both before and after collisions and ZAHNLEand GRINSPOON(1990) suggested this for the origin of extraterrestrial amino acids reported from the K/T boundary. It has also been proposed that large impacts might cause changes in the earth's rotation and tilt that could have global climatic effects (DAOHANet al., 1986), although major changes in orbital parameters from impact of a 10 km object seem unlikely.

Climatic changes from perturbations of biogeochemical cycles in the aftermath of impact-induced extinctions The Cretaceous/Tertiary boundary is marked by a major perturbation of the global carbon cycle, marked by severe reduction in pelagic carbonate deposition, decrease in biomass and

102

Theoretical estimates of impact-induced climate change oceanic productivity and changes in organic carbon deposition (Hsu and MCKENZIE, 1990). The carbon cycle involves the transfer of carbon between the solid Earth and the ocean/atmosphere system. Carbon dioxide releases associated with metamorphic decarbonation of sediments at subduction zones, mantle outgassing at mid-ocean ridges and the chemical weathering of carbonate rocks, represent the primary sources of carbon dioxide to the oceans and atmosphere. The primary source of alkalinity (excess positive charge) to the oceans is the riverine cation flux (primarily Ca 2§ and Mg 2§ derived from the chemical weathering of carbonate and silicate rocks. The two major sinks for alkalinity and carbon in the oceans and atmosphere are the deposition of shallow-water carbonate sediments and the accumulation of deep-water carbonates derived from the calcareous tests of pelagic plankton, primarily since their global dissemination in the Jurassic Period (~ 150 Ma) (BERNER et al., 1983). A major disruption of the pelagic alkalinity and carbon sinks may have occurred in the aftermath of the mass extinction and reduction in populations of calcareous plankton (~97% of planktonic foraminifera and ~88% of calcareous nannoplankton species disappeared) at, and/or very close to, the designated K/T boundary (THIERSTEIN, 1982a,b; KELLER, 1988a,b; BARRERA and KELLER, 1990; SMIT, 1990) and the apparent reduction of ocean primary productivity in the earliest Paleocene, the so-called "Strangelove Ocean" interval (HSu et al., 1982a,b; Hsu and MCKENZIE, 1985; ZACHOS and ARTHUR, 1986). KASTING et al. (1986) performed early model calculations with a 2-box ocean model of the carbon-cycle perturbations that might accompany a large planetesimal impact. They investigated Hsu's suggestion that an impact event would cause ocean surface waters to become undersaturated with respect to calcium carbonate, leading to the extinction of calcareous organisms and also to the delay in recovery of calcareous plankton during the Strangelove Ocean period of up to 500,000 years. Global darkening could have killed off most of the existing phytoplankton within several weeks to months, while deposition of atmospheric NOx created in the impact would have lowered the pH of ocean surface waters and released COa into the atmosphere (Hsu et al., 1982a,b). However, model results indicated that, because of mixing, ocean surface waters would only remain acidified for ~20 years or less. KASTING et al. (1986) also looked at the direct input of atmospheric CO2 from a comet or from oxidation of terrestrial and marine organic matter. The net results might be a greenhouse effect raising temperatures by several degrees and perhaps a surface ocean uninhabitable by calcareous organisms for a couple of decades, but the model runs suggested that the surface waters would not remain corrosive for long and thus could not have led to the Strangelove Ocean conditions of decreased carbonate productivity. However, a longer lived event is predicted if mixing of the deep ocean was involved. For example, if all life in the oceans were to die off suddenly and if sufficient time (> 1,000 years) were allowed for the deep waters to mix upward and equilibrate with the atmosphere, then the model predicted that the CO2 partial pressure of the atmosphere might increase ~2-3 times. Ocean mixing normally takes ~ 1,000 years (perhaps longer in the latest Cretaceous), but an oceanic impact(s) might temporarily speed up mixing. Model results indicated that these events could increase atmospheric temperatures by perhaps 3-4 ~, but still would not lead to a surface ocean undersaturated in calcium carbonate for a significant period of time.

103

Catastrophe: impact of comets and asteroids Hsu and MCKENZIE (1990) thought that blooms of opportunistic species of nannoplankton might have caused repeated rapid fluctuations in CO2 exchange between the ocean and atmosphere, creating unstable climatic conditions during the early Tertiary Strangelove Ocean period, although the blooms may have been partly in response to unstable climates, suggesting possible feedbacks between biota and climate.

Ocean alkalinity crisis at the K/T boundary ? At the time of the K/T mass extinctions, the rate of pelagic carbonate accumulation apparently fell by about a factor of three (ZACHOS and ARTHUR, 1986). The latest Cretaceous deep-water carbonate accumulation rate was about 20% of the total carbonate accumulation rate in the oceans (OPDYKE and WILKINSON, 1988). This contrasts markedly with modern carbonate sedimentation, in which more than 60% occurs in deep-water environments. The estimated reduction in calcium carbonate accumulation in the earliest Tertiary means that --6 • 1012 equivalents of alkalinity and 3 • 1012 mol of carbon could have accumulated in the oceans and atmosphere each year. CALDEIRA et al. (1990) utilized a carbonate-silicate cycle model coupled to a three-box ocean/atmosphere model to investigate the effects of the reduction in pelagic productivity at the K/T boundary. Their model results predicted that curtailed pelagic carbonate deposition could have caused ocean carbon and alkalinity concentrations to increase drastically, leading to an "alkalinity crisis" in the oceans and causing major fluctuations in atmospheric carbondioxide levels and climate. In order to investigate possible processes that might mitigate such perturbations, CALDEIRA and RAMPINO (1993) developed a five-box biogeochemical model of the oceans and atmosphere, incorporating a new parameterization for carbonate deposition in shallow waters, where the flux of calcium carbonate from the mixed layer to shallow-water sediments depends on the carbonate-ion concentration and temperature of the mixed layer. The model was run in a time-dependent mode, in which all pelagic carbonate and organic carbon production was instantaneously set to zero. Because the carbonate flux from the ocean-surface layer to the deep ocean was eliminated and because this flux carries more alkalinity than carbon, the short-term (less than -300 years) effect of this productivity crash was to cause a slight reduction in deep-ocean carbonate-ion concentration (Fig. 4a), which caused a reduction in the area of carbonate accumulation (Fig. 4b). This is consistent with evidence of increased dissolution of pelagic carbonate sediments just after the K/T boundary event (ZACHOS and ARTHUR, 1986; STOTT and KENNETT, 1990). In shallow continental shelf sections, carbonate dissolution at the K/T boundary seems to have ranged from negligible (KELLER, 1989a,b) to considerable, most likely depending on local conditions. In the model, as the pelagic biological carbon pump is turned off, CO2 begins to outgas from the surface oceans, increasing atmospheric CO2 levels (Fig. 4c). The elevated atmospheric CO2 causes an increase in the rate of chemical weathering of silicate and carbonate rocks. The elimination of the biological carbon pump in the model also increased total dissolved CO2 in the surface oceans causing carbonate-ion concentration to drop on the 1001,000 year time-scale. According to OPDYKE and WILKINSON (1990), this would act to reduce shallow-water carbonate accumulation and such conditions might also favour dolomitization in shallow-marine settings. Lowermost Tertiary sediments in the shallow-shelf

104

Theoretical estimates of impact-induced climate change



E 500 cq

9

c1,

400

~


8~

saline

waters in low to mid-latitude shelf and marginal sea regions. Some have suggested that the continental shelves of Antarctica were a source of Late Cretaceous ocean bottom waters in the Pacific (BARRERA et al., 1987), but study of benthic foraminifera from southern high latitudes argues against this idea (STOTTand KENNETT, 1990). Paleotemperatures for latest Cretaceous bottom waters in the South Atlantic range from -6.5 to 10~ at Deep Sea Drilling Project (DSDP) Site 527 on the Walvis Ridge in the South Atlantic (SHAcKLETON et al.,

108

Evidence of environmental perturbations at the K/T boundary 1984). Existing paleooceanographic evidence is insufficient to completely constrain Late Cretaceous/Early Tertiary ocean circulation and it is possible that patterns changed significantly in response to K/T boundary perturbations on the 105-106 year time-scale (KELLER, 1989a,b). Recently, KAJIWARA and KAIHO (1992) found evidence of positive sulphur isotope shift at the boundary indicating development of low oxygen conditions in the oceans.

Oxygen-isotopic paleotemperature studies at the K/T boundary The many published reports of oxygen-isotopic studies across the K/T boundary show a range of results. Some of the earliest oxygen-isotopic studies of pelagic sediments at the K/T boundary from Deep Sea Drilling Project (DSDP) Sites in the Atlantic Ocean (Sites 152, 356, 357, 384, Falkland and Agulhas Plateau Sites in the South Atlantic) were carried out by BOERSMA et al. (1979), who interpreted a persistent negative oxygen-isotopic excursion as evidence for a significant rise in ocean temperatures across the boundary, in both deep and surface waters. The temperature rise seems to have occurred before the early Tertiary "Globigerina" eugubina zone (beginning -50,000 years after the boundary), so that in the oldest Tertiary sample analysed (at DSDP Site 356), temperatures were already inferred to be 3~ higher than in the latest Cretaceous. However, part of the earliest Tertiary and the very latest Cretaceous section may be missing at this and other sites (BOERSMA and

SHACKLETON, 1981; MACLEOD and KELLER,1991 a,b). At DSDP Site 384, an inferred I~ by a 2~

surface-water cooling across the boundary is followed

warming by the end of the eugubina zone (-280,000 years after the K/T bound-

ary), whereas at Site 152, there is isotopic evidence for a net surface-water cooling of 2~ across the same sampling interval (BOERSMAand SHACKLETON, 1981), although the first 150,000 years of the Tertiary may be missing at Site 384 (MACLEOD and KELLER, 1991a,b). Site 357 shows no change in surface-water temperatures across the boundary, whereas the Agulhas Plateau Site shows an inferred net 2~ ferred to exhibit a cooling of 1-3~

cooling. Deep-water temperatures are in-

at Sites 384, 356 and 357, but a warming of 1-2~

at

Site 152 and on the Agulhas Plateau. THIERSTEIN (1980, 1981) reported that a 2 per mil shift towards lighter oxygen-isotopic values at the K/T boundary at South Atlantic Site 356 (over a thickness of 40 cm) indicates either a surface temperature increase of up to 10~ or alternatively a significant salinity decrease in surface waters. Note that this does not agree with the results of BOERSMA and SHACKLETON (1981) as quoted above and a hiatus may be present in the earliest Tertiary (MACLEOD and KELLER, 1991a,b). At Site 524, in the South Atlantic, Hsu et al. (1982) reported an average 6180 excursion in the carbonate fine fraction o f - 0 . 4 per mil between preand post-boundary samples, with a difference o f - 2 per mil between extremes in the standard deviation, suggesting warmings of at least 2~ and possibly up to 10~

although here

too there may be a minor hiatus. The maximum negative departure in 6180 occurs at about the same level as the most negative d l3C anomaly in the lowermost Tertiary. A negative departure in oxygen-isotopic ratios in benthic foraminifera of ~1 per mil at about the same time in some sections suggests a possible warming of ocean bottom waters by -4-5~ Late Cretaceous values (BOERSMA et al., 1979; HSU et al., 1982).

over

Boundary sections in Spain at Caravaca, Sopelana and Zumaya exhibit maximum negative 6180 excursions ranging from -1.4 to -2.4 per mil in bulk carbonate samples at or just

109

Catastrophe: impact of comets and asteroids above the K/T boundary, corresponding to possible warmings of ~7-12~

closely corre-

lated with the maximum negative 613C excursion (SMIT, 1990). Bulk carbonate from sections at Biarritz, France shows a maximum negative 6180 excursion of-1.5 per mil across the boundary and the boundary sequence at Lattengebirge in SW Germany, although possibly affected by diagenesis and containing a possible hiatus of ~200,000 years (MACLEOD and KELLER, 1991a,b), shows a negative 6180 excursion of-2.5 per mil, in both cases corresponding to the maximum 613C excursion. In the Pacific, DSDP Sites 47.2 and 577 at ~20~ paleolatitude and Site 465 at ~15~

pa-

leolatitude show little or no isotopic evidence for a significant temperature change across the KfF boundary (ZACHOS and ARTHUR, 1986). At Site 577, on the Shatsky Rise in the North Pacific, t5180 values from G. eugubina from below the KfF boundary are essentially the same as those above the boundary (GERSTEL et al., 1986). Some cooling of surface ocean waters is indicated ~20,000 years before the boundary, but some 50,000 years after the boundary may be missing in these sections (MACLEOD and KELLER, 1991a,b). It is possible that some important paleooceanographic changes preceded the recognized boundary and aberrant foraminifera in the early Tertiary may be evidence of ecological stress and/or instability of the marine environment after the boundary event. At southern high-latitude deep-water sites on the Maud Rise in the Weddell Sea (Ocean Drilling Project (ODP) Holes 689B and 690C), STOTT and KENNETT (1989, 1990) reported an oxygen-isotopic decrease of ~0.7-1.0 per mil in planktonic and benthic foraminifera beginning about 500,000 years before the KfI' boundary and a positive excursion of similar magnitude across the boundary (estimated from ~200,000 years before to ~100,000 years after). In this particularly detailed section, they interpreted the anomalies as representing a latest CretaceoUs year-round 4~

warming of Antarctic surface and intermediate waters,

followed by a similar cooling across the boundary, with return to previous paleotemperatures. The shifts in t5180 are mirrored by negative shifts in 613C and STOTT and KENNETT (1990) suggested that the two are related through feedback loops in the carbon cycle. In the Woodside Creek, New Zealand site (upper bathyal water depths, ~600-800 m), a very detailed study of the 0.6 cm thick boundary clay shows a -1.8 per mil shift in 6180 from 0 to 0.3 cm, followed by a + 1.3 per mil shift from 0.3 to 0.4 cm and a -1.0 per mil change at 0.4-0.6 cm (WOLBACHet al., 1988). The deposition time for the thin boundary clay is estimated by WOLBACH et al. (1988) at 1 per mil, indicating either cooling of deep waters, ice-volume increases, or some combination of the two, also marks the E/O transition interval (ZACHOS et al., 1992). The first indications of probable Antarctic glacierization at sea level are ice-rafted sediments in Middle Eocene (--45.5 Ma) deposits on the Maud Rise and Kerguelen Plateau, but a distinct strengthening of glacial conditions and the apparent onset of continental East Antarctic glaciation is recorded by abundant ice-rafted debris on the Kerguelen Plateau in the earliest Oligocene (--36 and 35.8 Ma). The occurrence of Nothofagus (southern beech) and relatively warm sea-surface temperatures around Antarctica suggest temperate conditions close to the ice sheet (EHRMANN and MACKENSEN, 1992; MACKENSENand EHRMANN, 1992). ZACHOS et al. (1992) found that the same transition may be a very rapid climatic change, marked by a spike of ice-rafted sediment in the southern Indian Ocean (ODP Site 748 on the Kerguelen Plateau) and by rapid shifts in the carbon- and oxygen-isotope records. Although the oxygen-isotopic shift is permanent, the pulse of ice rafting apparently lasted less than 100,000 years (Fig. 7), perhaps as little as 10,000 years! They interpreted this as indicating the brief appearance of a large Antarctic ice sheet during the transition from a non-glaciated to a glaciated world.

Late Triassic BICE et al. (1992) discovered two layers of shocked quartz at the Triassic-Jurassic (Tr/Jr) boundary (--205 Ma, but perhaps as young as 201 Ma), indicating multiple impacts, and the best estimated age of the --100 km wide Manicouagan crater in Quebec is indistinguishable from the estimated age of the boundary (OLSEN et al., 1990). Trace metal anomalies, including Ir (--200-400 ppt), are reported to occur in two European sections (Kendelbach, Austria and Audries Bay, England) (MCLAREN and GOODFELLOW, 1990), associated with a negative d 13C anomaly. The vertebrate faunal and floral transition are apparently quite marked in the Newark deposits of eastern North America, where the pollen and spore transition, showing a fern spike followed by gradual replacement of flora, took place in 106 km 3 are the largest known outpourings of basaltic magma and recent studies suggest that the eruptions are sudden, short-lived events, where the entire volume of the lava is erupted in a series of huge flows over a period of a few hundred thousand to perhaps a couple of million year. Although the convergence of evidence suggests that some (and possibly all) significant extinction events are correlated with extraterrestrial impacts, the K-Ar and other age data compiled by RAMPINO and STOTHERS (1988) showed a correlation between the mass extinctions and the times of continental flood-basalt eruptions over the past 250 Ma. More reliable 4~ and U-Pb age determinations that are now available for the flood basalt episodes support the initial dating and improve the correlation (COURTILLOTet al., 1986; BAKSI, 1988; BAKSI and FARRAR, 1990; DUNNING and HODYCH, 1990; RENNE and BASU, 1991; SEBAI et al., 1991; CAMPBELL et al., 1992; HEIMAN et al., 1992; RENNE et al., 1992; EBINGER et al., 1993; STOREVEDTet al., 1992), as shown in Table II. For example, the Deccan Flood Basalts of India (65.5 + 2.5 Ma) (VANDAMME et al, 1991) were erupted very close to the time of the end-Cretaceous mass extinction and large-body impact (64.5 _+0.1 Ma) and the Siberian Flood Basalts (248 + 2.3 Ma) correlate with the endPermian boundary clays (251 + 3 Ma) (CAMPBELLet al., 1992). Stratigraphic studies in India now place the Deccan eruptions near the paleontologically defined K/T boundary and the eruptions could have lasted only --250,000years (COURTILLOTet al., 1986). The most recent direct study of the Deccan lavas in relationship to foraminiferal changes at the K/T boundary in India (JAIPRAKASH et al., 1993) suggests that the first flows were erupted at the beginning of the faunal changes at the boundary; the first intertrappeans contain foraminiferal zones that begin up to ~350,000 years above the canonical K/T boundary, while the earliest Tertiary zone seems to be missing; and the last flows seem to have occurred about 500,000 years after the boundary. JAIPRAKASH et al. (1993) record that, within the stratigraphic resolution of the study, all Cretaceous planktonic

127

Catastrophe: impact of comets and asteroids foraminifera became extinct prior to or within the K/T transition interval marked by the first flows. Impacts large enough to form craters >100 km in diameter, flood-basalt eruptions and extinctions, are first-order geological events that apparently occur once every few tens of million of years. The recurrent close association in time of these major events over at least the last 250 Ma suggests that they are related (RAMPINO and STOTHERS, 1988) and recent statistical tests of the correlation exhibit a statistical significance of > 98% (STOTHERS,1993). Impacts of 10-km diameter asteroids or comets are estimated to produce earthquakes of Richter magnitude of--12, with large amplitude ground waves thousands of kilometres from the impact site that could deeply fracture and disturb the lithosphere and upper mantle. RAMPINO (1987) pointed out a possible mechanism of inducing flood-basalt eruptions at or near sites of large impacts through lithospheric fracturing and pressure-release melting in the upper mantle. Calculations suggest that large impacts (impactors of >10 km diameter) could excavate initial transient cavities deep enough to result in decompression melting in the upper mantle, with ensuing flood-basalt volcanism along deep impact-induced fractures that penetrated the lithosphere. WHITE and MCKENZIE (1988) raised objections to the impact-volcanism model, pointing to theoretical studies which suggested that large volumes of basaltic melt could only be produced by decompression melting of anomalously warm mantle (MCKENZIE and BICKLE, 1988), such as was inferred to exist primarily in the 2,000 km diameter regions of inferred hotspot swells over proposed mantle plume heads. Therefore, they inferred, impacts would have to preferentially hit these areas in order to trigger flood-basalt volcanism, which they considered very unlikely. However, calculations showed that the 2,000-km diameter hotspot swells related to the estimated 40 to ~ 100 present hotspots would cover a significant portion of the Earth (50 _ 25%), making impact into anomalously warm mantle surprisingly likely and it was thus concluded that impacts of large asteroids or comets might well be responsible for the initiation or triggering of flood-basalt volcanism and perhaps hotspot outbreaks, although this must be considered quite speculative at present (RAMPINOand STOTHERS, 1988). Moreover, the volcanism might be induced by lithospheric fracturing at the antipodes of the large impact sites and the Deccan and Siberian eruptions may have been near the reconstructed antipodes of the Chicxulub and proposed Falkland impact sites, respectively (RAMPINO and CALDEIRA, 1992). Seismic tomographic evidence now suggests that -50% of the global upper mantle is warm (possibly from broad mantle upwellings, within-mantle heating, or through insulation of the upper mantle by former continental lithosphere) (ANDERSON et al., 1992), providing temperature conditions under which large impacts might lead to significant decompression melting. In an impact-induced hotspot model, continued activity might be the result of a combination of impact heating and long-lasting perturbation of mantle geotherms. Examples of possible impact-related volcanism can be found in the earlier history of the Earth, for example in the Vredefort Dome and Bushveld Complex of South Africa, which have been interpreted as large impact basins (~400 km in diameter) created about 2 billion years ago (ELSTON and TWIST, 1990). Within the Bushveld, mafic rocks occur in overlapping ring complexes around of the central uplift of the basin (layered mafic rocks are apparently absent from the central part of the complex). ELSTONand TWIST (1990) interpret these as mantle melts induced by deep ring fracturing related to the impact structure.

128

Climatic changes caused by flood-basalt eruptions The Mackenzie igneous events in Canada represent one of the most widespread episodes of mafic magmatism on the continents. The mafic rocks consist of the Coppermine River and Ekalulia flood basalts (>140,000 km3), the Muskox layered intrusion and the spectacular Mackenzie dyke swarm that radiates from the Coronation Gulf across northwestern Canada to a distance of more than 2,400 km. The Muskox intrusion and the Mackenzie dykes have been dated by the U-Pb method using trace amounts of zircon or baddeleyite (ZrO2) with ages of 1270_ 4 and 1267 _+2 Ma BP, respectively (LECHEMINANT and HEAMAN, 1989). The contemporaneous flood basalts occur in the southern part of a large circular feature more than 500 km across, a portion of its perimeter outlining the Coronation Gulf itself. SEARS and ALT (1992) recently proposed that such Proterozoic mafic magmatism and layered intrusions are indicative of impact. The association of rapidly erupted flood basalts, a layered intrusion capped by granophyre (melted crustal rocks?) and radiating dykes, with a large circular structure supports the idea that the magmatism may have been generated by a large impact in the Middle Proterozoic (D. HYNDMAN, personal communication). However, despite these suggestive relationships, the geologic consensus is best summarized by the recent statement by MELOSH (1989) that, "to date, there is no firm evidence that impacts can induce volcanism."

Climatic changes caused by flood-basalt eruptions Flood basalt eruptions have been suggested as a cause of both cooling and warming of the climate and such climatic changes related to the eruptions have been proposed as a primary or contributing cause of mass extinctions - a "volcanist" alternative to the impact theory (e.g. MCLEAN, 1985; COURTILLOTet al., 1986; COURTILLOT, 1990; GLEN, 1990).

Cooling of the climate Climatic cooling at the earth's surface attributed to volcanic eruptions is primarily a result of the formation and spread of stratospheric HzSO4 aerosols. These aerosols are formed from sulphur volatiles (SO2, HzS, etc.) injected into the stratosphere by convective plumes rising above volcanic vents and fissures. Residence time of sulphuric acid aerosols in the stratosphere is several years; fine volcanic ash lofted into the stratosphere largely settles out of the atmosphere in 3-6 months (RAMPINO et al., 1988). Basaltic magmas are usually rich in dissolved sulphur (a function of the iron content of the magma) and commonly show concentrations of >1,500 ppm sulphur (SIGURDSSON, 1982, 1990). The largest historic basaltic eruption, the Laki fissure eruption in southern Iceland of 1783 A.D. (THORDARSON and SELF, 1993), produced an estimated 14.7 _ 1 km 3 of magma. SIGURDSSON (1982) developed techniques to estimate the sulphur release from magmas by petrological methods. According to their work, the Laki magma contained 800 to 1,000 ppm of sulphur prior to eruption and only ~ 150 ppm were retained in the erupted rocks, indicating an 85% release, equivalent to about 3 x 1013g of sulphur, or ~8 x 1013g of H2SO4 aerosols. The Crete, Greenland ice core contains a large acidity peak in 1783 A.D., equivalent to about 3 x 1013g of atmospheric sulphur in the Northern Hemisphere, but much of the

129

Catastrophe: impact of comets and asteroids H2SO4 (and HC1 and HF) contributing to the acidity peak could have been transported from Iceland to Greenland in the troposphere, so that total acidity should not be used as a direct measure of stratospheric loading. However, atmospheric effects of Laki were quite widespread. Haziness and dimming of sunlight were noticeable in 1783 A.D. in Europe. SIGURDSSON (1982) presented data that show 1783-1784 as the coldest winter ever recorded in the eastern United States (in a 225-year record), with a - A T = 4.8~ and he estimated a possible Northern Hemisphere -AT = 1~ The conversion of such a large amount of SO2 to H2SO 4 would have probably greatly depleted stratospheric H20 and OH, which at present consists of about 1015g globally (STOTHERS et al., 1986), but a large amount of water may have been injected into the stratosphere by the eruption plumes. A major unknown remains the possibility of strong selflimiting physical and chemical effects in very dense aerosol clouds, which may mitigate severe volcanic perturbations. More accurate modelling of dense stratospheric aerosol clouds and their effects on atmospheric dynamics and chemistry is needed. It is conventional wisdom that volcanic eruptions must inject a significant amount of sulphur volatiles into the stratosphere to have widespread climatic effects (RAMPINO and SELF, 1984). Basaltic eruptions are commonly effusive in character and much less explosive than silicic eruptions of comparable volume, so that they are commonly considered to be ineffective in lofting volatiles to significant altitudes (e.g. ALVAREZ, 1986). However, theoretical modelling of plume rise from fissure eruptions suggests that a fraction of sulphur volatiles from effusive eruptions can reach the stratosphere in eruption plumes driven by large basaltic fire-fountains (STOTHERS et al., 1986). Observations and model calculations suggest that the fire fountains in the Laki eruption were 800-1,400 m in height, with average masseruption rates during the most active phases of the eruption estimated from 4.5 to 5.6 • 103 kg s-1 per meter length of fissure (THORDARSON and SELF, 1993). According to theoretical models of convective plume rise from fissure eruptions (STOTHERS et al., 1986), the Laki plumes could have attained altitudes up to 12 km above sea level. In support of this estimate, observations of the widespread haze over Europe in 1783 suggest that the eruption plume may have just reached the tropopause over Iceland, which is located at a height of --11 km in Northern Hemisphere summer (STOTHERSet al., 1986). TRIPOLI and THOMPSON (1988) used a convective storm model to calculate plume growth, evolution and stabilization heights in a simulation of the very large (-700 km 3) Roza flow eruption (~ 15 Ma ago) in the Columbia River Flood Basalts. The complex model included effects such as wind shear, rotational effects, rainout and fallout. For a fissure assumed to be 90 km long, tremendous updrafts were produced that combined with the atmospheric circulation. As the surface winds settled into a large cyclonic circulation, the updrafts decreased by about a factor of 1/2. In this model 14% of the gases and ash were above the tropopause (11 km) after 3 h and 5% was above it after 6 h. This demonstrated that some material could be injected into the stratosphere by such large fissure eruptions, although soluble tracers used in this model did not make it past the tropopause. However, even 10% of the total release of SOa from a Roza-sized eruption is estimated to be equivalent to ~3 x 1014g of aerosols, more than three times the estimated output of Laki. Sulphate aerosols of modal diameter -0.5/tm should cool the earth's surface for all altitudes, with maximum cooling for low level aerosols. Therefore, volcanogenic aerosols in the troposphere will cool the surface as well. In most volcanic eruptions, tropospheric aero-

130

Climatic changes caused by flood-basalt eruptions sols have a very short lifetime of about a week, before they are washed out of the lower atmosphere. A flood basalt eruption, however, could release large amounts (1015 g) of sulphuric acid aerosols creating regional optical depths of > 10 and perhaps cooling the climate and suppressing convection, so that rainout might be less effective at removing the aerosols. A continuous curtain of aerosols from large, long-lived fire fountains might be a possibility. Scaling upward from the-15 km 3 Laki eruption, sulphur emissions from the entire sequence of Deccan Traps eruptions could have been equivalent of ~ 1019 g of sulphuric acid aerosols. Averaged over a 500,000 year eruption period, this is equal to -2 x 1013 g of sulphuric acid aerosols per year. Recently, COURTILLOT(1990) estimated that individual Deccan eruptions could have emitted 6 x 1018 g of sulphur and 6 x 1016 g of halogens into the atmosphere over a few hundred years. However, this sulphur release estimate is three orders of magnitude greater than that determined by using the Laki eruption as a scaling factor and seems to be much too high. Individual historic volcanic aerosol clouds in the stratosphere show an e-folding time of about 1 year, so that the ocean/atmosphere climate system does not have a chance to come to equilibrium with the typical volcanic perturbation. Historic eruptions that created -1013 g of aerosols in the stratosphere are associated with surface coolings of only a few 10ths of a ~ in the 1 or 2 years following the eruption. GCM model results, using the Goddard Institute for Space Studies (GISS) model, however, suggest that if maintained for 50 years, a stratospheric loading of about 1013 g of aerosols would lead to an equilibrium 5~ global cooling (D. RIND, personal communication, 1992). Therefore, if flood-basalt eruptions were continuous at the average activity level for 50 years or more, a significant cooling of climate is predicted. Eruptions in the Deccan show evidence of intermittent activity. Average thickness of individual flows in the Deccan is about 12 m (with a maximum of 160 m), suggesting that the total Deccan lava pile could be composed of--100 individual large flows or flow complexes. At Mahabaleshwar, the 1,200 m section contains about 50 flows and the flows are traceable for considerable distances within the Deccan (MAHONEY, 1988). Therefore, it is possible that flow complexes of up to -103 km 3 were erupted at intervals of several thousand years. Weathering profiles of 95% of calcareous plankton, a major floral transition and all of the extant dinosaurs. These extinctions can be explained by the severe conditions predicted at the boundary including a brief dark and cold "impact winter", global wildfires and climatic changes resulting from changes in greenhouse gases as a result of the marine extinctions. A review of the details of biotic changes and the environmental, isotopic and climatic indicators at times of recognized extinction events reveals some general patterns that may implicate large-body impacts and their climatic after-effects as a general cause of pulses of extinction. The major mass extinctions are accompanied by abrupt negative carbon-isotope shifts indicating sudden losses of biomass and reduction in marine productivity, abrupt negative shifts in oxygen-isotopic ratios indicating possible climatic warming, sudden crashes in global plankton and in organisms having planktonic growth stages, negative sulphur-isotope shifts suggesting development of anoxic conditions, and outbreaks of disaster fauna and flora suggesting a sudden ecological catastrophe. These conditions, combined with growing evidence of coincident large-body impacts from iridium anomalies, microtektites, shocked minerals and well-dated craters, lead to the working hypothesis that all extinction pulses are the result of the impacts of large asteroids or comets on the earth' s surface. Large impacts may trigger flood basalt volcanism, as indicated by a correspondence among mass extinctions, evidence of impact, and flood basalt eruptions over the last 250 Ma. Aerosols and carbon dioxide from massive flood basalt eruptions are predicted to produce climatic effects that could exacerbate environmental changes at mass extinction boundaries. Episodic delivery of impact energy to the Earth may even be great enough to have an effect on terrestrial geologic processes that determine long-term climate (cf. Chapter 2 by BERGER and Chapter 3 by BARRON). Although the chances of a large asteroid or comet collision in the near future are small such an occurrence is as likely as other "accepted" hazards of modern life. When the next largebody impact occurs the climatic effects, especially on human food resources, would be massive: crop failures would almost certainly result in mass starvation. In response to the possible threat, programmes to detect large asteroids and comets that are on collision courses with the Earth and to attempt to divert these bodies if necessary, have been proposed (CHAPMAN and MORRISON, 1994). If these are enacted the climatic impacts can be reduced or removed; if they are not, a massive climatic shock equivalent to that which caused the Cretaceous/Tertiary extinctions is a possible future climate.

References

AHARON, P., SCHIDLOWSKI,M. and SINGH, I. B., 1987. Chronostratigraphic markers in the endPrecambrian carbon isotope record of the Lesser Himalaya. Nature, 327: 699-702. ALVAREZ, L. W., ALVAREZ,W., ASARO, F. and MICHEL, H. V., 1980. Extraterrestrial cause for the Cretaceous-Tertiary extinction. Science, 208:1095-1108. ALVAREZ,W., 1986. Toward a theory of impact crises. Los, Trans. Am. Geophys. Union, 67: 649-658.

136

References ALVAREZ, W., ALVAREZ. L. W., ASARO, F. and MICHEL. H. V., 1982. Iridium anomaly approximately synchronous with terminal Eocene extinctions. Science, 216: 886-888. ALVAREZ, W. and ASARO, F., 1990. An extraterrestrial impact. Sci. Am., 263(4): 78-84. ALVAREZ, W., ASARO, F. and MONTANARI, A., 1990. Iridium profile for 10 million years across the Cretaceous-Tertiary boundary at Gubbio (Italy). Science, 250:1700-1702. ANDERS, E., WOLBACH, W. S. and LEWIS, R. S., 1986. Cretaceous extinctions and wildfires. Science, 234: 264. ANDERSON, D. L., TANIMOTO, T. and ZHANG, Y.-S., 1992. Plate tectonics and hotspots. The third dimension. Science, 256:1645-1651. ASARO, F., ALVAREZ,L. W., ALVAREZ,W. and MICHEL, H. V., 1982. Geochemical anomalies near the Eocene/Oligocene and Permian/Triassic boundaries. Geol. Soc. Am., Spec. Pap., 190:517-528. ASKIN, R. A., 1990. Campanian to Paleocene spore and pollen assemblages of Seymour Island, Antarctica. Rev. Palaeobot. Palynol., 65:105-113. BAKSI, A. K., 1988. Reply to comment on "Critical evaluation of the age of the Deccan traps, India. Implications for flood basalt volcanism and faunal extinctions". Geology, 16: 758-759. BAKSI, A. K. and FARRAR, E., 1990. Evidence for errors in the geomagnetic polarity time scale at 1715 Ma. 4~ dating of basalts from the Pacific Northwest, USA. Geophys. Res. Lett., 17: 1117-1120. BALME, B. E., 1979. Palynology of Permian-Triassic boundary beds at Kap Stosch, East Greenland. Meddelelser Om Gronland, Bd. 200, Nr. 6: 35. BARLOW, N. G., 1990. Application of the inner solar system cratering record to the earth. Geol. Soc. Am., Spec. Pap., 247:189-195. BARRERA, E, HUBER, B. T., SAVIN, S. M. and WEBB, P.-N., 1987. Antarctic marine temperatures. Late Campanian through early Paleocene. Paleoceanography, 2:21--47. BARRERA, E. and KELLER, G., 1990. Stable isotope evidence for gradual environmental changes and species survivorship across the Cretaceous/Tertiary boundary. Paleoceanography, 5: 867-890. BECKER, R. T., HOUSE, M. R., KIRCHGASSER,W. T. and PLAYFORD,P. E., 1991. Sedimentary and faunal changes across the Frasnian/Famennian boundary in the Canning Basin of Australia. Hist. Biol., 5: 183-196. BENTON, M. J., 1985. Mass extinction among non-marine tetrapods. Nature, 316:811-814. BENTON, M. J., 1986. More than one event in the late Triassic mass extinction. Nature, 321:857-861. BERNER, R. A., 1990. Atmospheric carbon dioxide levels of Phanerozoic time. Science, 249: 13821396. BERNER, R. A., LASAGA, A. C. and GARRELS, R. M., 1983. The carbonate-silicate geochemical cycle and its effect on atmospheric carbon dioxide over the past 100 million years. Am. J. Sci., 283: 641683. BHANDARI, N., SHUKLA, P. N. and PANDEY, J., 1991. Iridium enrichment at the Cretaceous/Tertiary boundary in Meghalaya. Curr. Sci., 56: 1003-1005. BHANDARI, N., SHUKLA, P. N. and AZMI, R. J., 1992. Positive Europium anomaly at the PermoTriassic boundary, Spiti, India. Geophys. Res. Lett., 19: 1531-1534. BHANDARI, N., SHtJI~A, P. N. and CINI CASAGNOLI, G., 1993. Geochemistry of some K/T sections in India. Palaeogeogr., Palaeoclimatol., Palaeoecol., 104:199-211. BICE, D. M., NEWTON, C. R., MCCAULEY, S., REINERS, P. W. and MCROBERTS, C. A., 1992. Shocked quartz at the Triassic-Jurassic boundary in Italy. Science, 259: 443-446. BOERSMA, A. and SHACKLETON, N. J., 1981. Oxygen- and carbon-isotope variations and planktonicforaminifer depth habitats, late Cretaceous to Paleocene, Central Pacific, Deep Sea Drilling Project Sites 463 and 465. Initial Reports of the Deep Sea Drilling Project, Vol. 62. Government Printing Office, Washington, DC, pp. 513-526. BOERSMA, A., SHACKLETON,N., HALL, M. and GIVEN, Q., 1979. Carbon and oxygen isotope records at DSDP site 384 (North Atlantic) and some Paleocene paleotemperatures and carbon isotope variations in the Atlantic Ocean. Initial Reports of the Deep Sea Drilling Project, Vol. 43. Government Printing Office, Washington, DC, pp. 695-717. BRAND, U., 1989. Global climatic changes during the Devonian-Mississippian. Stable isotope biogeochemistry of brachiopods. Palaeogeog., Palaeoclimatol., Palaeoecol., Global Planet. Change Sec, 75:311-329. BRENcHLEY, P. J., 1989. The Late Ordovician extinction. In: S. K. DONOVAN (Editor), Mass Extinctions, Processes and Evidence. Columbia University Press, New York, pp. 104--132.

137

Catastrophe: impact of comets and asteroids BUGGISCH, W., 1991. The global Frasnian-Famennian "Kellwasser Event". Geol. Rundsch., 80: 4972. CALDEIRA, K. and RAMPINO, M. R., 1990a. Carbon dioxide emissions from Deccan volcanism and a K/T boundary greenhouse effect. Geophys. Res. Lett., 17: 1299-1302. CALDEIRA, K. and RAMPINO, M. R., 1990b. Deccan volcanism, greenhouse warming and the Cretaceous/Tertiary boundary. Geol. Soc. Am., Spec. Pap., 247:117-123. CALDEIRA, K. and RAMPINO, M. R., 1993. The aftermath of end-Cretaceous mass extinction. Biogeochemical stabilization of the carbon cycle and climate? Paleoceanography, 8:515-525. CALDEIRA, K., RAMPINO, M. R., VOLK, T. and ZACHOS, J. C., 1990. Biogeochemical modelling at mass extinction boundaries. Atmospheric carbon dioxide and ocean alkalinity at the K/T boundary. In: E. G. KAUFFMANand O. H. WALLISER (Editors), Extinction Events in Earth History, SpringerVerlag, Berlin, pp. 333-345. CAMPBELL, I. H., CZAMANSKE,G. K., FEDORENKO,V. A., HILL, R. I. and STEPANOV,V., 1992. Syn chronism of the Siberian Traps and the Permian-Triassic boundary. Science, 258:1160-1163. CAPUTO, M., 1985. Late Devonian glaciation in South America. Palaeogeog., Palaeoclimatol., Palaeoecol., 51:291-317. CASSINIS, G. (Editor), 1988. Proceedings of the field conference on Permian and Permian-Triassic boundary in the South-Alpine segment of the western Tethys and additional regional reports. Mem. Soc. Geol. Ital., 34: 366. CHAPMAN,C. R. and MORRISON,D., 1994. Impacts on the Earth by asteroids and comets: assessing the hazard. Nature, 367: 33-40. CHIFANG, C., YAOQI, Z. et al., 1992. Geochemical constraints on the Permo-Triassic boundary event in South China. In: W. C. SWEETet al. (Editors), Permo-Triassic Events in the Eastern Tethys. Cambridge University Press, Cambridge, pp. 158-168. CLAEYS, P., CASIERJ-G. and MARGOLIS, S.V., 1992. Microtektites and mass extinctions. Evidence for a Late Devonian asteroid impact. Science, 257:1102-1104. COCKS, L. R. M. and RICKARDS, R. B. (Editors), 1988. A global analysis of the Ordovician-Silurian boundary. Bull. Br. Mus. (Nat. Hist.), Geol. No. 43: 394. COLBATH, G. K., 1986. Abrupt terminal Ordovician extinction in phytoplankton associations, southern Appalachians. Geology, 14: 943-946. CORFIELD, R. M., CARTLIDGE,J. E., PREMOLI-SILVA,I. and HOUSLEY,R. A., 1990. Oxygen and carbon isotope stratigraphy of the Palaeogene and Cretaceous limestones in the Bottaccione Gorge and the Contessa Highway sections, Umbria, Italy. Terra Nova, 3: 414--422. CORLISS, B. H., AUBRY, M.-P., BERGGREN,W. A., FENNER, J. M., KEIGWINJR., L. D. and KELLER, G., 1984. The Eocene/Oligocene boundary event in the deep sea. Science, 226: 806-810. COURTILLOT,V., 1990. A volcanic eruption. Sci. Am., October, 85-92. COURTILLOT, V., BESSE, J., VANDAMME, D., MONTIGNY, R., JAEGER, J.-J. and CAPPETTA, n., 1986. Deccan flood basalts at the Cretaceous-Tertiary boundary. Earth Planet. Sci. Lett., 80: 361374. COVEY, C., SCHNEIDER, S. n. and THOMPSON, S. L., 1984. Global atmospheric effects of massive smoke injection from a nuclear war: results from general circulation model simulations. Nature, 308: 21-25. COVEY, C., GHAN, S. J., WALTON,J. J. and WEISSMAN,P. R., 1990. Global environmental effects of impact-generated aerosols: results from a general circulation model. Geol. Soc. Am., Spec. Pap., 247: 263-270. CROCKET,J. n., OFFICER,C. B., WEZEL,F. C. and JOHNSON,G. O., 1988. Distribution of noble metals across the Cretaceous/Tertiary boundary at Gubbio, Italy. Iridium variation as a constraint on the duration and nature of Cretaceous/Tertiary boundary events. Geology, 16: 77-80. CROFT, S. K., 1982. A first-order estimate of shock heating and vaporization in oceanic impacts. Geol. Soc. Am., Spec. Pap., 190:143-151. CROWLEY, T. J. and NORTH, G. R., 1991. Environmental consequences of an asteroid impact. In: Paleoclimatology. Oxford University Press, New York, Ch. 9, pp. 171-182. CRUTZEN,P. J., 1987. Acid rain at the K/T boundary. Nature, 330: 108-109. DAOHAN, C., LINZHONG,L. and JIAQING, Z., 1986. Mass extinction, earth's rotating and periodic cratering. Sci. Sin., 29:531-539. DAO-YI, X., QIN-WEN, Z., YI-YIN, S., YAN, Z., ZHI-FANG, C. and JIN-WEN, H., 1989. Astrogeological Events in China. Van Nostrand Reinhold, New York, p. 264.

138

References DE SILVA, S. L., WOLFF,J. A. and SHARPTON,V. L., 1990. Explosive volcanism and associated pressures; implications for models of endogenically shocked quartz. Geol. Soc. Am., Spec. Pap., 247: 139-145. DAVENPORT, S. A. et al., 1990. Chondritic metal toxicity as a seed stock kill mechanism in impactcaused mass extinctions. Geol. Soc. Am., Spec. Pap., 247: 71-76. DELAUBENFELS, M. W., 1956. Dinosaur extinction, one more hypothesis. J. Paleontology, 30: 207212. D'HONDT, S. L., KELLER, G. and STALLARD,R. F., 1987. Major element compositional variation within and between different late Eocene microtektite strewnfields. Meteoritics, 22:61-79. DICKINS, J. M., 1993. Climate of the Late Devonian to Triassic. Palaeogeog., Palaeoclimatol., Palaeoecol., 100: 89-94. DONOVAN, S. K., 1989. Introduction. In: S. K. DONOVAN (Editor), Mass Extinctions, Processes and Evidence. Columbia University Press, pp. xi-xiv. DOWSETT, H. J. and LOUBERE, P., 1992. High resolution late Pliocene sea-surface temperature record from the Northeast Atlantic Ocean. Mar. Micropaleontol., 20:91-105. DUNNING, G. R. and HODYCH, J. P., 1990. U/Pb zircon and baddeleyite ages from the Palisades and Gettysburg sills of the northeastern United States. Implications for the age of the Triassic/Jurassic boundary. Geology, 18: 795-798. EBINGER, C. J., YEMANE, T., WOLDEGABRIEL,G., ARONSON, J. L. and WALTER, R. C., 1993. Late Eocene-Recent volcanism and faulting in the southern main Ethiopian rift. J. Geol. Soc., London, 150: 99-108. EHRMANN, W. U. and MACKENSEN, A., 1992. Sedimentological evidence for the formation of an East Antarctic ice sheet in Eocene/Oligocene time. Palaeogeog., Palaeoclimatol., Palaeoecol., 93: 85112. ELSTON, W. E. and TWIST, D., 1989. Is the Rooiberg Felsite (Bushveld Complex) an impact melt? Unpublished. EMILIANI, C., 1980. Death and renovation at the end of the Mesozoic. Los, Trans. Am. Geophys. Union, 61: 504-506. EMILIANI, C., KRAUS, E. B. and SHOEMAKER,E. M., 1982. Sudden death at the end of the Mesozoic. Earth Planet. Sci. Lett., 55:317-334. ENGELHARDT, W. v., MATTHAI, S. K. and WALZEBUCK, J., 1992. Araguinha impact crater, Brazil. I. The interior part of the uplift. Meteoritics, 27: 442-457. ERICKSON, D. J. and DICKSON, S. M., 1987. Global trace-element biogeochemistry at the K/T boundary. Oceanic and biotic response to a hypothetical meteorite impact. Geology, 15:101 4-1017. ESHET, Y., 1990. The palynostratigraphy of the Permian-Triassic boundary in Israel. Two approaches to biostratigraphy. Isr. J. Earth-Sci., 39: 1-15. FAN, D., YANG, R. and HUANG, Z., 1984. The lower Cambrian black shale series and the iridium anomaly in South China. In: Academia Sinica, Developments in Geoscience. Contribution to 27th Geological Congress, Moscow. Science Press, Beijing, pp. 215-224. GANAPATHY,R., 1980. A major meteorite impact on the earth 65 million years ago. Evidence from the Cretaceous-Tertiary boundary clay. Science, 209:921-923. GANAPATHY, R., 1982a. Evidence for a major meteorite impact on the Earth 34 million years ago. Implication for Eocene extinctions. Science, 216: 885-886. GANAPATHY, R., 1982b. Evidence for a major meteorite impact on Earth 34 million years ago. Implication on the origin of North American tektites and Eocene extinction. Geol. Soc. Am., Spec. Pap., 190: 513-516. GELDSETZER, n. H. J., GOODFELLOW, W. D., MCLAREN, D. J. and ORCHARD, M. J., 1987. Sulfurisotope anomaly associated with the Frasnian/Famennian extinction, Medicine Lake, Alberta, Canada. Geology, 15: 393-396. GERLACH, T. M. and GRAEBER, E. J., 1985. Volatile budget of Kilauea volcano. Nature, 313: 273277. GERSTEL, J., THUNELL, R. C., ZACHOS, J. C. and ARTHUR, M. A., 1986. The Cretaceous/Tertiary boundary event in the North Pacific. Planktonic foraminiferal results from Deep Sea Drilling Project Site 577, Shatsky Rise. Paleoceanography, 1:97-117. GERSTL, S. A. and ZARDECKI, A., 1982. Reduction of photosynthetically active radiation under extreme stratospheric aerosol loads. Geol. Soc. Am., Spec. Pap., 190:201-210. GILMOUR, I., WOLBACH, W. S. and ANDERS, E., 1989. Major wildfires at the Cretaceous-Tertiary

139

Catastrophe: impact of comets and asteroids boundary. In: S. V. M. CLUBE (Editor), Catastrophes and Evolution, Astronomical Foundations. Cambridge University Press, Cambridge, pp. 195-213. GILMOUR, I., WOLBACH, W. S. and ANDERS, E., 1990. Early environmental effects of the terminal Cretaceous impact. Geol. Soc. Am., Spec. Pap., 247: 383-390. GLASS, B. P., 1982. Possible correlations between tektite events and climatic changes? Geol. Soc. Am., Spec. Pap., 190:251-256. GLEN, W., 1990. What killed the dinosaurs? Am. Sci., 78: 354-371. GOODFELLOW, W. O., GELDSETZER, n. n. J., MCLAREN, O. J., ORCHARD, M. J. and KLAPPER, G., 1989. Geochemical and isotopic anomalies associated with the Frasnian-Famennian extinction. Hist. Biol., 2:51-72. GOODFELLOW, W. D., NOWLAN, G. S., MCCRACKEN, A. D., LENZ, A. C. and GREGOIRE, D. C., 1992. Geochemical anomalies near the Ordovician-Silurian boundary, Northern Yukon Territory, Canada. Hist. Biol., 6: 1-23. GRATZ, A. J., NELLIS, W. J. and HINSEY, N. A., 1992. Laboratory simulations of explosive volcanic loading and implications for the cause of the K/T boundary. Geophys. Res. Lett., 19: 1391-1394. GRIEVE, R. A. F., 1991. Terrestrial impact, the record in the rocks. Meteoritics, 26: 175-194. GRIFFIS, K. and CHAPMAN, D. J., 1988. Survival of phytoplankton under prolonged darkness. Implications for the Cretaceous-Tertiary boundary darkness hypothesis. Palaeogeog., Palaeoclimatol., Palaeoecol., 67: 305-314. GRUSZCZYNSKI, M., HALAS, S., HOFFMAN, A. and MALKOWSKI,K., 1989. A brachiopod calcite record of the oceanic carbon and oxygen isotope shifts at the Permian-Triassic transition. Nature, 337: 64-68. GRUSZCZYNSKI, M., HOFFMAN, A., MALKOWSKI,K. and VEIZER, J., 1992. Seawater strontium isotopic perturbation at the Permian-Triassic boundary, West Spitsbergen and its implications for the interpretation of strontium isotopic data. Geology, 20: 779-782. HALLAM, A., 1989. The case for sea-level change as a dominant causal factor in mass extinction of marine invertebrates. Phil. Trans. R. Soc. London, B325: 437-455. HALLAM, A., 1990. The end-Triassic mass extinction event. Geol. Soc. Am., Spec. Pap., 247: 577-583. HEIMANN, A., FLEMING, T. n., ELLIOT, O. n. and FOLAND, K. A., 1994. A short interval of Jurassic continental flood basalt volcanism in Antarctica as demonstrated by 4~ geochronology. Earth Planet. Sci. Lett., 141: 19-41. HEUER, K., 1953. The End ofthe World. Rinehart & Co., New York, 220 pp. HICKEY, L. J., 1980. Paleocene stratigraphy and flora of the Clark's Fork Basin. Univ. Mich. Pap. Paleontol., 24: 33-49. HILDEBRAND, A. R. et al., 1991. Chicxulub Crater. A possible Cretaceous/Tertiary boundary impact crater on the Yucatan Peninsula, Mexico. Geology, 19:867-871. HOLSER, W. T. and MAGARITZ, M., 1992. Cretaceous/Tertiary and Permian/Triassic boundary events compared. Geochim. Cosmochim. Acta, 56: 3297-3309. HOLSER, W. T., SCHONLAUB, n. P., BOECKELMANN, K. and MAGARITZ, M., 1991. The PermianTriassic of the Gartnerkofel-1 Core (Carnic Alps, Austria). Synthesis and conclusions. In: W. T. HOLSER and H. P. SCHONLAUB (Editors), The Permian-Triassic boundary in the Carnic Alps of Austria (Gartnerkofel Region). Abhandlung Geol. Bundesanst., 45: 213-232. Hsu, K. J., 1980. Terrestrial catastrophe caused by cometary impact at the end of the Cretaceous. Nature, 285:201-203. Hso, K. J. and MCKENZIE J. A., 1985. A "strangelove" ocean in the earliest Tertiary. In: E. T. SUNDQUIST and W. S. BROECKER (Editors), Carbon Cycle and Atmospheric CO2, Natural Variations Archean to Present. Geophysical Monograph 32, Am. Geophys. Union, Washington, DC, pp. 487-492. Hsu, K. J. and MCKENZIE, J. A., 1990. Carbon-isotope anomalies at era boundaries; global catastrophes and their ultimate cause. Geol. Soc. Am., Spec. Pap., 247:61-70. Hsu, K. J., MCKENZIE,J. A. and HE, Q. X., 1982a. Terminal Cretaceous environmental and evolutionary changes. Geol. Soc. Am., Spec. Pap., 190:317-328. nsu, K. J., MCKENZIE, J. A., WEISsERT, n. et al., 1982b. Mass mortality and its environmental and evolutionary consequences. Science, 216: 249-256. HSU, K. J., OBERHANSLI,n. et al., 1985. "Strangelove ocean" before the Cambrian explosion. Nature, 316:809-811. HUT, P., ALVAREZ,W. et al., 1987. Comet showers as a cause of mass extinctions. Nature, 329:118-

140

References 126. JABLONSKI, D., 1986. Evolutionary consequences of mass extinctions. In: D. M. RAUP and D. JABLONSKI (Editors), Patterns and Processes in the History of Life. Springer-Verlag, Berlin, pp. 313-329. JAIPRAKASH, B. C., SINGH, J. and RAJU, D. S. N., 1993. Foraminiferal events across K/T boundary and age of Deccan volcanism in Palakollu area, Krishna-Godavari Basin, India. J. Geol. Soc. India, 41: 105-117. JOHNSON, A. L. A. and SIMMS, M. J., 1989. The timing and cause of Late Triassic marine invertebrate extinctions. Evidence from scallops and crinoids. In: S. K. DONOVAN(Editor), Mass Extinctions, Processes and Causes. Columbia University Press, New York, pp. 174--194. JOHNSON, K. R., 1992a. Leaf-fossil evidence for extensive floral extinction at the Cretaceous-Tertiary boundary, North Dakota, USA. Cretaceous Res., 13:91-117. JOHNSON, K. R., 1992b. High latitude deciduous vegetation and muted floral response to the Cretaceous-Tertiary boundary event in New Zealand. Geol. Soc. Am., Abstr. Prog., 24: A333. JOHNSON, K. R. and HICI~Y, L. J., 1990. Megafloral change across the Cretaceous/Tertiary boundary in the northern Great Plains and Rocky Mountains, U.S.A. Geol. Soc. Am., Spec. Pap., 247: 433444. JOHNSON, K. R., NICHOLS, D. J., ATTREPJR, M. and ORTH, C. J., 1989. High-resolution leaf-fossil record spanning the Cretaceous/Tertiary boundary. Nature, 340:708-711. JONES, D. S., MUELLER, P. A. et al., 1987. Biotic, geochemical and paleomagnetic changes across the Cretaceous/Tertiary boundary at Braggs, Alabama. Geology, 15:311-315. JONES, D. S. et al., 1991. Integrated geochronology of a classic Pliocene fossil site in Florida. Linking marine and terrestrial biochronologies. J. Geol., 99: 637-648. KAIHO, K., 1992. A low extinction rate of intermediate-water benthic foraminifera at the Cretaceous/Tertiary boundary. Mar. Micropaleont., 18: 229-259. KAJIWARA, Y. and KAIHO, K., 1992. Oceanic anoxia at the Cretaceous/Tertiary boundary supported by the sulfur isotopic record. Palaeogeog., Palaeoclimatol., Palaeoecol., 99:151-162. KASTING, J. F., RICHARDSON, S. M., POLLACK, J. B. and TOON, O. B., 1986. A hybrid model of the CO 2 geochemical cycle and its application to large impact events. Am. J. Sci., 286:361-389. KAUFFMAN, E. G., 1988. The dynamics of marine stepwise mass extinction. Rev. Esp. Paleontol., no. Extraordinario, 57-71. KELLER, G., 1986. Stepwise mass extinctions and impact events. Late Eocene to Early Oligocene. Mar. Micropaleontol., 10: 267-293. KELLER, G., 1988a. Biotic turnover in benthic foraminifera across the Cretaceous/Tertiary boundary at E1 Kef, Tunisia. Palaeogeog., Palaeoclimatol., Palaeoecol., 66: 153-171. KELLER, G., 1988b. Extinction, survivorship and evolution of planktic foraminifera across the Cretaceous/Tertiary boundary at E1Kef, Tunisia. Mar. Micropaleontol., 13: 239-263. KELLER, G., 1989a. Extended Cretaceous/Tertiary boundary extinctions and delayed population change in planktonic foraminifera from Brazos River, Texas. Paleoceanography, 4: 287-332. KELLER, G., 1989b. Extended period of extinctions across the Cretaceous/Tertiary boundary in planktonic foraminifera of continental-shelf sections. Implications for impact and volcanism theories. Geol. Soc. Am. BulL, 101: 1408-1419. KELLER, G. and BARRERA, E., 1990. The Cretaceous/Tertiary boundary impact hypothesis and the paleontological record. Geol. Soc. Am., Spec. Pap., 247: 563. KRASSILOV, V. A., 1975. Climatic changes in Eastern Asia as indicated by fossil floras. II. Late Cretaceous and Danian. Palaeogeog., Palaeoclimatol., Palaeoecol., 17: 157-172. KUMP, L., 1991. Interpreting carbon-isotope excursions. Strangelove oceans. Geology, 19: 299-302. KYTE, F. T., 1988. The extraterrestrial component in marine sediments. Description and interpretation. Paleoceanography, 3: 235-247. KYTE, F. T. and WASSON, J. T., 1986. Accretion rate of extraterrestrial matter. Iridium deposited 33 to 67 million years ago. Science, 232: 1225-1229. KYTE, F. T., ZHOU, Z. and WASSON, J. T., 1981. High noble metal concentrations in a late Pliocene sediment. Nature, 292:417-420. KYTE, F. T., ZHOU, Z. and WASSON, J. T., 1988. New evidence on the size and possible effects of a late Pliocene oceanic asteroid impact. Science, 241: 63-65. LEARY, P. N. and RAMPINO, M. R., 1990. A multi-causal model of mass extinctions. Increase in trace metals in the oceans. In: E. G. KAUFFMAN and O. H. WALLISER (Editors), Extinction Events in

141

Catastrophe: impact of comets and asteroids Earth History. Springer-Verlag, Berlin, pp. 45-55. LECHEMINANT, A. N. and HEAMAN, L. M., 1989. Mackenzie igneous events, Canada. Middle Proterozoic hotspot magmatism associated with ocean opening. Earth Planet. Sci. Lett., 96: 38-48. LEHMAN, T. M., 1987. Late Maastrichtian paleoenvironments and dinosaur biogeography in the western interior of North America. Palaeogeog., Palaeoclimatol., Palaeoecol., 60:189-217. LEWIS, J. S. et al., 1982. Chemical consequences of major impact events on Earth. Geol. Soc. Am, Spec. Pap., 190:215-222. LI., Z., ZHAN, L., YAO, J. and ZHOU, Y., 1991. On the Permian-Triassic events in South China- Probe into the end Permian abrupt extinction and its possible causes. Saito Ho-on Kai Special Publication No. 3, pp. 371-385. LOWRm, W., ALVAREZ,W. and ASARO, F., 1990. The origin of the White Beds below the CretaceousTertiary boundary in the Gubbio section, Italy. Earth Planet. Sci. Lett., 98:303-312. LYELL, C., 1830-33. Principles of Geology, Being an Attempt to Explain the Former Changes of the Earth's Surface by Processes Still in Operation, 3 Vols. John Murray, London, 1700 pp. LYONS, J. B. and OFFICER, C. B., 1992. Mineralogy and petrology of the Haiti Cretaceous/Tertiary section. Earth Planet. Sci. Lett., 109: 205-224. MACDOUGAL, J. D., 1988. Seawater strontium isotopes, acid rain and the Cretaceous-Tertiary boundary. Science, 239: 485-487. MACKENSEN, A. and EHRMANN,W. U., 1992. Middle Eocene through Early Oligocene climate history and paleoceanography in the Southern Ocean. Stable oxygen and carbon isotopes from ODP Sites on Maud Rise and Kerguelen Plateau. Mar. Geol., 108: 1-27. MACLEOD, N. and KELLER G., 1991a. Hiatus distributions and mass extinctions at the Cretaceous/Tertiary boundary. Geology, 19: 497-501. MACLEOD, N. and KELLER, G., 1991b. How complete are Cretaceous/Tertiary boundary sections? A chronostratigraphic estimate based on graphic correlation. Geol. Soc. Am. Bull., 103" 14391457. MAGARITZ, M., 1989. 13C minima follow extinction events. A clue to faunal radiation. Geology, 17: 337-340. MAGARrrz, M., BAR, R., BAUD, A. and HOLSER, W. T., 1988. The carbon-isotope shift at the Permian/Triassic boundary in the southern Alps is gradual. Nature, 331: 337-339. MAGARITZ, M., HOLSER, W. T. and KIRSCHVL~K, J. L., 1986. Carbon-isotope events across the Precambrian/Cambrian boundary on the Siberian Platform. Nature, 320: 258-259. MAHONEY, J. J., 1988. Deccan Flood Basalts. In: J. D. MACDOUGALL(Editor), Continental Flood Basalts. Kluwer, Dordrecht, pp. 151-194. MARGOLIS, S. V., CLAEYS, P. and KY'rE, F. T., 1991. Microtektites, microkrystites and spinels from a Late Pliocene asteroid impact in the Southern Ocean. Science, 251:1594-1597. MARGOLIS, S. V., MOUNT, J. F., DOEHNE, E., SHOWERS, W. and WARD, P., 1987. Late CretaceousTertiary boundary oxygen and carbon isotope stratigraphy, diagenesis and biotic extinctions at Zumaya, Spain. Paleoceanography, 2:361-377. MARVIN, U. B., 1990. Impact and its revolutionary implications for geology. Geol. Soc. Am., Spec. Pap., 247: 147-154. MAXWELL, W. D., 1989. The end Permian mass extinction. In: S. K. DONOVAN(Editor), Mass Extinctions. Processes and Evidence. Columbia University Press, New York, pp. 152-173. MAXWELL, W. D., 1992. Permian and Early Triassic extinction of non-marine tetrapods. Palaeontology, 35: 571-583. MCGHEE JR, G. R., 1988. The Late Devonian extinction event. Evidence for abrupt ecosystem collapse. Paleobiology, 14: 250-257. MCGHEE JR, G. R. et al., 1986. Late Devonian "Kellwasser Event" mass-extinction horizon in Germany. No geochemical evidence for a large-body impact. Geology, 14: 776-779. McKAY, C. P. and THOMAS, G. E., 1982. Formation of noctilucent clouds by an extraterrestrial impact. Geol. Soc. Am., Spec. Pap., 190:211-214. MCKENZIE, D. P. and BICKLE, M. J., 1988. The volume and composition of melt generated by extension of the lithosphere. J. Petrol., 29: 625-679. MCKINNEY, M. L., 1987. Taxonomic selectivity and continuous variation in mass and background extinctions of marine taxa. Nature, 325: 143-145. MCLAREN, D. J., 1970. Presidential address. Time, life and boundaries. J. Paleontol., 44: 801-815. MCLAREN, D. J., 1985. Mass extinction and iridium anomaly in the Upper Devonian of Western Aus-

142

References tralia. A commentary. Geology, 13:170-172. MCLAREN, D. J. and GOODFELLOW,W. D., 1990. Geological and biological consequences of giant impacts. Annu. Rev. Earth Planet. Sci., 18: 123-171. MCLEAN, D. M., 1978. A terminal Mesozoic "greenhouse": lessons from the past. Science, 201" 401406. MCLEAN, D. M., 1985. Deccan Traps mantle degassing in the terminal Cretaceous marine extinctions. Cretaceous Res., 6: 235-259. MELOSH, n. J., 1982. The mechanics of large meteoroid impacts in the Earth's oceans. Geol. Soc. Am., Spec. Pap., 190:121-128. MELOSH, n. J., 1989. Impact Cratering, A Geologic Process. Oxford, New York, 245 pp. MELOSH, n. J., SCHNEIDER,N. M., ZAHNLE, K. J. and LATHAM, O., 1990. Ignition of global wildfires at the Cretaceous/Tertiary boundary. Nature, 343:251-254. MILLER, K. G., FAIRBANKS,R. G. and MOUNTAIN,G. S., 1987. Tertiary oxygen isotope synthesis, sea level history and continental margin erosion. Paleoceanography, 2: 1-19. MILNE, D. H. and McKAv, C. P., 1982. Response of marine plankton communities to a global atmospheric darkening. Geol. Soc. Am., Spec. Pap., 190: 297-304. MONTANARI, A., 1990. Geochronology of terminal Eocene impacts; an update. Geol. Soc. Am., Spec. Pap., 247:607-616. NAPIER, W. M. and CLUBE, S. V. M., 1979. A theory of terrestrial catastrophism. Nature, 282: 455459. NICHOLS, D. J., JARZEN, D. M., ORTH, C. J. and OLIVER, P. Q., 1986. Palynological and iridium anomalies at Cretaceous-Tertiary boundary, south-central Saskatchewan. Science, 231: 714717. O'KEEFE, J. D. and AHRENS, T. J., 1982a. Impact mechanics of the Cretaceous-Tertiary extinction bolide. Nature, 298: 123-127. O'KEEFE, J. D. and AHRENS, T. J., 1982b. The interaction of the Cretaceous/Tertiary extinction bolide with the atmosphere, ocean and solid earth. Geol. Soc. Am., Spec. Pap., 190: 103-120. O'KEEFE, J. D. and AHRENS, T. J., 1989. Impact production of CO 2 by the Cretaceous/Tertiary extinction bolide and resultant heating of the Earth. Nature, 338: 247-249. OLSEN, P. E., FOWELL, S. J. and CORNET, B., 1990. The Triassic/Jurassic boundary in continental rocks of eastern North America; a progress report. Geol. Soc. Am., Spec. Pap., 247: 585-593. OPDYKE, B. N. and WILKINSON,B. H., 1988. Surface area control of shallow cratonic and deep marine carbonate accumulation. Paleoceanography, 3: 685-703. OPDYKE, B. N. and WILKINSON, B. H., 1990. Ocean temperature, saturation state and Phanerozoic/Holocene carbonate accumulation rates. Geol. Soc. Am. Abstr. Prog., 22: A131. ORTH, C. J., 1989. Geochemistry of the biD-event horizons. In: S. K. DONOVAN(Editor), Mass Extinctions. Processes and Evidence. Columbia University Press, New York, pp. 37-72. ORTH, C. J., ATTREPJR., M. and QUINTANA, L. R., 1990. Iridium abundance patterns across biD-event horizons in the fossil record. Geol. Soc. Am., Spec. Pap., 247: 45-59. ORTH, C. J., GILMORE, J. S., QUINTANA,L. R. and SHEEHAN, P. M., 1986. Terminal Ordovician extinction. Geochemical analysis of the Ordovician/Silurian boundary, Anticosti Island, Quebec. Geology, 14: 433-436. PIASECKI, S., 1984. Preliminary palynostratigraphy of the Permian-Lower Triassic sediments in Jameson Land and Scoresby Land, East Greenland. Bull. Geol. Soc. Den., 32:139-144. POLLACK, J. B. et al., 1982. Environmental effects of an impact generated dust cloud. Implications for the Cretaceous-Tertiary extinctions. Science, 219: 287-290. PRINN, R. G. and FEGLEY, B., 1987. Bolide impacts, acid rain and biospheric traumas at the Cretaceous/Tertiary boundary. Earth Planet. Sci. Lett., 88: 1-15. RAMPINO, M. R., 1987. Impact cratering and flood basalt volcanism. Nature, 327: 468. RAMPINO, M. R., 1992. A major Late Permian impact event on the Falkland Plateau. Los, Trans. Am. Geophys. Union, 73: 336. RAMPINO, M. R. and CALDEIRA, K., 1992. Antipodal hotspot pairs on the earth. Geophys. Res. Lett., 19:2011-2014. RAMPINO, M. R. and CALDEIRA, K., 1993. Major episodes of geologic change. Correlations, time structure and possible causes. Earth Planet. Sci. Lett., 114:215-227. RAMPINO, M. R. and HAGGERTY, B. M., 1994. Extraterrestrial impacts and mass extinctions. In: T. GEHRELS (Editor), Hazards Due to Comets and Asteroids. University of Arizona Press, Tucson,

143

Catastrophe: impact of comets and asteroids AZ.

RAMPINO, M. R. and SELF, S., 1984. Sulphur-rich volcanic eruptions and stratospheric aerosols. Nature, 310: 677-679. RAMPINO, M. R., SELF, S. and STOTHERS,R. B., 1988. Volcanic winters. Annu. Rev. Earth Planet. Sci., 16: 73-99. RAMPINO, M. R. and STOTHERS,R. B., 1984a. Terrestrial mass extinctions, cometary impacts and the Sun's motion perpendicular to the galactic plane. Nature, 308: 709-712. RAMPINO, M. R. and STOTHERS,R. B., 1984b. Geologic rhythms and cometary impacts. Science, 226: 1427-1430. RAMPINO, M. R. and STOTHERS, R. B., 1986. Geologic periodicities and the Galaxy. In: R. SMOLUCHOWSKI,J. N. BAHCALLand M. S. MATTHEWS(Editors), The Galaxy and the Solar System. University of Arizona Press, Tucson, AZ, pp. 241-259. RAMPINO, M. R. and STOTHERS,R. l . , 1988. Flood basalt volcanism during the past 250 million years. Science, 241: 663-668. RAMPINO, M. R. and VOLK, T., 1988. Mass extinctions, atmospheric sulphur and climatic warming at the K/T boundary. Nature, 332: 63-65. RAUP, D. M., 1979. Size of the Permo-Triassic bottleneck and its evolutionary implications. Science, 206:217-218. RAUP, D. M., 1990. Impact as a general cause of extinction; a feasibility test. Geol. Soc. Am., Spec. Pap., 247: 27-32. RAUP, D. M., 1991. A kill curve for Phanerozoic marine species. Paleobiology, 17: 37-48. RAUP, D. M., 1992. Extinctions. Bad Genes or Bad Luck? Norton, New York, 210 pp. RAYMO, M. E. et al., 1989. Late Pliocene variation in northern hemisphere ice sheets and North Atlantic deep water circulation. Paleoceanography, 4: 413-446. RENNE, P. R. and BASU, A. R., 1991. Rapid eruption of the Siberian Traps flood basalts at the PermoTriassic boundary. Science, 253:176-179. RENNE, P. R. et al., 1992. The age of Parana flood volcanism, rifting of Gondwanaland and the Jurassic-Cretaceous boundary. Science, 258: 975-981. RETALLACK, G. J., LEAHY, G. D. and SPOON, M. D., 1987. Evidence from paleosols for ecosystem changes across the Cretaceous/Tertiary boundary in eastern Montana. Geology, 15:1090-1093. ROBERT, C. and CHAMLEY, n., 1990. Paleoenvironmental significance of clay mineral associations at the Cretaceous-Tertiary passage. Palaeogeog., Palaeoclimatol., Palaeoecol., 79:205-219. ROCCHIA, R., BOCLET, D., BONTE, P., BUFFETAUT,E., ORUE-ETXEBARRIA,JAEGER, J-J. and JEHANNO, C., 1988a. Structure of the iridium anomaly at the K-T boundary of the Sopelana section (Spanish Basque country). C. R. Acad. Sci. Paris, 307, serie II: 1217-1223. ROCCHIA, R., BOCLET, O., COURTILLOT,V. and JAEGER, J.-J., 1988b. A search for iridium in the Deccan traps and inter-traps. Geophys. Res. Lett., 15:812-815. SAITO, Z., YAMANOI, Z. and KAIHO, K., 1986. End-Cretaceous devastation of terrestrial flora in the boreal Far East. Nature, 323: 253-255. SANCETTA, C., HEUSSER, L. and HALL, M. A., 1992. Late Pliocene climate in the Southeast Atlantic. Preliminary results from a multi-disciplinary study of DSDP Site 532. Mar. Micropaleontol., 20: 59-75. SANDBERG, C. A., ZIEGLER, W., DREESEN, R. and BUTLER, J. L., 1988. Late Frasnian mass extinction. Conodont event stratigraphy, global changes and possible causes. Cour. Forsch.-Inst. Senckenberg, 102: 263-307. SCHNEIDER, S. n. and THOMPSON, S. L., 1988. Simulating the climatic effects of nuclear war. Nature, 333: 221-227. SCHUBERT,J. K. and BOTTJER,D. J., 1992. Early Triassic stromatolites as post-mass extinction disaster forms. Geology, 20: 883-886. SCHULTZ, P. n. and GAULT, O. E., 1990. Prolonged global catastrophes from oblique impacts. Geol. Soc. Am., Spec. Pap., 247: 239-261. SEARS, J. W. and ALT, D., 1992. Impact origin of large intracratonic basins, the stationary Proterozoic crust and the transition to modern plate tectonics. In: M. J. BARTHOLOMEWet al. (Editors), Basement Tectonics 8, Kluwer, Dordrecht, pp. 385-392. SEBAI, A., FERAUD, G., BERTRAND, n. and HANES, J., 1991. 4~ dating and geochemistry of tholeiitic magmatism related to the early opening of the Central Atlantic rift. Earth Planet. Sci. Lett., 104: 455-472.

144

References SEPKOSKI JR., J. J., 1990. The taxonomic structure of periodic extinction. Geol. Soc. Am., Spec. Pap., 247: 33-44. SHACKLETON, N. J., BACKMAN, J. et al., 1984. Oxygen isotope calibration of the onset of ice-rafting and history of glaciation in the North Atlantic region. Nature, 307: 620-623. SHARPTON, V. L., DALRYMPLE,G. B. et al., 1992. New links between the Chicxulub impact structure and the Cretaceous/Tertiary boundary. Nature, 359:819-821. SHEEHAN, P. M. et al., 1991. Sudden extinction of the dinosaurs. Latest Cretaceous, Upper Great Plains, U.S.A. Science, 254: 835-839. SHOEMAKER, E. M., WOLFE, R. F. and SHOEMAKER,C. S., 1990. Asteroid and comet flux in the neighborhood of Earth. Geol. Soc. Am., Spec. Pap., 247: 155-170. SIGNOR III, P. W. and LIPPS, J. H., 1982. Sampling bias, gradual extinction patterns and catastrophes in the fossil record. Geol. Soc. Am., Spec. Pap., 190:291-296. SIGURDSSON, H., 1982. Volcanic pollution and climate: the 1783 Laki eruption. Los, Trans. Am. Geophys. Union, 63:601-602. SIGURDSSON, H., 1990. Assessment of the atmospheric impact of volcanic eruptions. Geol. Soc. Am., Spec. Pap., 247:99-110. SIGURDSSON, H., D'HONDT, S., ARTHUR, M. A. et al., 1991. Glass from the Cretaceous/Tertiary boundary in Haiti. Nature, 349: 482-487. SIGURDSSON, H., D'HONDT, S. and CAREY, S., 1993. The impact of the Cretaceous/Tertiary bolide on evaporite terrane and generation of major sulfuric acid aerosol. Earth Planet. Sci. Lett., 109: 543559. SIMMS, M. J., 1990. Climatic and biotic change in the Late Triassic. J. Geol. Soc., London, 147: 321327. SIMMS, M. J. and RUFFELL, m., 1990. Synchroneity of climatic change and extinctions in the Late Triassic. Geology, 17: 265-268. SMrr, J., 1990. Meteorite impact, extinctions and the Cretaceous-Tertiary boundary. Geol. Mijnbouw, 69:187-204. SMrr, J. and HERTOGEN, J., 1980. An extraterrestrial event at the Cretaceous-Tertiary boundary. Nature, 285: 198-200. SM1T, J., MONTANARI, A. et al., 1992. Tektite-bearing, deep-water clastic unit at the CretaceousTertiary boundary in northeastern Mexico. Geology, 20:99-103. STANLEY, S. M., 1984. Marine mass extinctions. A dominant role for temperature. In: M. H. NITECKI (Editor), Extinctions. University of Chicago Press, Chicago, IL, pp. 69-117. STANLEY, S. M., 1988. Paleozoic mass extinctions. Shared patterns suggest global cooling as a common cause. Am. J. Sci., 288: 334-352. STANLEY, S. M., 1992. An ecological theory for the origin of Homo. Paleobiology, 18: 237-257. STANLEY, S. M. and CAMPBELL,L. D., 1985. Neogene mass extinction of western Atlantic molluscs. Nature, 293: 457-459. STORETVEDT, K. M., MrrCHELL, J. G., ABRANCHES, M. C., MAALOE, S. and ROBUIN, G., 1992. The coast-parallel dolerite dykes of east Madagascar; age of intrusion, remagnetization and tectonic aspects. J. Afr. Earth Sci., 15: 237-249. STOTHERS, R. B., 1993. Flood basalts and extinction events. Geophys. Res. Lett., 20: 1399-1402. STOTHERS, R. B., WOLFE,J. A., SELF, S. and RAMPINO, M. R., 1986. Basaltic fissure eruptions, plume heights and atmospheric aerosols. Geophys. Res. Lett., 13: 725-728. STOTT, L. D. and KENNETT, J. P., 1989. New constraints on early Tertiary palaeoproductivity from carbon isotopes in foraminifera. Nature, 342: 526-529. STo'rr, L. D. and KENNETT, J. P., 1990. The paleooceanographic and paleoclimatic signature of the Cretaceous/Paleogene boundary in the Antarctic. Stable isotope results from ODP Leg 113. Proc. Ocean Drilling Prog., Sci. Results, 113: 829-848. SWISHER III, C. C. and PROTHERO, D. R., 1990. Single-crystal 4~ dating of the EoceneOligocene transition in North America. Science, 249: 760-762. THIERSTEIN, n. R., 1980. Selective dissolution of late Cretaceous and earliest Tertiary calcareous nannofossils. Experimental evidence. Cretaceous Res., 2:165-176. THIERSTEIN, n. R., 1981. Late Cretaceous nannoplankton and the change at the Cretaceous-Tertiary boundary. Soc. Econ. Paleont. Mineral., Spec. Publ., 32: 355-394. THIERSTEIN, H. R., 1982a. The terminal Cretaceous extinction event and climatic stability. In: W. H. BERGER and J. C. CROWELL (Editors), Climate in Earth History. National Academy Press, Wash-

145

Catastrophe: impact of comets and asteroids ington, DC, pp. 90-96. THIERSTEIN, H. R., 1982b. Terminal Cretaceous plankton extinctions. A critical assessment. Geol. Soc. Am., Spec. Pap., 190: 385-399. THORDARSON, TH. and SELF, S., 1993. The Laki (Skaftar Fires) and Grimsvotn eruptions in 1783-85. Bull. Volcanol., 55: 233-263. THUNELL, R. C., 1981. Cenozoic palaeotemperature changes and planktonic foraminiferal speciation. Nature, 289: 670-672. TINus, R. W. and RODDY, D. J., 1990. Effects of global atmospheric perturbations on forest ecosystems in the Northern Temperate Zone; predictions of seasonal depressed-temperature kill mechanisms, biomass production and wildfire soot emissions. Geol. Soc. Am., Spec. Pap., 247: 77-86. TIWARI, R. S. and VIJAYA, 1992. Permo-Triassic boundary on the Indian peninsula. In: W. C. SWEET et al. (Editors), Permo-Triassic Events in the Eastern Tethys. Cambridge University Press, Cambridge, pp. 37-45. TOON, O. B. et al., 1982. Evolution of an impact-generated dust cloud and its effects on the atmosphere. Geol. Soc. Am., Spec. Pap., 190:187-200. TRIPOLI, G. J. and THOMPSON, S. L., 1988. A three-dimensional numerical simulation of the atmospheric injection of aerosols by a hypothetical basaltic fissure eruption. Global Catastrophes in Earth History, Abstracts. Lunar and Planetary Institute, Houston, TX, pp. 200-201. TURCO, R. P. et al., 1981. Tunguska meteor fall of 1908. Effects on stratospheric ozone. Science, 214: 19-23. TURCO, R. P., TOON, O. B., ACKERMAN,T. P., POLLACK,J. B. and SAGAN, C., 1983. Nuclear winter. Global consequences of multiple nuclear explosions. Science, 222:1283-1292. TURCO, R. P., TOON, O. B., ACKERMAN, T. P., POLLACK, J. B. and SAGAN, C., 1990. Climate and smoke. An appraisal of nuclear winter. Science, 247: 166-176. UPCHURCH, G. R., 1989. Terrestrial environmental changes and extinction pattems at the CretaceousTertiary boundary, North America. In: S. K. DONOVAN(Editor), Mass Extinctions, Processes and Evidence. Columbia University Press, New York, pp. 195-216. UREY, n. C., 1973. Cometary collisions and geological periods. Nature, 242: 32-33. VANDAMME, D., COURTILLOT,V., BESSE, J. and MONTIGNY, R., 1991. Paleomagnetism and age determinations of the Deccan traps (India). Results of a Nagpur-Bombay traverse and review of earlier work. Rev. Geophys., 29: 159-190. VALENTINE, J. W., 1986. The Permian-Triassic extinction event and invertebrate developmental modes. Bull. Mar. Sci., 39:607-615. VENKATESAN, M. I. and DANE, J., 1989. Organic geochemical evidence for global fires at the Cretaceous/Tertiary boundary. Nature, 338: 57-60. VICKERY, A. M. and MELOSH, n. J., 1990. Atmospheric erosion and impactor retention in large impacts, with applications to mass extinctions. Geol. Soc. Am., Spec. Pap., 247: 289-300. VISSCHER, n. and BRUGMAN, W. A., 1988. The Permian-Triassic boundary in the Southern Alps. A palynological approach. Mere. Soc. Geol. Ital., 34: 121-128. WANG, K., 1992. Glassy microspherules (microtektites) from an Upper Devonian Limestone. Science, 256: 1547-1550. WANG, K., ORTH, C. J., ATTREPJR, M., CHATTERTON,B. D. E., nou, n. and GELDSETZER,n. n. J., 1991. Geochemical evidence for a catastrophic biotic event at the Frasnian/Famennian boundary in south China. Geology, 19: 776-779. WANG, K., ORTH, C. J., ATTREPJR, M., CHATTERTON,B. O. E., WANG, X. and LI, J-J., 1993. The great latest Ordovician extinction on the South China Plate. Chemostratigraphic studies of the Ordovician-Silurtian boundary interval on the Yangtze Platform. Palaeogeogr., Palaeoclimatol., Palaeoecol., 104: 61-79. WANG, J-X., CHAI, Z-F. et al., 1992. Osmium abundance and its correlation with iridium at some geological boundaries. Chin. Sci. Bull., 37: 1984-1987. WATERHOUSE,J. B., 1973. The Permian-Triassic boundary in New Zealand and New Caledonia and its relationship to world climatic changes and extinction of Permian life. In: A. LOGAN and L. V. HILLS (Editors), The Permian and Triassic Systems and their Mutual Boundary. Canadian Society of Petroleum Geologists, Mem. 2, pp. 445-464. WEEMS, R. E., 1992. The "terminal Triassic catastrophic extinction event" in perspective, a review of Carboniferous through Early Jurassic terrestrial vertebrate extinction pattems. Palaeogeog., Palaeoclimatol., Palaeoecol., 94: 1-29.

146

References WHISTON, W., 1696. A New Theory of the Earth, From its Origin to the Consummation of All Things. R. Roberts, London, 635 pp. WHITE, R. and MCKENZIE, D., 1989. Magmatism at rift zones. The generation of volcanic continental margins and flood basalts. J. Geophys. Res., 94: 7685-7729. WIDMARK, J. G. V. and MALMGREN, B., 1992. Benthic foraminiferal changes across the CretaceousTertiary boundary in the deep sea; DSDP Sites 525, 527 and 465. J. Foraminiferal Res., 22: 81113. WILDE, P., BERRY, W. B. N., QUINBY-HUNT, M. S. et al., 1986. Iridium abundances across the Ordovician-Silurian stratotype. Science, 233: 339-341. WISE, K. P. and SCHOPF,T. J. M., 1981. Was marine faunal diversity in the Pleistocene affected by changes in sea level? Paleobiology, 7: 394-399. WOLBACH, W. S., GILMOUR, I., ANDERS, E., ORTH, C. J. and BROOKS, R. R., 1988. Global fire at the Cretaceous-Tertiary boundary. Nature, 334: 670. WOLBACH, W. S., LEWIS, R. S. and ANDERS, E., 1985. Cretaceous extinctions. Evidence for wildfires and search for meteoritic material. Science, 230:167-170. WOLFE, J. A., 1987. Late Cretaceous-Cenozoic history of deciduousness and the terminal Cretaceous event. Paleobiology, 13:215-226. WOLFE, J. A., 1990. Palaeobotanical evidence for a marked temperature increase following the Cretaceous/Tertiary boundary. Nature, 343: 153-156. WOLFE, J. A., 1991. Palaeobotanical evidence for a June 'impact winter' at the Cretaceous/Tertiary boundary. Nature, 352: 420-423. WOLFE, J. A. and UPCHURCH JR, G. R., 1986. Vegetation, climatic and floral changes at the Cretaceous-Tertiary boundary. Nature, 324: 148-152. WOLFE, J. A. and UPCHURCH JR, G. R., 1987. North American nonmarine climates and vegetation during the Late Cretaceous. Palaeogeog., Palaeoclimatol., Palaeoecol., 61: 33-77. Xu, D., MA, S., CHAI, Z, MAO, X., SUN, Y., ZHANG, Q. and YANG, Z., 1985. Abundance variation of iridium and trace elements at the Permian/Triassic boundary at Shangsi in China. Nature, 314: 154-156. YABUSHITA, S. and ALLEN. A. J., 1989. On the effect of accreted interstellar matter on the terrestrial environment. Mon. Not. R. Astron. Soc., 238: 1465-1478. ZACHOS, J. C. and ARTHUR, M. A., 1986. Paleoceanography of the Cretaceous/Tertiary boundary event. Inferences from stable isotopic and other data. Paleoceanography, 1: 5-26. ZACHOS, J. C., ARTHUR, M. A. and DEAN, W. E., 1989a. Geochemical and paleoenvironmental variations across the Cretaceous/Tertiary boundary at Braggs, Alabama. Palaeogeog., Palaeoclimatol., Palaeoecol., 69: 245-266. ZACHOS, J. C., ARTHUR, M. A. and DEAN, W. E., 1989b. Geochemical evidence for suppression of pelagic marine productivity at the Cretaceous/Tertiary boundary. Nature, 337:61-64. ZACHOS, J. C., BREZA, J. R. and WISE, S. W., 1992. Early Oligocene ice-sheet expansion on Antarctica. Stable isotope and sedimentological evidence from Kerguelen Plateau, southern Indian Ocean. Geology, 20: 569-573. ZAHNLE, K. J., 1990. Atmospheric chemistry by large impacts. Geol. Soc. Am., Spec. Pap., 247:271288. ZAHNLE, K. and GRINSPOON, D., 1990. Comet dust as a source of amino acids at the Cretaceous/Tertiary boundary. Nature, 348:157-160. ZHOU, L. and KYTE, F. T., 1988. The Permian-Triassic boundary event: a geochemical study of three Chinese sections. Earth Planet. Sci. Lett., 90:411-421.

147

This Page Intentionally Left Blank

II. Observations of Climate and Future Projection

This Page Intentionally Left Blank

Chapter 5

Observations from the surface" projection from traditional meteorological observations P. D. JONES

Introduction Although meteorological recording began earlier, the period since the middle of the 19th century is that traditionally associated with instrumental records. This chapter examines in detail the reliability of temperature, precipitation and sea-level pressure data, concentrating on the construction of the hemispheric and global average temperature series. The key piece of observational evidence in the "global warming debate" is the "global" temperature series (JONES and WIGLEY,1990). How much has the temperature risen and how certain are we of the causes? The remainder of the chapter considers these issues, together with an assessment of the record in a longer paleoclimatic context (see also Chapter 6 by DIAZ and KILADIS) and the representativeness of the longer surface record with respect to free-atmosphere observations since the late 1950s. The chapter concludes by considering possible explanations of the instrumental record (for more discussion of the possible explanations see also Chapter 4 by RAMPINO, Chapter 6 by DIAZ and KILADIS, Chapter 9 by WANG et al., Chapter 10 by ANDREAE, Chapter 11 by BRASSEUR et al. and Chapter 12 by HENDERSON-SELLERS)and indicates what each suggests concerning the course of climate change in the next century.

Development of instrumentation and meteorological observational networks The instrumental period of weather recording began in the late 17th century with the development of barometers and thermometers. Up until the late 18th century, the early observers were dedicated amateurs, mostly physicians, parsons and country squires and their equivalents in countries other than England. Careful analyses are necessary to enable the early records to be used. Calibration scales and the exposure of the early instruments must be scrutinized with, in some cases, experiments performed to duplicate the conditions (CHENOWETH, 1992, 1993). Long records have been developed for some European regions (e.g. MANLEY (1974) and PARKER et al. (1992) for Central England; DETTWILLER (1981) for Paris and SCHAAKE (1982) for Berlin). There was some correspondence and liaison between the early observers through societies such as the Royal Society in London, but little central organisation. Many of the original data have been lost and are only available in society proceedings. Sometimes the entire records of early observers have been published (e.g. KINGTON(1988a) for Lyndon in England). The first real weather observing network was set up in 1778 by the Socigtd Royale de

151

Observations from the surface: projection from traditional meteorological observations M~dicine in France under the patronage of Louis XVI. The French meteorologist Louis Cotte was appointed to establish and maintain the network of observers. Meanwhile in Mannheim, then the capital of the Palatinate of the Rhine, the Elector Karl Theodor founded the Societas Meteorologica Palatina in 1780. In both cases detailed instructions on how and when to record information were sent to observers. In the German case, standard instruments were also issued. The networks, however, did not survive the change in the political climate in Europe in the late 18th and early 19th centuries and both ceased in the 1790s. KINGTON (1988b) has produced, from the station data, weather maps for most of Europe for the 1780s. Further expansion of the meteorological recording network began following the Vienna Meteorological Congress of 1873. This meeting provided the impetus for the expansion of recording networks worldwide and for the international exchange of data that continues today under the auspices of the World Meteorological Organization (WMO). When national meteorological centres were established in the late 19th century, observing networks expanded, instruments and instrument shelters were standardised, and uniform instructions to observers were issued. As a result, a reasonably comprehensive global network of temperature, precipitation and pressure recording stations emerged, many of which have continued in operation to the present. Unfortunately, standardisation of instruments and exposure did not extend to a universally adopted policy so that, even today, vastly different protocols are followed. The effects of this are discussed in the remainder of this section. Further discussion on the history of instrumentation has been presented by LAMB and JOHNSON (1966), VON RUDLOFF(1967) and MIDDLETON(1966, 1969). Land areas represent only 30% of the Earth's surface and it is important to consider marine weather observations. Both merchant and naval vessels have taken observations of weather, including measurements of the temperature of the sea surface (SST), since the 1850s. Data were stored in log books by the major maritime nations following an international agreement signed in Brussels in 1853. The force behind this was an American naval captain, Matthew Fontaine Maury. In the last 20 years, major international efforts have been made to put all this log-book information into computer data banks. Two such compilations exist, the Comprehensive Ocean-Atmosphere Data Set (COADS) produced at Boulder, Colorado

(WOODRUFFet al., 1987) and the UK Meteorological Office marine data bank (BOTTOMLEY et al., 1990). About 80 million individual observations of SST exist between 1854 and 1990. Homogeneity

of temperature

records

In this section we define homogeneity according to CONRAD and POLLAK (1962), who call a numerical series representing the variations of a climatological element homogeneous if the variations are caused only by variations of weather and climate. It is vital that all meteorological observations contain as few non-meteorological variations as possible. Although observers may take readings with meticulous care, other influences can easily affect the readings. The following sections discuss the most important causes of non-homogeneity.

152

-

changes in instruments, exposure and measurement technique;

-

changes in station location;

-

changes in observation times and the methods used to calculate monthly averages;

-

changes in the station environment with particular reference to urbanisation.

Development of instrumentation and meteorological observational networks Changes in instrumentation, exposure and measuring techniques For land-based temperature, the effects of changes in thermometers have been slight so far this century. More important to the long-term homogeneity is the exposure of the thermometers. Readings now are generally from thermometers located in louvered screened enclosures, which are generally painted white. Earlier readings often were from shaded wall locations, sometimes under roof awnings. These need corrections which can only be estimated by parallel measurements or by reconstructing past locations (CHENOWETH,1992, 1993). The greatest potential discontinuity in temperature time series is currently being induced in first order stations in the United States. Here mercury-in-glass thermometers have been replaced by a thermistor installed in a small louvered housing (model HO-83). The change was made by the National Weather Service in the USA at most first order sites during early 1982. There appears to have been little thought given to effects on the homogeneity of the record. The continuity of the record for climatic purposes was only a minor consideration of the weather service. Parallel measurements were not made as is generally recommended. The first signs of a problem were noticed by the frequent breaking of records for extreme daily maximum temperatures. A warm bias was clearly evident, possibly being indicative of an aspiration problem in the replacement instrumentation (GALL et al., 1992; KESSLER et al., 1993). As more developed countries might follow the lead of the United States, the problem is likely to recur. However, with good design and adequate testing through parallel observations, the problems induced by any changes should be surmountable. Marine temperature data are seriously affected by homogeneity problems. SST data were generally taken before World War II by collecting some seawater in an uninsulated canvas bucket and measuring the temperature. As there was a short time between sampling and reading, the water in the bucket cooled slightly due to evaporation. Since World War II most readings are made in the intake pipes which take water on to ships to cool the engines. The change appears to have been quite abrupt, occurring within a month during December 1941. The sources of ships' observations in the two marine data banks (COADS and UKMO) changed from principally European to mainly American at this time (see also FOLLAND and PARKER, 1995). Studies of the differences between the two methods indicate bucket temperatures are cooler than intake ones by 0.3-0.7~ (JAMESand Fox, 1972). A correction technique for bucket measurements has been derived based on the physical principles related to the causes of the cooling (BOTTOMLEY et al., 1990; FOLLAND and PARKER, 1995). The cooling depends on the prevailing meteorological conditions, and so depends on the time of year and on location. Although a day-to-day phenomenon, the various influences on the bucket are basically linear, so cooling amounts can be calculated on a monthly basis. Assuming that there has been no major change in the seasonal cycle of SSTs over the last 100 years, the time between sampling and reading can be estimated as that which best minimises the residual seasonal cycle in the pre-World War II data. The estimate of between 3 and 4 min is both quite consistent spatially and agrees with instructions given to marine observers (BOTToMLEY et al., 1990; FOLLAND and PARKER, 1991, 1995). Coverage of SST data is generally dictated by shipping routes and is, therefore, poor over the Southern and tropical oceans. From the late 1970s, it has been possible to include infor-

153

Observations from the surface: projection from traditional meteorological observations mation for all oceanic areas using satellite measurements. Co-located in-situ and satellite values do not agree quantitatively because the satellite is measuring the skin temperature and the in-situ observations, the average temperature of the top few metres. However, the anomaly patterns of the two are in excellent agreement and a blending technique (REYNOLDS and MARSlCO, 1993) has been developed to anchor the globally complete satellite patterns to the in-situ observations which act as the ground truth. The blending technique also makes use of sea-ice extents, which are also determined from satellite measurements. It will be some time, however, before there is a long record of complete SST fields to assess the significance of the much sparser coverage, prior to the late 1970s, for hemispheric and large scale averages.

Changes in station locations It is rare that a temperature recording site will have remained in exactly the same position throughout its lifetime. Information concerning station moves is of primary importance to homogenisation. Such information about a station's history is referred to as metadata. Assessment of station homogeneity requires nearby station data and appropriate metadata.

Changes in observation time and the methods used to calculate monthly averages In the late 19th and early 20th centuries, there was considerable discussion in the climatological literature concerning the best method of calculating true daily and, hence, monthly average temperatures (e.g. ELLIS, 1890; DONNELL, 1912; HARTZELL, 1919; RUMBAUGH, 1934). Many schools of thought existed and, unfortunately, no one system prevailed in all regions. At present, there is no uniform system, although the majority of nations use the average of the daily maximum and minimum temperatures to estimate monthly mean temperatures. The remainder use a variety of formulae based on observations at fixed hours, some of which are designed to simulate the true daily mean that would be estimated from readings every hour. There would be no need to correct readings to a common standard, however, provided the observing systems had remained internally consistent. This is because analysts are generally interested in relative changes rather than the absolute values. Unfortunately, only in a few countries has the methodology remained consistent since the 19th century (BRADLEY et al., 1985). If absolute temperatures are important for some study then a common standard will be necessary. In the routine exchange of data, countries exchange monthly mean temperatures, using whatever observational practice is applied in the country. Proposals currently before WMO seek to change the number of meteorological variables routinely exchanged. When this takes place, maximum and minimum temperatures will be exchanged and it is likely these will become the preferred method of calculating mean temperature. Due to data limitations, studies of past temperature variations have had to concentrate on mean temperature. Maximum and minimum temperatures will be extremely useful when considering both the causes and impacts of climate changes (see later). Using maximum and minimum temperature, however, to calculate monthly means is extremely susceptible to the times at which the observations are made (BAKER, 1975; BLACKBURN, 1983). In the United States, KARL et al. (1986) have developed a model to correct observations made in the

154

Development of instrumentation and meteorological observational networks morning and afternoon to the 24-h day ending at midnight. The corrections to the monthly means are called time-of-observation biases.

Changes in the station environment The growth of towns and cities over the last 200 years around observing sites can seriously impair the temperature record. Affected sites will appear warmer than adjacent rural sites. The problem is serious at individual sites but appears relatively small in averages calculated for large continental-scale areas (JONES et al., 1990). A correction procedure has been developed for the contiguous United States by KARL et al. (1988). It depends on population size. Population is generally used as a ready surrogate for urbanisation but other indices such as paved area (OKE, 1974), although more useful, are only available for specific cases. The individual nature of the effect means that the corrections applied to US data by KARL et al. (1988) will only be reliable on a regional rather than a site basis. The corrections are also specific to North America and would not apply to Europe, for example, where urban development has taken place over a much longer period of time. The problem is likely to become important in the future in developing countries as cities expand rapidly (OKE, 1986).

Precipitation and pressure homogeneity Both of these variables are also affected by severe homogeneity problems. For precipitation, there are problems associated with the gauge size, shielding, height above the ground and the growth of vegetation near the gauge. All can impair the performance of the gauge, affecting the efficiency of the catch. The biggest problem concerns the measurement of snowfall, where continual attempts to improve gauge performance affect the long-term homogeneity of the station records. RODDA (1969) has reviewed the possible errors and inhomogeneities that can result in precipitation records. The total of these problems, coupled with the greater spatial variability of precipitation data, makes the problem of homogeneity much more difficult to deal with than for temperature. Precipitation networks are rarely dense enough to allow for the homogeneity checks discussed in the next section to be undertaken. Instead of studying individual records, analyses of regional precipitation series have been made (BRADLEYet al., 1987a; DIAZ et al., 1989; FOLLAND et al., 1990, 1992). Precipitation climatologies for land regions have been developed (LEGATESand WILLMOTT, 1990; HULME, 1992) partly for the evaluation of the performance of General Circulation Models (GCMs) (see Chapter 7 by DICKINSON). The Hulme climatology is based on the 1951-1980 period and has values for each month of the period allowing model variability as well as model precipitation totals to be compared. The Legates and Willmott climatology uses all available precipitation data and is not geared to any specific time period. This means that temporal variability variations cannot be considered. These authors have, however, attempted to allow for gauge undercatch using correction methods devised by SEVRUK (1982). Taken globally they estimate that precipitation is underestimated by about 10% varying from 40% near the poles in winter to less than 5% at the tropics. In 1986, the World Climate Research Programme (WCRP) established the Global Precipitation Climatology Project (GPCP). The aim of GPCP is to produce a global 10-year precipi-

155

Observations from the surface: projection from traditional meteorological observations tation data set for the decade 1986-1995, 2.5~ 2.5 ~ box resolution. Year-to-year estimates over ocean areas have never been attempted before (RUDOLF et al., 1991). Initial results for 1987 have so far been published (RUDOLF et al., 1992). Recently WCRP has extended the project indefinitely. The GPCP will integrate various satellite and gauge measurements of precipitation and should provide a truer representation of the spatial and temporal variability of precipitation than previously achieved (ARKIN and ARDANUY, 1989). Various precipitation climatologies, including ground-based data, satellite precipitation estimates and model precipitation forecasts have been compared by JANOWIAK (1992) over the tropics. Differences are sometimes large, although the intercomparisons made by SPENCER (1993), between the new estimates from microwave-sounding unit MSU (see section on Upper air data during the last 40 years for a description of this source) data and the satellite infrared index are very encouraging, indicating that the GPCP aims will be achieved, even if the initially published data have to be revised. Pressure data have an advantage over the temperature and precipitation databases because they are routinely analysed on to regular grid networks for weather forecasting purposes. Monthly mean data for the Northern Hemisphere (N of 15~ extend back to 1873 and for the Southern Hemisphere routine analyses began in the early 1950s. Neither analyses are complete for the entire period, particularly during the earliest years. The sources of the data are discussed and their quality assessed in a number of papers (Northern: WILLIAMSand VAN LOON, 1976; TRENBERTH and PAOLINO, 1980; JONES, 1987; Southern: VAN LOON, 1972; KAe,OLY, 1984; JONES, 1991). Global analyses have been produced at the major operational weather centres since 1979. The most widely available analyses are those from the European Centre for Medium-Range Weather Forecasts (ECMWF) and the National Meteorological Center (NMC) of NOAA. However, because of changes to the model and to data assimilation schemes, the homogeneity of the analyses for many climate applications is questionable (TRENBERTH,1992). A consistent set of analyses will soon be produced as part of the World Climate Research Programme (WCRP) re-analysis project at ECMWF and NMC (WCRP, 1993).

Data homogenisation techniques Many methods have been proposed for testing the homogeneity of station records relative to those at adjacent stations (CONRADand POLLAK, 1962; WMO, 1966). Generally, tests involve the null hypothesis that a time series of differences (ratios for precipitation) between neighbouring observations will exhibit the characteristics of a random series (e.g. MITCHELL, 1961; CRADDOCK, 1977; POTTER, 1981; ALEXANDERSSON, 1986; JONES et al., 1986a,c; YOUNG, 1993). The methods work with or without metadata. The scope of the method and the ability to explain and correct inhomogeneities is dramatically improved, however, with detailed metadata. The most comprehensive homogeneity exercise performed produced the United States Historic Climate Network (KARLand WILLIAMS,1987).

What surface observations tell us about the last 140 years

The longest and most spatially extensive record of surface variations is available for three meteorological variables: temperature, precipitation and mean sea-level pressure (MSLP).

156

What surface observations tell us about the last 140 years

Long records often exist for other variables such as sunshine duration, cloudiness, humidity, wind speeds etc., but these are rarely "global" in extent and these variables can have severe problems related to long-term homogeneity due to changing instrumentation (e.g. sunshine recorders; KARL and STEURER, 1990) or to changing practices (e.g. clouds; HENDERSON-

SELLERS, 1990). Coverage for the three principal variables is, however, never truly global. Even now, in-situ observations are rare over the Southern Oceans and routine observations are only available over Antarctica since 1957. Improvements are being made and since the late 1970s it has been possible to integrate satellite and surface measurements.

Surface temperature Hemispheric averages

Various research groups have synthesised the individual site temperature series into gridded and/or hemispheric average time series. The three main groups recently involved are the Climatic Research Unit/UK Meteorological Office (CRUAJKMO) (JONES, 1988; FOLLAND NORTHERN HEMISPHERE LAND+MARINE TEMPERATURES

0.5 0.0 -0.5 -1.0 I

a

I

I

I

i

I

I

I

I

I

l

I

o

I

I

I

I

I

i

I

i

I

I

I

i

I

I

I

i

!

i

I

I

I

o

I

o

I

o

0.5 0.0 -0.5

0.5 0.0

-0.5 Summer I

I

0.5 0.0 -0.5 Autumn

-1.0 t

J

I

i

I

J

I

I

I

I

I

!

I

!

18;0

'

18;0

'

19()0

'

1920

'

1940

'

19;50

'

19~}0

'

0.5 0.0 -0.5

Fig. 1. Northern Hemisphere surface air temperatures for land and marine areas by season, 18541994. Standard meteorological seasons for the Northern Hemisphere are used: winter is December to February, dated by the year of the January. Data are expressed as anomalies from 1950 to 1979. Time series in this and subsequent plots have been smoothed with a 10-year Gaussian filter.

157

Observations from the surface: projection from traditional meteorological observations et al., 1990, 1992; JONES et al., 1991), the Goddard Institute for Space Studies (HANSEN and LEBEDEFF, 1987, 1988) and the State Hydrometeorogical Institute in St. Petersburg

(VINNIKOVet al., 1990). The differences between the groups arise both from the data used and from the method of interpolation. Intercomparison of the results of the three groups can be found in various publications (e.g. ELSAESSER et al., 1986; FOLLAND et al., 1990, 1992; JONES et al., 1991; ELSNER and TSONIS, 1991a). Here we consider only CRU/UKMO because they are the only group to incorporate marine data. Time series of the mean hemispheric annual and seasonal temperature anomalies are shown in Figs. 1 and 2. Both series have been widely discussed (WIGLEY et al., 1986; FOLLAND et al., 1990, 1992; JONES and WIGLEY, 1990). The accuracy of the individual seasonal and annual estimates undoubtedly varies with time. The most reliable period is from about 1950. Various techniques have been tried to assess error estimates, particularly for years prior to the 1920s. Frozen grid analyses (JONES et al., 1985, 1986a,b, 1991; BOTTOMLEY et al., 1990) indicate that hemispheric estimates can be well estimated from the sparse data available during the late nineteenth and early twentieth centuries. Individual annual estimates appear to have about twice the variability of the post 1950 data but the average temperature anomaly during the late 19th century is well approximated by the sparse network. This reSOUTHERN HEMISPHERE LAND+MARINE TEMPERATURES

0.0 -0.5

0.0 -0.5 Autumn

0.5 0.0 -0.5 Winter

0.0 -0.5 Spring

0.0 -0.5 Annual 1860

1880

1900

1920

1940

1960

1980

Fig. 2. Southern Hemisphere surface air temperatures for land and marine areas by season, 18541994. Standard meteorological seasons for the Southern Hemisphere are used: summer is December to February, dated by the year of the January. Data are expressed as anomalies from 1950 to 1979.

158

What surface observations tell us about the last 140 years sult has been confirmed by more sophisticated techniques which also use GCM results to estimate the additional uncertainty associated with the lack of data for the Southern Oceans (TRENBERTH et al., 1992; MADDEN et al., 1993). GUNST et al. (1993) estimate uncertainty using optimal statistical sampling techniques. In Figs. 1 and 2 warming begins in the hemispheric series around the turn of the century in all seasons. Only in summer in the Northern Hemisphere are the 1980s not the warmest decade. In this season, temperatures during the 1860s and 1870s were of a comparable magnitude to the most recent decade. In the Northern Hemisphere, interannual variability is greater in winter compared to summer. No such seasonal contrast is evident in the Southern Hemisphere. The hemispheric annual time series and their average (the global series) (Fig. 3) represent the major observational evidence in the "global warming" debate. Before moving to explanations of the past record (see section on Explanations of the instrumental temperature record) and to projections of it into the future (see section on Projections to the future), it is essential to consider both the spatial and diurnal nature of the changes and the longer-term context within which the instrumental record is perceived (section on The last 140 years in a longer term context), and temperature change throughout the atmospheric column (section on Upper air data during the last 40 years). Spatial patterns of change Although hemispheric and global series are heavily relied upon as important indicators of past climatic change, they mask the spatial details of temperature variability across the globe. Annual time series for 11 regions are given in WMO's report on the Global Climate

0.5

i

,

~

0.0 --0.5 ~ I 0.5

Northern Hemisphere

0.0 --0.5 0.5

I

~

I

I

I

I

I

I

I

0

I

I

I

I

I

i

I

i

I

i

I

i

I

i

I

i

I

i

0.0 --0.5 1860 1880 1900 1920 1940 1960 1980 Fig. 3. Global and hemispheric annual surface air temperatures for land and marine areas, 1854-1994. Data are expressed as anomalies from 1950 to 1979.

159

Observations from the surface: projection from traditional meteorological observations

(

:

tO

0.5

~

~3

Fig. 4. Seasonal temperature anomalies for the Northern Hemisphere, 1981-1990. The anomalies are based on the 1950-1979 period. The contour interval is 0.5~ with negative anomalies dashed. Shaded areas indicate regions with insufficient data.

System (WMO and UNEP, 1993). The various curves highlight different levels of high-tolow frequency variability in each region. On a priori grounds there are only weak arguments to suggest why any regional series should be more indicative of global scale change than any other of comparable size. The issue has been addressed by a number of workers (see e.g. BARNETT, 1978; JONES and KELLY, 1983; FOLLAND et al., 1990, 1992; JONES and BRIFFA, 1992; BRIFFA and JONES, 1993; PARKER et al., 1994). Here we consider the spatial pattern of warmth during the 1981-90 decade. Seasonal temperature anomalies for the Northern and Southern Hemispheres, relative to the 1950-1979 reference period, are plotted in Fig. 4 (NH) and Fig. 5 (SH). Annual temperature anomalies are shown in Fig. 6. Most areas of the globe experienced above normal (1950-1979) temperatures for the 1981-1990 decade, although there was great variability from season to sea-

160

What surface observations tell us about the last 140 years

0S

M

os

Fig. 5. Seasonal temperature anomalies for the Southern Hemisphere, 1981-1990. The anomalies are based on the 1950-1979 period. The contour interval is 0.5~ with negative anomalies dashed. Shaded areas indicate regions with insufficient data.

son in the Northern Hemisphere (Fig. 4). Much of the decadal warmth was apparent during the DJF and MAM seasons, with the warmth located over northern and central Asia and over northwestern North America. During JJA, anomalies were smaller in magnitude over North America and negative over parts of Asia. During SON, temperatures were below normal over western Canada and Alaska but slightly above normal over Siberia. Over northern Africa temperatures were above normal during MAM, JJA and SON but below normal in DJF. Major areas of relatively consistent below normal temperatures included the central North Pacific Ocean and the northern North Atlantic region encompassing Iceland and Greenland. Over the Southern Hemisphere (Fig. 5) anomalies were generally smaller in magnitude except over parts of Antarctica. Almost all of the hemisphere was warmer than normal. The

161

Observations from the surface: projection from traditional meteorological observations

Fig. 6. Annual temperature anomalies for the Northern and Southern Hemispheres. The anomalies are based on the 1950-1979 period. The contour interval is 0.5~ with negative anomalies dashed. Shaded areas indicate regions with insufficient data.

major exception to this was a part of eastern Antarctica during MAM and SON. In tropical and temperate latitudes Australia, the Indian Ocean and the eastern equatorial Pacific were warmer than normal. The regions with the greatest warming over Antarctica were the Australian sector in JJA, SON and DJF and the Antarctic Peninsula in MAM and JJA. The annual anomaly maps for the two hemispheres (Fig. 6) obviously encompass the average features of the four seasons. Over the Northern Hemisphere relative warmth prevailed over most of Eurasia, especially Siberia, the western half of North America and western Africa. Temperatures were cooler during the decade over the North Pacific Ocean and the northern Atlantic region surrounding Greenland and Iceland. Most regions of the Southern Hemisphere have experienced warmth during the decade except for parts of eastern Antarctica, the Amazon Basin and part of the southwestern Pacific. Warmth during the decade was

162

What surface observations tell us about the last 140 years

greatest over the Indian Ocean, southern Africa, the eastern equatorial Pacific Ocean, the Antarctic Peninsula and the Australian sector of Antarctica. Although this decade has clearly been the warmest recorded since the beginning of the instrumental era, it is clear that the warming has not been global in extent. The term "global warming" therefore, is something of a misnomer. The warmth has also varied seasonally, being more significant in the boreal spring and winter and less in the boreal autumn and summer. Assessment of the spatial patterns of change is particularly important when considering possible causal factors for the changes. Diurnal temperature change

Most of the focus of past and future temperature change has been concerned with mean temperature. For past changes, this has resulted from a lack of readily available databases of monthly mean maximum and minimum temperatures. These data, since November 1994, are routinely exchanged between countries. For models, the diurnal cycle in the control runs is often poorly simulated. Few GCM experiments have so far considered maximum and minimum temperatures and the associated diurnal range. The issue is, however, particularly important with regard to climate change impact studies. The impacts of a change will, for example, be very different if the change affects only minimum temperatures rather than affecting both minimum and maximum equally. Recently developed data sets (KARL et al., 1993) have permitted analyses of maximum and minimum temperatures to be made for 37% of the global land mass (encompassing the contiguous United States, Canada, Alaska, the former Soviet Union, China, Japan, the Sudan, South Africa and eastern Australia). The analyses indicate (Fig.7) that over the 1952-1989 period, minimum temperatures have risen at a rate three times that of maximum temperatures. The reduction in the diurnal temperature range is approximately equal to the temperaI

Y

I

0.5 0.0 -0.5 -1.0 1.0 0.5 0.0 -0.5 -1.0 0.0 -0.5 -1.0

Diurnel Temperature Range 1960

1970

1980

Fig. 7. Annual maximum, minimum temperatures and the diurnal temperature range for 1952-1989.

163

Observations from the surface: projection from traditional meteorological observations ture increase. The changes are detectable in all of the regions studied and in all seasons. Extension of the analyses to other regions is planned over the next few years. Results for New Zealand (SALINGERet al., 1993) suggest less organised changes compared to those for the major northern mid-latitude land masses seen by KARL et al. (1993). The causes of the differential increases in maximum and minimum temperatures are closely related to changes in cloudiness. Increases in cloud cover have been reported from Europe, Canada and Australia (HENDERSON-SELLERS, 1986, 1989; JONES and HENDERSON-SELLERS, 1992, respectively) although there has been marked station-to-station variability. Although decreasing diurnal temperature range is a characteristic signal of increasing urbanisation, irrigation usage and desertification, the widespread nature of the response and the lack of any seasonal bias suggests a large-scale climatic forcing. Possible factors are anthropogenic increases in sulphate aerosols and/or greenhouse gases. Neither factor is entirely convincing as more regional variability in the changes might be expected with sulphate aerosols (CHARLSONet al., 1991) while, for greenhouse gases, GCMs indicate a much smaller level of response (RIND et al., 1989; CAO et al., 1992). Natural variability is always a candidate for the explanation of the changes. Some long individual European records reveal large changes in diurnal temperature range with the recent decline being a relatively minor feature. With increasing data availability, studies in this area will become increasingly important in the issue of the detection of climatic change (see section on The next century).

Precipitation Analyses of precipitation data have tended to focus on regional rather than hemispheric scales. The reasons for this stem from two factors: precipitation data have much greater spatial variability compared to temperature and precipitation data from oceanic areas are virtually non-existent. For most areas of the world, precipitation variability, both in space and time, is so large that it will be generally difficult to detect trends statistically until after significant impacts have been felt in the agricultural and water resource sectors. Various regional time series were developed for the IPCC Scientific Assessments (FOLLAND et al., 1990, 1992). All of the studied regions in North America, Eurasia, Africa and Australia show large year-to-year variability in either annual or growing-season precipitation totals (see also WMO and UNEP, 1993). All show marked decadal-scale variation but with little long-term change on the century time-scale. There are two major exceptions to this. Average annual series for the former Soviet Union show a gradual increase in precipitation since the beginning of this century (Fig.8). Much of the increase has occurred in the non-summer season. Some of the increase may be related to the underestimation of snowfall during winter but even allowing for this a significant increase over the former USSR south of 70~ is still evident (GROISMAN et al., 1991). The most significant and dramatic changes in precipitation have occurred over the Sahel region of sub-Saharan Africa. Here the reduction of precipitation during the rainy season in recent years has been highly statistically significant. At some sites in the region there has been a decline of about 30% between averages for the two standard WMO periods of 19311960 and 1961-1990. The steep change would be even greater if the periods 1941-1970 and 1971-1990 were compared as the 1960s were relatively wet. Figure 9 plots the average time series for the region using the two standard WMO reference periods. Using the most recent

164

What surface observations tell us about the last 140 years

2 "E "~ -6

I

E o ~

0

~3

~-I

8 ~

_

- 3 ~......__L~L_..__J~~___L____J~ 1860

1880

1900

1920

1940

1960

1980

2000

yeor

Fig. 8. Standardised regional annual precipitation anomaly time series for the territory of the former Soviet Union. Standardisation of each station time series is achieved by subtracting the reference period precipitation and dividing by the standard deviation. The regional series is the average of all available station series. No form of areal weighting is used. The reference period used for calculating the annual means and standard deviations was 1951-1980. 1940

1860 1880 1900 1920 -~ r----r T r----r ~ I I_ Sahel (I0-15~ 1900-1992 2 t with r e s p e c t t o 1 9 3 1 - 6 0 m e a n

1960

1980

2000

/

1

0

o

o "ID

-2

.~ "o

k_ o tO 01

-3

2

t-

a~

1

0 -1

smoothed ~

1860

1

~

~

1880

~

values

1

1900

I

1920

I

1940

i_......._J.........__J~l..._.----J

1960

1980

t

2000

Fig. 9. Standardised regional (annual) precipitation anomaly time series for the Sahel regions (1015ON; 20~ Two reference periods 1931-1960 and 1961-1990 are shown for comparison of the results.

165

Observations from the surface: projection from traditional meteorological observations period changes the appearance of the time series from the "Sahel drought" of the last 20years (relative to 1931-1960) to the "Sahel wet period" of the 1920s to the 1960s (relative to 1961-1990). The difference between the two analyses is not a simple change in level because of the use of standardised anomalies. Interannual variability is greater at most stations during the 1961-1990 period compared to 1931-1960. The reasons for the decline in rainfall have been widely discussed and a variety of causes have been postulated (LAMB, 1978; LOUGH, 1986; NICHOLSON, 1989). Rainy season (June-September) totals for the region appear to be partially predictable from the previous seasons' (March-May) sea surface temperatures in the Atlantic Ocean (FOLLAND et al., 1986; ROWELL et al., 1992). Global precipitation data sets have been developed recently, principally to assess the performance of GCMs (see section on Precipitation and pressure homogeneity). Assessment is achieved by comparing spatial patterns of model (control) and real-world precipitation climatologies (HULME, 1991). GCM climatologies, however, with 2x CO2 levels (perturbed) indicate an enhanced hydrological cycle with between a 7 and 15% increase in global total precipitation (MITCHELL et al., 1990). A reason for analysing global-scale precipitation, therefore, is to see if any increase is occurring. Detection of future climate change is developed in more detail later. Although large-scale precipitation series are a poor detection variable, because of greater uncertainties in both observed and model data, the impacts of climate change will be principally felt through changes in precipitation. Runoff, for example, is much more sensitive to precipitation than temperature (WIGLEYand JONES, 1985). Hemispheric- and global-scale precipitation have been studied in two papers (BRADLEY et al., 1987a; DIAZ et al., 1989). In both studies, the non-normal distribution of precipitation and its large spatial variability were accounted for by transforming all seasonal and annual precipitation totals using the gamma distribution. With this, precipitation totals are scaled between 0 and 1, median values that occur half the time having a value of 0.5. The distribution was fitted over the 1921-1960 period, and for each station, seasonal and annual estimates of the scale and shape parameters of the distribution were made. Large scale averages were calculated after the transformed precipitation data were interpolated onto an equal-area grid. For the Northern Hemisphere (BRADLEY et al., 1987a), the most interesting finding was a decline in precipitation during the last 40 years over subtropical areas (5-35~ and an increase in mid-latitudes (35-70~ These trends are consistent with many other analyses and are partially accounted for by the Sahel and the former Soviet Union trends discussed earlier. For the NH as a whole, there is considerable decade-to-decade variability but little overall trend. Northern summer and autumn seasons show a slight decline whereas winter and spring seasons show an increase. These trends are in accord with hemispheric temperature changes, providing evidence for a more intense hydrological cycle when average temperatures are higher. The Southern Hemisphere analyses (DIAZ et al., 1989) indicate increases in precipitation in both subtropical and temperate latitudes. The rise in precipitation is primarily due to increases in autumn and spring over South America. A slight increase is evident over Australasia and no change over southern Africa. The changes in precipitation in both hemispheres are broadly consistent with a number of GCM results with doubled CO2 levels (MITCHELL et al., 1990). The consistencies, however, are only on the largest of scales and do not survive scrutiny on small spatial and seasonal scales.

166

What surface observations tell us about the last 140 years

!

I

!

I

i

I

!

,

,

1

i

Azores

30

I

w

I

!

i

I

i

-Iceland

25 20 15 10 i

i

i

i

i

I

i

I

i

I

1880 1900 1920 1940 1960 1980

Fig. 10. Winter (November-March) pressure difference (mb) between Ponta Delgada, Azores and Stykkisholmur, Iceland. Year dated by the January. Higher values indicate stronger westerlies over the North Atlantic.

Pressure

The interrelationships of the climate system mean that past changes in the patterns of temperature and precipitation patterns will undoubtedly also be accompanied by changes in circulation throughout the troposphere. Changes that have been noted range from local-scale studies such as the decline in westerlies over the British Isles (LAMB, 1972) to much larger regional scale indices such as the North Atlantic Oscillation (VANLOON and ROGERS, 1978), North Pacific winter pressure (TRENBERTH, 1990) and various Southern Hemisphere indices (including the Southern Oscillation) (KAROLY, 1995). The North Atlantic/European region has probably been studied more than other parts of the world, almost certainly because of the availability of long (>150 years) records. The wellknown out-of-phase relationship between winter temperatures over northern Europe (especially Fennoscandia) and Greenland has long been recognised (VAN LOON and ROGERS, 1978). The link between surface temperature anomalies in the region and the circulation of the North Atlantic region is evident in indices such as the North Atlantic Oscillation (NAO). Figure 10 shows seasonal (November-March) values of the NAO (difference in pressure between Iceland and the Azores). Stronger westerlies prevailed during the first two decades of the 20th century and again during the 1970s and 1980s. The weakest westerlies occurred during the 19th century and the 1960s. In the North Pacific region a major change in the strength and position of the Aleutian Low took place during the mid-1970s. After about 1977 the strength of the low pressure centre

i

,

1014

North

,

|

'

!

'

i

,

Pocific

1012 1010 1008 1006 t

I

1920

i

I

1940

i

i

19t50

i

1980

Fig. 11. Winter (November-March) pressure average (mb) for the North Pacific 30-70~ l 50~ Year dated by the January.

167

150~

Observations from the surface: projection from traditional meteorological observations increased and the centre moved slightly south particularly in winter (Fig. 11). In many respects the change may have been a return to the conditions that prevailed during the 1920s, 1930s and the early 1940s. Some of the change since the 1970s has reversed in the most recent years. The effects of the change in the mid-1970s in the climate of the whole N. Pacific region have been discussed by numerous workers (e.g. DOUGLAS et al., 1982; NITTA and YAMADA, 1989; TRENBERTH, 1990). The decreased pressures led to greater advection of warm maritime air on to the northwest of North America leading to enhanced precipitation and warmer temperatures. Alaska has experienced the greatest regional warmth over the Northern Hemisphere with average decadal temperatures I~ warmer than the 1951-1980 period during the 1980s. In contrast sea temperatures have been much cooler than normal in the central North Pacific and cooler air temperatures have been experienced over Japan and the Okhotsk/Kamchatka region. Although the region only explains part of the hemispheric increase, the change in climate of the region, particularly in the winter, has been abrupt. Some of the cause of the extratropical circulation change may be tropical in origin due to a lack of any significant La Nifia (see PHILANDER 1985 and Chapter 6 by DIAZ and KILADIS) events (TRENBERTH, 1990). Modelling studies support the tropical Pacific link (GRAHAM, 1995).

The last 140 years in a longer term context

The purpose of this section is to address, partly, the question of how much of the rise of temperatures can be explained by natural fluctuations. In addressing this issue we need to consider what is known about climatic variations of the last 1,000 years. Prior to the development of instruments, climatic information must be inferred from various naturally recording phenomena as well as from written historical records. The reconstruction of past climates is generally referred to as paleoclimatology. Numerous indicators exist, e.g. trees, ice cores, historical records, corals, varves, glacier movements, etc. (see BRADLEY, 1985 for a complete discussion of the methods and their strengths and weaknesses). In the following sections the changes in climate inferred, from the proxy climate sources, for the last 1,000 years are surveyed (see WILLIAMS and WIGLEY, 1983; JONES and BRADLEY 1992; BRADLEY and JONES, 1992a, 1993; see Chapter 6 by DIAZ and KILADISfor more details). Over the last 1,000 years three main episodes are evident in many, but not all, proxy records from the Northern Hemisphere. Although less well studied, Southern Hemisphere evidence from South America, Australia and New Zealand suggests that these main events also occurred there. The first episode is a cool period spanning approximately the 9th and 10th centuries. Following this a warm period, the "Medieval Warm Period", occurred which was at its maximum during the early 12th century A.D. After this period, climate entered the period known as the "Little Ice Age". This period was marked by cooler conditions between the 14th and 19th centuries, although not always. It was probably cool from 1300 to 1450 and again from 1550 to 1850 (WILLIAMS and WIGLEY, 1983). A general warming of the climate from the 10th to the 12th century is evident from both North America and Europe. By the 12th century, trees were living beyond their present limits in Alaska and the Yukon. In Europe winter temperatures were generally mild. Conditions in Greenland were best for human exploitation during the 10th and 1 lth centuries and this

168

Upper air data during the last 40 years allowed the establishment of two main Norse settlements in western and southern Greenland with combined populations of over 5,000 people (McGOVERN, 1981). The ensuing cooler conditions are confirmed by glacier advances in North America and Europe during the 14th, 16th and 17th centuries. The advances in different regions were never synchronous, but they point to cooler conditions for about 300 years interspersed with a few warmer episodes in different regions at different times (GROVE, 1988). Paintings of glaciers in Switzerland

(ZUMBUHL, 1976) show evidence of cooler summers, with the glaciers reaching their maximum extent in the late 17th century. Tree ring evidence in Switzerland and Scandinavia indicates similar timings. Hill farms were abandoned in many areas of upland Europe. It is impossible to say when the coldest periods were globally, because of different seasonal timings of the coldest phases in different parts of the world. From an analysis of 16 "summer-responding" proxy records in the middle-to-high latitudes of the Northern Hemisphere BRADLEY and JONES (1993) conclude that the coldest conditions of the last 560 years were between 1570 and 1730 and in the 19th century. Longer single-site and composite instrumental records confirm the annual warming from the late 19th century. The long European records, however, show that the 1880s were the coldest decade, at least since 1700, so part of the warming since then may reflect this unusually low starting point.

Upper air data during the last 40 years There are two methods for measuring temperatures at different levels in the atmosphere. These are the conventional radiosonde network (ANGELL, 1988) and the microwavesounding unit (MSU) on the TIROS-N series of satellites (SPENCER and CHRISTY, 1990 and see also Chapter 7 by DICKINSON). The conventional network extends back to 1958 and the MSU data to 1979. Before comparing the two sets of analyses with each other and with the surface, their advantages and limitations are discussed. Both sets of information are used routinely in global scale analyses undertaken at several operation centres (e.g. ECMWF, NMC, World Meteorological Centre-C at Melbourne), but changes to the operational models and data assimilation schemes mean that the homogeneity of long-time series (as for surface pressures also, see above) must be questionable (TRENBERTH, 1992). In the following it is preferable to use the original station and satellite observations rather than the analysed fields. There are many potential problems even with the original station measurements due to changes in sonde design and the timing of observations (see GAFFEN, 1994; PARKER and Cox, 1994) The conventional radiosonde network began in the late 1940s, but an important instrument alteration in the mid-1950s and a change of observation timing means that the most consistent long series begin in 1958 (ANGELL, 1988). Extension back prior to 1958 may be possible but would be limited spatially to the mid-to-high latitudes of the Northern Hemisphere because of data availability. The most widely studied data set is that produced by Angell and Korshover (see ANGELL, 1988) using a network of 63 stations spread as evenly over the world as data availability allows. The data are available seasonally from 1958 for seven zones of the world, each the average of the stations in each zone. Using up to four times as many stations as ANGELL (1988) for the northern polar zone, KAHL et al. (1993) have confirmed the earlier results. The spatial representativeness of the Angell network on a global

169

Observations from the surface: projection from traditional meteorological observations scale has been examined by TRENBERTH and OLSON (1991). They compared complete global analysed fields from the ECMWF for 1979-1987 and found that correlations between the two were quite high, although root-mean-square errors for Angell's data were high in zones with fewer stations particularly the Southern Extratropics. The advantage of the 63 site network is that averages are easy to compute. The disadvantages are that the network is sparse and only the very largest of spatial features will be measured. Recently OORT and LIU (1993) have interpolated monthly radiosonde data onto a regular latitude/longitude grid from over 800 radiosonde sites. Although there is a considerably greater amount of data used, agreement with the Angell data set is extremely high (FOLLAND et al., 1992). In the following comparisons we use the Angell data set because it has been more widely used. Two levels are used, lower troposphere (850-300 mb) and lower stratosphere (100-50 mb). The MSU data are direct satellite measurements of microwave emission of radiation from space due to molecular oxygen in the atmosphere (SPENCERand CHRISTY, 1990). The satellite measures four channels but most analyses have concentrated on channels 2 (lower troposphere) and 4 (lower stratosphere). Because channel 2 is contaminated to some extent by variations in the upper troposphere, a modification placing greater emphasis on the lower troposphere, channel 2R was developed (SPENCERand CHRISTY, 1992b). This is used here and in other related studies (FOLLAND et al., 1992). The advantages of MSU are a single instrument and global coverage which means lower standard errors than conventional radiosonde data (SPENCERand CHRISTY, 1992a,b). The major disadvantages are that the series is short (1979 to the present) and is developed from a number of instruments on different satellites. Although overlaps enable corrections to be made, these may slightly affect the long" term homogeneity of the record.

Comparisons of lower tropospheric data sets Comparisons of the data sets have been made by a number of other studies (WIGLEYet al., 1985, 1986; FOLLAND et al., 1990, 1992; TRENBERTH et al., 1992). This section is focussed upon the period since 1979. Correlations between the three data sets are shown in Table I, on the monthly, seasonal and annual time-scales. ANGELL's (1988) data are only available for the traditional seasons and his annual series is for a December to November year. Correlations are always higher for the Northern than for the Southern Hemisphere, and higher the greater the averaging period. MSU2R data correlations with Land/Marine, Land Only and 850-300 mb data have similar values increasing slightly for each time-scale. Lower correlations between MSU2R and the surface over oceanic areas (which dominate the SH record) have been noted by TRENBERTH et al. (1992). The highest correlations of all are found between Land Only and Land/Marine data, but this is to be expected as the Land Only data set is the most variable part of the combined data set. Monthly time series of the Land/Marine and MSU2R data are shown in Fig. 12 for the 1979-1992 period. The interhemispheric correlations for MSU2R are clearly higher than for Land/Marine (see Table II). Variations of MSU2R between the two hemispheres are highly correlated on all time-scales. At the surface, the variability is clearly greater over the Northern Hemisphere. The marked cooling of the records during 1991 is discussed in the next section and in the section on External forcing of the climate.

170

Upper air data during the last 40 years TABLE I INTERCORRELATIONS BETWEENRADIOSONDE (850-300MB), M S U 2 R AND SURFACE DATA (1979--1992)

Time period

Correlation pair

NH

SH

Global

Monthly (164) d

MSU2R v LM a MSU2R v Land b LM v Land MSU2R v LM MSU2R v Land MSU2R v 850-300 c LM v Land LM v 850-300 Land v 850-300 MSU2R v LM MSU2R v Land MSU2R v 850-300 LM v Land LM v 850-300 Land v 850-300

0.62 0.58 0.96 0.69 0.63 0.76 0.98 0.67 0.65 0.84 0.81 0.89 0.99 0.93 0.92

0.45 0.56 0.85 0.57 0.67 0.74 0.83 0.51 0.63 0.59 0.74 0.88 0.97 0.63 0.74

0.62 0.62 0.94 0.70 0.70 0.80 0.96 0.69 0.70 0.79 0.81 0.93 0.99 0.89 0.89

Seasonal (52) d

Annual (13) d

a LM, land + marine (JONES et al., 1991) b Land, land only (JONES et al., 1986a,b) c 850-300, 850-300mb (ANGELL, 1988) d No. of cases.

Comparisons of lower stratospheric data sets FOLLAND et al. (1990, 1992) have performed a number of comparisons between MSU channel 4 (MSU4) data and conventional radiosonde data from ANGELL (1988) and OORT and LIU (1993). Table III gives correlations between MSU4 data (SPENCER and CHRISTY, 1993) and the 100-50 mb data from ANGELL (1988). These correlations are slightly higher than the similar comparisons for the lower troposphere. Figure 13, which shows the monthly time series of MSU4 data for the two hemispheres, clearly suggests a different pattern between the two hemispheres. Interhemispheric correlations are significantly lower than for the lower troposphere (see Table II). The warming and cooling episodes in the lower stratosphere are related to the increase in aerosol loading resulting from the E1 Chich6n and Mt. Pinatubo volcanic eruptions (see section on External forcing of the climate).

TABLE II INTERHEMISPHERIC CORRELATIONSFOR RADIOSONDE (850-300MB AND 100-50MB), MSU2R, M S U 4 AND SURFACE DATA (ABBREVIATIONSAS IN TABLE I)

MSU2R LM Land 850-300 mb MSU4 100-50 mb

171

Monthly (164)

Seasonal (52)

Annual (13)

0.58 0.26 0.22 0.39 -

0.76 0.47 0.45 0.56 0.54 0.67

0.89 0.64 0.74 0.71 0.78 0.85

Observations from the surface: projection from traditional meteorological observations 1980 1982 1984 1986 1988 1990 1992 1994 1.0 0.5 0.0 -0.5 -1.0 l,

i

,

i

,

i

,

i

,

i

,

i

,

i

''Northern ~ H e m i s p l ~ ~e ' r e

1.0

,

i

,

i

,

i

,

i

,

i

i

i

,

i

,

i

,

i

,

i

Lower tl~ I 1Tro' . ~IIpos' ILp,J4,#o" h'.e.;e' ~ '(1'98:;~--'1;9i) ~, I~.~ ' ' J

0.5 0.0 -0.5 -1.0

i

: : ',

1.0

:

: : ', : : : ', : ', I = I = I :

' : I ', I : : : : I I o

Southern Hemisphere Lond + Morine (1950-1979)

0.5 0.0 -0.5 f,

1.0 0.5 0.0

|

,

|

,

i

i

i

|

!

,

|

,

i

,

|

,

|

,

|

,

!

,

i

|

i

i

i

,

i

,

i

,

|

,

Southern Hemisphere Lower Troposphere (1982-1991)

t...................................t

-0.5 ,

i

,

i

,

i

i

i

,

i

,

i

i

i

,

i

|

i

,

i

i

I

i

i

i

i

i

i

,

i

,

i

i

i

1980 1982 1984 1986 1988 1990 1992 1994 Fig. 12. Hemispheric monthly averages of land and marine temperatures (surface; data from Figs. 1 and 2) and MSU2R data (lower troposphere; SPENCERand CHRISTY, 1992b). The time intervals given in each panel are the reference intervals from which anomalies have been calculated.

Explanations of the instrumental temperature record Explanations of changes in hemispheric- and global-mean climate over the past 100 years or so have concentrated on temperature variations. Temperature is the most reliably measured variable, in terms of accuracy of regional averages, and the most widely studied. Changes in precipitation and pressure, particularly on the regional scale, have undoubtedly occurred and variations in all three are strongly interrelated. It might be suspected that precipitation totals over global land-areas have increased over the 20th century but we are not as confident of the data as we are for the 0.5~

rise in global temperatures. Despite the fact that precipita-

tion variations have a greater impact on human activities than variations in other variables, TABLE III INTERCoRRELATIONS BETWEEN RADIOSONDE (100-50MB) AND M S U 4 DATA (ABBREVIATIONSAS IN TABLE I)

172

Time period

Correlation pair

NH

SH

Global

Seasonal (52) Annual (13)

MSU4 v 100-50 MSU4 v 100-50

0.89 0.93

0.78 0.97

0.86 0.99

Explanations of the instrumental temperature record

1980 1982 1984 1986 1988 1990 1992 1994 2.0

NorthernHemisphere Lower Stratosphere (1982-1991)

1.5 1.0 0.5 0.0 -0.5 -1.0 2.0 1.5 1.0 0.5 0.0 -0.5 -1.0 1980 1982 1984 1986 1988 1990 1992 1994

Fig. 13. Hemispheric monthly averages of MSU4 data (lower stratosphere; SPENCERand CHRISTY, 1992b). The time intervals given in each panel are the reference intervals from which anomalies have been calculated. we concentrate here on global-mean temperature, which is a fundamental measure of the state of the climate system. All variations in the historic and geologic past are considered as cold or warm epochs. The possible causal factors affecting global-mean temperature are conveniently grouped as either "internal" or "external" (ROBOCK, 1978). Internal factors are those which must be considered even in the absence of potentially larger changes in external forcing. Factors include natural changes in planetary albedo resulting from changes in cloudiness or in surface characteristics and changes in the circulation of the atmosphere and/or the ocean. The latter factor determines the lower boundary conditions for the atmosphere and the rates of heat exchange between the ocean and atmosphere. The former factors determine the spatial distribution of heat sources and sinks, as well as the horizontal and vertical fluxes of sensible and latent heat. Man-induced surface albedo changes should strictly be classified as external but, although their effects may be important at the regional scale, they are unlikely to be at the global scale. The main external factors are changes in the Sun's output or luminosity, changes in the fraction of shortwave solar radiation reaching the troposphere (due to injection of particulate matter and sulphate aerosols in the stratosphere following volcanic eruptions), changes in the reflectivity of clouds due to industrial activity and changes in the vertical distribution of outgoing longwave radiation flux in the troposphere due to increasing concentrations of greenhouse gases. Internal forcing of the climate It is likely that most of the interannual variability shown in Figs. 1-3 is the result of natural variability, even though the precise causes cannot be quantified. On slightly longer timescales, such as 2-8 years (but still classed as natural variability as it results from internal forcing), the ocean circulation and related changes in sea surface temperatures in the tropical Pacific Ocean associated with the E1 Nifio/Southern Oscillation (ENSO) phenomenon

(RASMUSSEN and CARPENTER, 1982; PHILANDER, 1983; see also Chapter 6 by DIAZ and

173

Observations from the surface: projection from traditional meteorological observations KILADIS) have a noticeable effect on hemispheric and global temperatures. This phenomenon, along with the similar long-term trends, is responsible for much of the high annual correlation ( r = 0.79, 1901-1990) between the two hemispheres. Many warm years relative to the filtered curves have been shown to be E1 Nifio or "warm event" years (BRADLEY et al., 1987b; VAN LOON and SHEA, 1985). In contrast many cold years relative to the filtered curve have been previously classified as La Nifia (PHILANDER, 1985) or "cold event" years. Even during the last 12 years the warm and cold event years can be seen in the MSU2R record (see Fig. 12). Using the Southern Oscillation Index (SOI), the pressure difference between Tahiti and Darwin, (see ROPELEWSKI and JONES, 1987; ALLAN et al., 1991; and Chapter 6 by DIAZ and KILADIS for more details), JONES (1989) has removed the ENSO influence from the ,annual hemispheric and global temperature series. The SOI explains about 25-30% of the high frequency (

i

i

i

i

~

1

120W

1

I

60W

0

60E

120E

180

Fig. 5. The global distribution of the annual mean aerial run-off determined by PEIXOTOand OORT (1983) together with corresponding stream lines. Each barb represents 2 g kg-i m s-1.

ally averaged distributions of water vapour transport although as shown in Fig. 5, they were able to deduce global distributions of the mean aerial run-off which support reasonable expectations regarding the sources and sinks of moisture. Using global data assimilation analyses for a single year (FGGE), CHEN (1985) found that the stationary non-divergent mode shown in Fig. 6(a) describes most of the total water vapour transport. The stationary divergent mode composed of a combination of local Hadley circulation and the local Walker circulation, is responsible for the maintenance of high regional values of the water vapour content in equatorial regions. In mid-latitudes, the transient divergent mode shown in Fig. 6(b) is an important component of the poleward transient water vapour transport in the storm tracks. Problems associated with difficulties in moisture analysis, especially over the Southern Oceans, suggest some caution is necessary in assessing the reliability of these computations of water vapour transport. Further analysis is needed over a longer time period. It is interesting, however, to note that many of the models used in AMIP show quite large interannual variations in the stationary divergent mode of water vapour transport in the tropics which is strongly linked to the E1 Nifio-Southern Oscillation (ENSO). An example from the BMRC model is shown in Fig. 7. Large

scale

interactions

Teleconnections across the globe are a fundamental component of the climate system. Canonical examples include the ENSO phenomenon (PHILANDER, 1983, 1990); the PacificNorth American (PNA) pattern (HOREL and WALLACE, 1981; HOSKiNS and KAROLY, 1981; WALLACE and GUTZLER, 1981) and the Madden Julian Oscillation (MADDEN and JULIAN,

297

Dynamics

of future

climates

9, (DECEMBER-FEBRUARY) 90N

-

'

'1

"

I

....

I

'

-'I

I

I

I

I

_ _ . _ _ _ ~ _ _I _ ~ 1

I

II

9

60N

.

. ~ ' .

30N

',

,

.

.

.

.

.

~.t~."=r

.

.

.

.

.

: : "_2 L L ~ 2 _ 2 ~

........

9

~ ~~. . *t. . . , ~9. . . , . .

.

.

~. , , ) , ~ , ) , , ~ , 1 - - - " ' . ' ' ~

:::-::: : "''1'

. . . .

~ .. ,

.

.

.

.

, I

.

-

.

.

~a,

. . . . .

60S 908

.

22.22Z2222

'9. |

1~

30S

.

: : 7 7 7 -- -- : -- .-~ . . >.,-.,(. . . . . . . . . . . . .

9 ' , ' " ~ ' : ' ' \ / ' -Vx :t--.-.~:~ ~-~- - . . . . . . . . 9 "~.t~ff.-'~.,,/x:".~.~-~'-': 7.. _

'~

EQ

T

. .~. .

. \ 1. . . . 4. ,

/ -!

~-u.222_22~_.-; ? ~ = . ,. \..I ~ . . . . " '/; ~.. . . . / ~ . . . . . . . "

r

. ..,.--

.._-_.~-~ .

.

..

9

."

. . . . . . .

~

,~

.

. )

.

.

"[/

.

"

-

-

.,

-

...-, .~:.-...,.-.2-1.;-.-~-:i~-.'.', ? ;,i,VI/~, ,.....

L

.....-. .- .-.-.:.a:.; " : : : ~---: : : :..-_ ~ . . . . .. . . . . 60E 90E 120E150E 180E

Fig. 9. The mean January-February-March Rossby wave source term deduced from the NMC, Washington 200 hPa analyses from 1987-1989. The contour interval is 6 x 10-11 s-2 and negative contours are dashed (taken from RASMUSSON and Mo, 1993). scale climate. Whilst a full debate on this topic is beyond the scope of this chapter, we illustrate some of the problems by way of a discussion of the synoptic climatology of storm tracks and a brief s u m m a r y of the interactions that contribute to maintenance of the s u m m e r m o n s o o n circulation in the western North Pacific. AMIP R31 KU0 -180

-

ENSEMBLE MEAN DJF ( 8 0 - 8 7 ) ROSSBY S 2 0 0 h P a

-1120

-90

?

610

1 ~0

11 =0

~ O"~~~

_ ,,.,~

"='..~

;'-'d

-~,~.~

9

.~:

9

-30

(~

r ~.

(

-

tl

~d

~

( 0 S (_ ~~I . l~.

60

>~

~

,

,:,

-

'

'

"

--

l

-60

0

O

o

(

30

--

-30

--o _-',-:-'" ,'q~ :_.- 12o-

. --

-

0

--

~

"

~

~ C

-60

o I=

-90

-180

-120

-60

0

60

120

180

-90

Contour f r o m - 4 8 0 to 600 by 60 (x 10 -12)

Fig. 10. The mean of the Rossby wave source at 200 hPa for December-January-February for 19801987 from the Australian Bureau of Meteorology Research Centre AMIP experiment. The contour interval is 60 x 10-12 s-2 and negative contours are dashed. Negative areas are also shaded.

301

Dynamics offuture climates Storm tracks

The synoptic climatology of storms, cyclone and anticyclone tracks, fronts and "storm tracks" (i.e. regions of maximum variance in geopotential at synoptic time scales less than a week) are important components of climate, especially because of their social and economic impacts on affected populations. The interesting "weather" experienced in mid-latitudes is primarily associated with "storminess" due to the passage of cyclones, anticyclones and their accompanying fronts. The usual interpretation of such systems is that they are a manifestation of the conversion of the potential energy associated with the north-south temperature gradient into the energy of atmospheric eddies through the process of baroclinic instability. The more traditional synoptic approach of "storm track" analysis usually only considers the low centres when individual "storms" are tracked (e.g. MURRAY and SIMMONDS, 1991; KONIG et al., 1993). As TRENBERTH (1991a) argues, it is probably more meaningful to include both highs and lows since "weather" events are associated not only with the approach of low pressure systems but also with the withdrawal of high pressure systems. WALLACEet al. (1988) present a good summary of the relationships between traditional indicators of storm activity (tracks of cyclones and anticyclones) and the distribution of time-filtered geopotential height. Figure 11, from TRENBERTH (1991a), illustrates the link between the storm track and the mean flow through the polar jet streams; Figure 11 (a) shows the zonal mean of the standard deviation of the geopotential height at 300 hPa as a direct measure of the storm track activity over the annual cycle, Fig. 1 l(b) shows the zonal average of the zonal wind at 300 hPa to illustrate the annual cycle of the subtropical and polar jet streams and Fig. 11 (c) shows the meridional gradient of temperature at 700 hPa as a direct measure of baroclinicity in the lower troposphere. One of the most striking things about storm track activity is the persistent storminess in the Southern Hemisphere throughout the year compared to the Northern Hemisphere where the storminess shows a considerable seasonal variation in both intensity and location. The local meridional temperature gradient is a reasonable indicator of the location and strength of the storm track. There is also a strong relationship between the storm track and the major tropospheric jet streams. In the Southern Hemisphere, there is a noticeable semiannual variation with a slight shift of the polar jet and the storm tracks towards the pole during the transition seasons and is most probably related to the existence of the strong heat sink of Antarctica. TRENBERTH (1991a) has demonstrated that the observed relationships amongst bandpass filtered eddy quantities can be understood in terms of geostrophic theory and perturbation analysis applied to baroclinic systems. A detailed study of the interannual variability of storm tracks and relations with the mean flow must await the emergence of global data from data assimilation "re-analysis projects". However, since large interannual variations in the mean flow are well documented and these variations have a profound influence on high frequency storms, it is to be expected that there will be significant interannual variations in storm tracks. TRENBERTH (1991a) has also investigated the question of the impact of the storm track eddies on the mean flow itself using a localised Eliassen-Palm flux. He finds that baroclinic effects from the poleward heat flux dominate in winter and the eddies act so as to decelerate the mean westerlies in the upper troposphere. The barotropic effect due to the flux convergence of meridional momentum contributes to the maintenance of the mean

302

Regional-scale interactions

90N

~

6ON

~ 2 0

"~--

---~~

>

9

3ON o 13 .

m

_J

9

JAN FEB MAR APR MAY JUN

.-..

4

JUL AUG SEP OCT NOV DEC JAN Month

U Zonal Average 1979-1989 =t 500 mo

~20~~~30~[ "---..

13 4

"

/

6OS-

'. . . . .

- - - - q

--s

('~L

I

I

-

905 JAN

FEB

MAR

APR MAY

JUN

JUL AUG

SEP

OCT NOV

DEC

JAN

Month

dT /dy Zonal Average 1979-1989 at 700 mb 90N

9

60N- ......... 4 - - ; : ~:-~::-:-.-:T-"-:: 50N-

.

',- ........... : ............. _ - - 4 ..... : _ ~ . . . . . . . . . . . . '-........ :~"

. ....

..~::-::.-:-:---:4:(-:.:!:~:;::: .. . . . . . . . .

(1) 13

.

~-------4~'~o

0 _ _ 0 . _ ~ _ _ 3OS,-' ~

' " ~ "

60'S~ _ . . . ~ 9

O

~

JAN

j

~

_

FEB

MAR

APR

MAY

JUN

JUL

AUG

SEP

OCT

NOV DEC

JAN

Month

Fig. 11. The seasonal variation in monthly mean zonal means for the years 1979-1989 from ECMWF analyses. (a) Time filtered zonal average of the standard deviation in 300 hPa geopotential, contour interval 10 m, values greater than 80 m are stippled; (b) zonal mean of the zonal component of the wind at 300 hPa, contour interval 5 m s-1, values greater than 25 m s-1 are stippled and negative values (easterlies) are dashed; (c) meridional temperature gradient at 700 hPa, contour interval 2 x 106 K m -1, magnitudes grater than 6 x 10 -6 K m -1 are stippled and negative values are dashed ( f r o m TRENBERTH, 1991).

303

Dynamics offuture climates westerlies by accelerating the main polar jet in the upper troposphere. The role of eddy-induced diabatic heating in offsetting the effects of eddies on the mean flow is not resolved. HOSKINS and VALDES (1990) and HALL et al. (1994) have related the storm tracks to the baroclinicity of the mean flow and have also gone on to show how the mean equator-to-pole temperature gradient which is producing the baroclinicity can be maintained by local eddyinduced diabatic heating despite its continual erosion by the baroclinic instability process itself. S u m m e r monsoon circulation in the western North Pacific

A regular feature of the western North Pacific is the presence of a summer-monsoon circulation with a monsoon trough and a broad belt of equatorial westerly flow extending eastwards several thousand kilometres from the Asian subcontinent (Fig. 12). This circulation lies between broad low level easterlies in the central Pacific and westerlies associated with the Asian monsoon and is noted for its stable nature once established (TANAKA, 1992; HOLLAND, 1994). It is an active region of extratropical interaction, lying at the western end of the Pacific subtropical-storm track identified in the previous section, and in a region of regular subtropical transition by intense tropical cyclones. Whilst the convection has a transitory nature, it is located for sufficient time to have an impact by the non-linear rotational-flow response discussed by SARDESHMUKH and HOSKINS (1988) and is the origin of the Pacific-North American (PNA) teleconnection (HOREL and WALLACE, 1981; HOSKINS

)

) SUBTROPICAL

ASIAN MONSOON

~/ i \ DEPRESSION /

_

PACIFIC EASTERUES

Fig. 12. Schematic of the major lower-tropospheric components of the western North Pacific summer monsoon. Diabatic heating by moist convection is maintained in the confluence zone east of the monsoon depression and, in turn maintains the overall flow pattern. Complex scale interactions also occur and lead to development of tropical cyclones.

304

Regional-scale interactions and KAROLY, 1981). The region also is west of the major, upper-tropospheric accumulation zone of wave energy identified by WEBSTER and CHANG (1988) and directly in a potential lower-tropospheric accumulation region. HOLLAND (1994) describes the establishment of the monsoon as a result of enhanced convection over the warm western Pacific Ocean. The convective region provides a Rossbywave response at the scale of the Rossby deformation radius (section on Diabatic heating). The non-linear phase propagation in the lower troposphere is poleward and westward (e.g. FIORINO and ELSBERRY, 1989) so that a well-defined monsoon gyre is established to the northwest of the region of enhanced cloudiness. Rossby waves at this scale also have strong eastward group speed (GILL, 1982). The associated energy propagation develops the anticyclonic gyre northeast of the monsoon gyre and the equatorial eddy to the southeast. The response in the upper troposphere is similar, although of opposite sign and considerably more diffuse due to the processes discussed in the previous section. The confluence of the low-level flow east of the monsoon gyre enhances the potential for sustained moist convection in this region, thus leading to a stable feedback cycle and maintenance of the overall circulation (Fig. 2). This convective enhancement may arise from extratropical forcing, or eastward movement of a Madden Julian Oscillation (MJO1 MADDEN and JULIAN, 1972), both of which are within the capacity of climate models. However, it is highly likely that the convective organisation also is strongly influenced by interactions between gravity modes (YAMASAKI, 1984, 1988, 1989; MAPES, 1993), by coalescence of mesoscale complexes (SIMPSONet al., 1993) and by the development of long-lived mesoscale clusters identified for the vertical windshear conditions in this region by YAMASAKI (1984). Westward travelling Rossby modes, often referred to as easterly waves (REED and RECKER, 1971), can be expected to slow down and accumulate in the confluence zone (FARRELL and

WATTERSON, 1985; WEBSTER and CHANG, 1988; CHANG and WEBSTER, 1990). The result will be short-term enhancements of convection, which may produce one of the tropical cyclones that frequently develop in this region (BRIEGEL, 1993). HOLLAND (1994) further shows that phase locking between the energy dispersion from one tropical cyclone with subsequent easterly waves moving into the confluence region may substantially enhance the potential for secondary cyclone development several days later. Vortex-vortex interaction, which is known to be highly chaotic under many conditions occurs regularly in this region, indeed, interacting tropical cyclones were first identified here (FuJIWHARA, 1921). Recent studies have shown that this interaction is complex and highly non-linear (HOLLANDand DIETACHMAYER, 1993; LANDER and HOLLAND, 1993; RITCHIE and HOLLAND, 1993; WANG and HOLLAND, 1995). Mesoscale vortices can be expected to occur frequently in the vicinity of long-lived mesoscale convective clusters (RAYMONDand JIANG, 1990) and RITCHIE et al. (1993) have suggested that merger and growth of these vortices may be associated with formation of tropical cyclones in this region. HOLLAND and LANDER (1993) and WANG and HOLLAND (1995) also indicate that the movement and development of tropical cyclones may be substantially affected by the presence of other mesoscale vortices in their vicinity. Of considerable importance to climate conditions is the manner in which a small vortex may rapidly merge with the large monsoon gyre to produce a breakdown of the overall monsoon circulation (HOLLAND, 1994). The presence of tropical upper tropospheric troughs (TUTTS) (SADLER, 1978) provides an upper-tropospheric influ-

305

Dynamics offuture climates ence on the monsoon region that is considered to be of importance by forecasters in the region, but is poorly defined. Identifying future regional climates requires a consideration of scale interactions such as those identified in previous paragraphs. Of particular importance is the identification of stable processes and those that are likely to produce unstable conditions, which are inherently difficult to forecast adequately. Unfortunately, our detailed knowledge of the full range of interactions is still far from complete. Furthermore, current climate models are capable of capturing only some of the known mechanisms, other processes, such as mesoscale vortex interaction, being neglected. Special problems arise from the impact of convective and mesoscale processes on the larger scales. For these reasons, objective specification of future regional climates is not at present possible except in the most general and uncertain of terms.

Future climate dynamics We conclude this overview of climate dynamics by examining the potential changes that might occur in climate over the next millenium. Barring major catastrophes, such as an impact by a large extraterrestrial object (Chapter 4 by RAMPINO), we expect that the basic dynamics of future climates will be essentially the same as those described here for the current climate. For example, substantial changes should not be expected in the rotation rate of the Earth, the basic moist thermodynamics, the energy balance of solar warming at the equator and cooling at the poles, and the related baroclinic processes (but cf. Chapter 14 by PENG). As a result, the tropics will remain a region of significant moist convection, but with weak mass-wind balance and high dispersion of Rossby modes. Hadley and Walker circulations and equatorially trapped systems, such as the Madden-Julian Oscillation, will remain. Higher latitudes will continue to be dominated by baroclinic processes, with strong mass wind balance. The separation of scales, and resulting complex interaction between ocean and atmospheric processes, will continue as will the basic dynamics of interaction with both the ocean and land surfaces. Changes can be expected in the climate system response to the underlying dynamics, some of which will be subtle, others substantial. For example, significant changes in the composition of atmospheric gases and particulates will occur from both anthropogenic (e.g. CO2 emissions) and natural (e.g. changes to frequency of volcanic eruptions) causes (Chapters 9 by WANG et al. and Chapter 10 by ANDREAE); the distribution of the major land masses and orographical features and the magnitude of solar heating will evolve over millennia (cf. Chapter 3 by BARRON). The climate system response may be stable or it may involve positive feedback and evolution to a new, or substantially different climate. The types of climate that may result are difficult to predict, but we know that the climate has evolved through major cycles in the past and can be expected to continue to do so in the distant future (cf. Chapter 15 by KUMP and LOVELOCK). We draw on three examples of the types of detailed response that can occur to illustrate the complexity of future climate responses to subtle changes in the dynamical and thermodynamical balance and forcing. The linkage of the meridional temperature gradient to the extra-tropical atmospheric eddies that dominate the transport of heat and momentum is a fundamental part of tropospheric dy-

306

Future climate dynamics namics. From studies of baroclinic instability, any increase (reduction) in this temperature gradient should cause these eddies to grow (shrink) and the associated transport to increase (decrease). Many different Atmospheric General Circulation Models have been used to determine the equilibrium response of climate to a doubling of the CO2 concentration in the atmosphere. These models generally show a warming which is larger at high latitudes than at the equator (this effect is due primarily to the positive feedback from the overall retreat of snow and ice). Thus the pole to equator temperature gradient is weakened in both hemispheres and a change in the behaviour of baroclinic eddies is to be expected. However, the situation is made more complicated because the models all show an enhanced warming in the tropical upper troposphere (this is a consequence of the fact that warmer and moister air parcels lifted upwards from the surface have more latent heat to release in the upper troposphere) thus increasing the equator to pole temperature gradient in the upper troposphere. This behaviour is illustrated in Fig. 13 where the equilibrium response in annual mean temperature simulated by the Australian Bureau of Meteorology Research Centre (BMRC) Atmospheric General Circulation Model (AGCM) is shown (MCAVANEY et al., 1994). It is not clear whether the baroclinic eddies are more sensitive to the decreased lower tropospheric temperature gradient or the increased upper tropospheric temperature gradient. Further complications ensue if we consider the results from transient CO2 modelling experiments. In this case, most coupled ocean-atmosphere models show a very much reduced surface warming in the Southern Hemisphere due to oceanic mechanisms that promote cooling of the ocean in high southern latitudes (HOUGHTON et al., 1992; HELD, 1993), while the tropical upper

8.900~ 7 4 . lOS

188.6~

3 3 6 . 0 .c

500. O.c

663.9,{

811.4. c

925.9 S 991.1,c

Contour from-7

to 3.5 by .25

Fig. 13. The ensemble average of the zonal average annual change (2x CO2 - 1 x CO2) in air temperature (K) for an equilibrium 2 x (202 experiment conducted by the Australian Bureau of Meteorology Research Centre. The ensemble average was for 15 years. The contour interval is 0.25 K and negative contours are dashed. Areas of warming are shaded.

307

Dynamics offuture climates B. 900S 74. lOS

nnMIIIIHMUlnlIMIHmnmnIIIMnM' [ImlMllnlllllllll[llllllllllllllllll[ll

18B. 6S

- 1 .... .... "" _ -----"

~

--.". . . . . . .

~

""

'~7-. -:::::: . . . . . . . ~ :

~

.........

~

.... +..+~ ................+...+..... ~ ~

"-_-. . . . . . . . . . . . . . . . . . . . . :.-:_::z:.:::.:.=.::: :--'_-": ~C~: ='-- 1 .~" ...................................t

~

l

+ +

~+++.+~.+~

......

~----.-__-_-_-+

o

--.:.._..:

"+: ::.. ""i'+'?!'!'~'i'i'!'~'!'! ............. ' .

+

~J"

~

~

~:..+

~ : ' : + : + : ' : " +:' ++:':':':':':'.

i.+~

"':': . 5 ......

+ ~.+;-+~i~'::::::%.+:.+:%+i+++.~......:...... ~, ,|~ ~i+:..."::::::~::::+:#:

500. OS

663. gs

BI I. 4 S

~--

991.;s, ~

~--~i

.

.

.

.

.

~i+~+i.~

.

I

+

+ re+N+ ,

70

'

+o

'

3'o

'

~b

'

Contour from-6.5

-Io

'

-+o

'

-!io

'

-+o

'

-4o

to 3 b y .25

Fig. 14. The ensemble average of the zonal average annual change (transient - control2) in air temperature (K) for an transient increase in CO 2 of 1% per year in an experiment conducted by the Australian Bureau of Meteorology Research Centre using a coupled ocean-atmosphere model without flux correction. The ensemble average was for 20 years centred around the time of doubling of CO 2. The contour interval is 0.25 K and negative contours are dashed. Areas of warming are shaded.

troposphere is still warmed. This slightly modified behaviour is illustrated in Fig. 14 which shows the annual mean temperature response at the time of doubling in the transient CO2 increase experiment conducted by BMRC (POWER et al., 1993). This transient experiment, and other similar experiments conducted by other modelling groups, indicates that the response in the baroclinic eddies may be different between the two hemispheres in enhanced greenhouse conditions. HELD (1993) suggests, on the basis of linear instability theory, that the temperature gradient in the lower troposphere should dominate the response under greenhouse warming and hence eddies in the Northern Hemisphere should transport less energy while those in the Southern Hemisphere should transport more. However, as HELD (1993) points out, no definitive statement can be made since, even if the eddy transport of heat responds more to changes in the temperature gradient in the lower troposphere, it may be that the eddy transport of momentum is also sensitive to the gradient in the upper troposphere resulting in eddies with a larger fraction of their total energy in the middle and upper troposphere. Further complications arise due to consideration of the role of moist processes in conjunction with the growth and decay of the baroclinic eddies. Any warming of the atmosphere will tend to increase the moisture holding capacity in the eddies. The intensity of the resulting storms will depend on the balance between the competing effects of direct enhancement of the energy of eddies by latent heat release versus the determination of the size and strength of the eddies through the balance between the poleward transport of energy by the eddies and the gradient in the heating (HELD, 1993). Much further work is required before a more complete

308

Future climate dynamics understanding of what controls the size and strength of the eddies in mid-latitudes. A further issue that requires further study is the relationship between the cloud field that is associated with regions of high eddy activity and the temperature gradient which is responsible for their formation. Consideration of potential changes in the distribution of the rainfall associated with a movement of mid-latitude depressions is also of great importance since in most regions much of the rain falls from the equatorward edge of these depressions. The way in which the latitudinal distribution of eddy activity is affected by changes in the horizontal temperature gradient is not clear. Preferred regions for the growth and decay of depressions exist through zonal asymmetries of the large scale temperature and flow fields. These asymmetries arise in turn from major asymmetries in the lower boundary either through topographic features or contrasts between the heating over land compared to an adjacent ocean. Linear stability analyses of realistic planetary scale flows (e.g. FREDERIKSEN, 1985) and sensitivity studies with AGCMs are gradually leading to a better understanding of the factors which control the location of the tracks of depressions. However, because of the fundamental non-linearities involved, it is very difficult to predict how the dominant planetary-wave pattern and associated storm tracks might change in response to a given change in forcing. The potential changes in the circulation of the Atlantic Ocean due to changes in freshwater input at high latitudes are discussed in Chapter 14 by PENG. The stability of North Atlantic deep water formation and the possibility of different modes of oceanic circulation have been the subject of much recent research using ocean only models. In most of these studies, feedback mechanisms involving salinity operate. The robustness of distinct climatic ocean states remains to be fully explored with three-dimensional coupled ocean-atmosphere models. However, in the context of global warming, it is possible that, with higher temperatures, a modest freshening of high latitude waters is possible due to increased moisture convergence producing enhanced excess of precipitation over evaporation (a requirement of moisture balance). Any freshening of polar water would weaken vertical overturning in the ocean and thereby increase the time that the surface waters were exposed to the freshening, reducing the overturning still further. Such a weakening in the overturning is found in many coupled ocean-atmosphere experiments. Whether an irreversible transition to a state where there is no deep-water formation will occur remains an open question. Extensive public discussion on climatic changes to severe weather systems, such as hail storms, mid-latitude wind storms and tropical cyclones have engendered much debate because of the potentially catastrophic consequences to the global insurance industry and economies of individual countries. Much of the discussion is conjectural as it is difficult to consider regional-scale changes let alone changes in very rare severe weather events, for which there is often not even suitable records available for current climate (HOLLAND, 1981; NICHOLLS, 1992). We therefore examine the potential for changes associated with tropical cyclones, as an example of the problems and solutions to be applied. Some studies have shown skill in climate model predictions of systems that have some of the large-scale characteristics of tropical cyclones (BENGTSSENet al., 1994). However, we have shown that tropical cyclone development is fundamentally linked to complex scaleinteraction processes at the regional scale. Since these scale interactions cannot be confidently predicted in future climate scenarios, and since climate models cannot resolve the cyclone core and actual intensity, we need to utilise some other measure for estimating cli-

309

Dynamics of future climates

mate changes. Whilst there are some creative attempts at applying climatic indices to model fields (e.g. RYAN et al., 1992), most methods have relied on application of thermodynamic approximations, such as by EMANUEL (1988) to tropical cyclones or to extrapolating current SST relationships (GRAY, 1968) to changes in ocean temperatures. These methods were investigated in a review paper by LIGHTHILL et al. (1994), who concluded that there were no consistent statistical relationships between cyclone numbers and SST changes. For example the study of E1 Nifio effects by HOLLAND et al. (1988) showed that the direct effects of SST changes were negligible and that it was the subtle circulation changes, and scale interactions that were responsible for any observed relationship. A careful study by LANDSEA and GRAY (1994) also found no relationship between the proportion of intense cyclones and surface temperature anomalies over the entire northern hemisphere. The thermodynamic model predictions of maximum potential intensity (EMANUEL, 1987) indicate quite strong sensitivity to small ocean changes at the very intense cyclone stage. The observational evidence for such changes is equivocal, with some conclusions of no change (EVANS, 1993) and other evidence of some change (DEMARIA and KAPLAN, 1994), albeit with less amplitude than predicted from purely thermodynamic considerations. Other theoretical studies of the impact of spray and ocean energy transfers (FAIRALL et al., 1994) indicates that there may be counteracting processes occurring for extreme cyclones. The best conclusion for future climates is that the regional characteristics can be expected to change in manners that are difficult to predict with current knowledge and techniques. We support the conclusions by LIGHTHILL et al. (1994), that SST changes of a few degrees will have negligible direct effect on tropical cyclones. Any small changes in cyclones are likely to be difficult to ascertain above the high natural variablity that occurs. However, substantial indirect changes in tropical cyclones at a regional level can be expected, and have been observed over recent decades. These changes arise from readjustments in the regional response to the global climate system.

References

BENGTSSON, L. and SHULKA, J., 1988. Integration of space and in situ observations to study climate change. Bull. Am. Meteorol. Soc., 69:1130-1143. BENGTSSON, L., BOTZET, M. and ESCH, M., 1994. Hurricane-type vortices in a general circulation model, Part I. Rep. No. 123, Max-Planck-Institut fiJr Meteorologie, Hamburg, Germany, ISSN 0937-1060, 42 pp. BOER, G. J., 1982. Diagnostic equations in isobaric coordinates. Mon. Weather Rev., 110:1801-1820. BOER, G. J., 1986. A comparison of mass and energy budgets from two FGGE datasets and a GCM. Mon. Weather Rev., 114: 885-902. BOER, G. J., 1990. Earth-atmosphere exchange of angular momentum simulated in a general circulation model and implication for the length of day. J. Geophys. Res., 95:5511-5531. BOER, G. J. and SARGENT,N. E., 1985. Vertically integrated budgets of mass end energy for the globe. J. Atmos. Sci., 42:1592-1613. BRIEGEL, L. M., 1993. An observational study of tropical cyclogenesis in the western Pacific Ocean. Proc. 20th Conf. on Hurricanes and Tropical Meteorology, San Antonio, May. American Meteorological Society, Boston, MA, pp. 397-400. CHANG, H.-R. and WEBSTER,P. J., 1990. Energy accumulation and emanation at low latitudes. Part II: nonlinear response to strong episodic equatorial forcing. J. Atmos. Sci., 47: 2624-2644. CHARNEY, J. G., 1947. The dynamics of long waves in a baroclinic westerly current. J. Meteorol., 4: 135-162.

310

References

CHEN, T.-C., 1985. Global water vapour flux and maintenance during FGGE. Mon. Weather Rev., 113: 1801-1819. DEMARIA, M. and KAPLAN, J., 1994. Sea surface temperature and the maximum intensity of Atlantic tropical cyclones. J. Climate, 7: 1324-1334. EMANUEL, K. A., 1987. The dependence of hurricane intensity on climate. Nature, 326: 483-485. EMANUEL, K. A., 1988. The maximum intensity of hurricanes. J. Atmos. Sci., 45:1143-1155. EVANS, J. L., 1993. Sensitivity of tropical cyclone intensity to sea surface temperature. J. Climate, 6: 1133-1140. FAIRALL, C. W., KEPERT, J. D. and HOLLAND, G. J., 1994. The effect of sea spray on surface energy transports over the ocean. The Atmosphere Ocean System, in press. FARRELL, B. and WArrERSON, I., 1985. Rossby waves in opposing currents. J. Atmos. Sci., 42: 428449. FIORINO, M. J. and ELSBERRY, R. L., 1989. The role of vortex structure in tropical cyclone motion. J. Atmos. Sci., 46: 975-990. FORTELIUS, C. and HOLOPAINEN, E., 1990. Comparison of energy source estimates derived from atmospheric circulation data with satellite measurements of net radiation. J. Climate, 3: 646-660. FRASER, J. R., 1994. Vorticity and momentum budget analysis of ENSO anomalies in a GCM. M. Sc. Thesis, Monash University, Melbourne. FREDERIKSEN, J. S., 1985. The geographic locations of Southern Hemisphere storm tracks: linear theory. J. Atmos. Sci., 42:710-723. FUJIWHARA, S., 1921. The natural tendency towards symmetry of motion and its application as a principle of motion. Q. J. R. Meteorol. Soc., 47: 287-293. GILL, A. E., 1982. Atmosphere-Ocean Dynamics. International Geophysics Series, Vol. 30, Academic Press, New York, ISBN 0-12-283520-4, 662 pp. GRAY, W. M., 1968. Global view of the origin of tropical disturbances and storms. Mon. Weather Rev., 96: 669-700. HALL, N. M. J., HOSKINS, B. J., VALDES, P. J. and SENIOR, C. A., 1994. Storm tracks in a high resolution GCM with doubled CO 2. Q. J. R. Meteorol. Soc., 120: 1209-1230. HALTINER, G. J. and WILLIAMS, R. T., 1979. Numerical Prediction and Dynamic Meteorology, 2nd Edition. Wiley, New York, ISBN 0-471-05971-4, 477 pp.. HELD, I. M., 1993. Large-scale dynamics and global warming. Bull. Am. Meteorol. Soc., 74: 228241. HIDE, R., DICKEY, J. O., MARCUS, S. L., ROSEN, R. D. and SALSTEIN, D. A., 1994. Diagnostic SubProject (No 15) for atmospheric angular momentum of the Atmospheric Model Intercomparison Project (AMIP). In: G. J. BOER (Editor), Research Activities in Atmospheric and Oceanic Modelling. WMO/TD 592, Geneva. HOLLAND, G. J., 1981. On quality of the Australian tropical cyclone data base. Aust. Meteorol. Mag., 29: 169-181. HOLLAND, G. J., 1994. Scale interaction in the western Pacific monsoon. Meteorol. Atmos. Phys., in press. HOLLAND, G. J. and DIETACHMAYER,G. S., 1993. On the interaction of tropical-cyclone scale vortices. II: Interacting vortex patches. Q. J. R. Meteorol. Soc., 119: 1363-1379. HOLLAND, G. J. and LANDER, M., 1993. The meandering nature of tropical cyclone tracks. J. Atmos. Sci., 50:1254-1266. HOLLAND, G. J., MCBRIDE, J. L. and NICHOLLS, N. N., 1988. Australian tropical cyclones and the Greenhouse effect. In: Greenhouse. Planning for Climate Change. E. J. Brill, Leiden, The Netherlands, pp. 438-455. HOLTON, J. R., 1992. An Introduction to Dynamic Meteorology, 3rd Edition. Academic Press, New York. HOREL, J. D. and WALLACE,J. M., 1981. Planetary-scale atmospheric phenomena associated with the southern oscillation. Mon. Weather Rev., 109:813-829. HOSKINS, B. J., 1983. Dynamical processes in the atmosphere and the use of models. Q. J. R. Meteorol. Soc., 109: 1-21. HOSKINS, B. J. and KAROLY, D., 1981. The steady linear response of spherical atmosphere to thermal and orographic forcing. J. Atmos. Sci., 38:1179-1196. HOSKINS, B. J., MCINTYRE, M. E. and ROBERTSON, A. W., 1985. On the use and significance of isentropic potential vorticity maps. Q. J. R. Meteorol. Soc., 111: 877-946.

311

Dynamics o f future climates HOSKINS, B. J. and VALDES, P. J., 1990. On the existence of storm tracks. J. Atmos. Sci., 47: 18541864. HOSKINS, B. J., Hsu, H. H., JAMES, I. N., MASUTANI, P. D, SARDESHMUKH,P. D. and WHITE, G. H., 1989. Diagnostics of the global atmospheric circulation based on ECMWF analyses, 1979-1989. UGAMP Technical Report No. 7, Department of Meteorology, University of Reading, Reading, UK, 217 pp. HOUGHTON, J. T., CALLANDER,B. A. and VARNEY, S. K. (Editors), 1992. Climate Change 1992. The Supplementary Report to The IPCC Scientific Assessment. World Meteorological Organisation and United Nations Environment Programme, Cambridge University Press, 200 pp. JOHNSON, D. R., 1980. A generalised transport equation for use with meteorological coordinate systems. MuD. Weather Rev., 108: 733-745. JOHNSON, D. R., 1989. The forcing and maintenance of global monsoonal circulations: an isentropic analysis. Adv. Geophys., 31: 43-329. JOHNSON, D. R. and DOWNEY,W. K., 1975. Azimuthally averaged transport and budget equations for storms: quasi-Lagrangian diagnostics 1. MuD. Weather Rev., 103: 967-979. KILADIS, G. N. and FELDSTEIN, S. B., 1994. Rossby wave propagation into the tropics in two GFDL general circulation models. Clim. Dyamics, 9: 245-252. KILADIS, G. N. and WEICKMANN, K. M., 1992. Extratropical forcing of tropical Pacific convection during northern winter. MuD. Weather Rev., 120:1924-1938. KONIG, W., SAUSEN,R. and SIELMANN,F., 1993. Objective identification of cyclones in GCM simulations. J. Climate, 6:2217-2231. LANDER, M. and HOLLAND,G. J., 1993. On the interaction of tropical-cyclone scale vortices. I: Observations. Q. J. R. Meteorol. Soc., 119:1347-1361. LANDSEA, C. W. and GRAY, W. M., 1994. Predicting Atlantic basin seasonal tropical cyclone activity. Weather Forecasting, 9:103-115. LIGHTHILL, J., HOLLAND, G. J., GRAY, W. M., LANDSEA, C., CRAIG, G., EVANS, J. L., KURIHARA, Y. and GUARD, C. P., 1994. Global climate change and tropical cyclones. Bull. Am. Meteorol. Soc., submitted. MADDEN, R. A. and JULIAN, P. R., 1972. Description of global circulation cells in the tropics with a 40-5-day period. J. Atmos. Sci., 29:1109-1123. MAPES, B. E., 1993. Gregarious tropical convection. J. Atmos. Sci., 50: 2026-2037. MCAVANEY, B. J. and COLMAN, R. A., 1993. The AMIP Experiment: the BMRC AGCM configuration. BMRC Research Report No 38. October. MCAVANEY, B. J., DAHNI, R. R., COLMAN, R. A., FRASER, J. R. and POWER, S. B., 1994. The dependence of the climate sensitivity on convective parameterisation, statistical evaluation. Global Planet. Change, in press. Mo, K. and RASMUSSON,E. M., 1993. The 200-mb climatological vorticity budget during 1986-1989 as revealed by NMC analyses. J. Climate, 6: 577-594. MURRAY, R. J. and SIMMONDS,I., 1991. A numerical scheme for tracking cyclone centres from digital data. Part 1: development and operation of the scheme. Aust. Meteorol. Nag., 39:155-166. NICHOLLS, N., 1992. Recent performance of a method for forecasting Australian region tropical cyclone activity. Aust. Meteorol. Mag., 40:105-110. OORT, A. H., 1983. Global atmospheric circulation statistics, 1958-1973. NOAA Professional Paper No. 14. US Government Printing Office, Washington, DC, 180 pp. + 47 microfiche. OORT, A. H. and PEIXOTO, J. P., 1983. Global angular momentum and energy balance requirements from observations. Adv. Geophys., 25: 355-490. PEIXOTO, J. P. and OORT, A. H., 1983. The atmospheric branch of the hydrological cycle and climate. In: A. STREET-PERROTTet al. (Editors), Variation in the Global Water Budget. Reidel, Dordrecht, pp. 5-65. PEIXOTO, J.P and OORT, A. H., 1984. Physics of Climate. Rev. Mud. Phys., 55: 365-429. PEIXOTO, J.P and OORT, A. H., 1992. Physics of Climate. American Institute of Physics, New York. PHILANDER, S. G. H., 1983. E1 Nifio Southern Oscillation phenomena. Nature, 302: 295-301. PHILANDER, S. G. H., 1990. El NiTro, La Nina and the Southern Oscillation. Academic Press, San Diego, CA, 293 pp. POWER, S. B., COLMAN, R. A., MCAVANEY, B. J., DAHNI, R. R., MOORE, A. M. and SMITH, N. R., 1993. The BMRC coupled atmosphere-ocean-sea ice model. BMRC Research Report No 37.

312

References RAYMOND, D. J. and JIANG, H., 1990. A theory for long-lived mesoscale convective systems. J. Atmos. Sci., 47: 3067-3077. REED, R. J. and RECKER, E. E., 1971. Structure and properties of synoptic scale wave disturbances in the equatorial tropical Pacific. J. Atmos. Sci., 28:1117-1133. RrrCnlE, E. A. and HOLLAND, G. J., 1993. On the interaction of tropical-cyclone scale vortices. II. Interacting vortex patches. Q. J. R. Meteorol. Soc., 119:1363-1379. RITCHIE, E. A., HOLLAND, G. J. and LANDER, M., 1993. Contributions by mesoscale convective systems to movement and formation of tropical cyclones. In: J. LIGHTHILL, K. EMANUEL, G. J. HOLLAND and Z. ZHENG (Editors), Tropical Cyclone Disasters. pp. 286-289. ROSEN, R. D., SALSTEIN, D. A. and WOOD, T. M., 1990. Discrepancies in the Earth-atmosphere angular momentum budget. J. Geophys. Res., 95: 265-279. ROSEN, R. D., SALSTEIN, D. A. and WOOD, T. M., 1991. Zonal contributions to global momentum variations on intraseasonal through interannual time scales. J. Geophys. Res., 96:5145-5151. RYAN, B. F., WATFERSON, I. G. and EVANS, J. L., 1992. Tropical cyclone frequencies inferred from Gray's Yearly Genesis Parameter. Validation of GCM tropical climates. Geophys. Res. Lett., 19: 1831-1834. SADLER, J. C., 1978. Mid season typhoon-development and intensity changes and the tropical upper tropospheric trough. Mon. Weather Rev., 106:1137-1152. SARDESHMUKH, P. O. and HOSKINS, B. J., 1988. The generation of global rotational flow by steady idealised tropical divergence. J. Atmos. Sci., 45: 1228-1251. SCHUBERT, W. H., HACK, J. J., SILVADIAS, P. L. and FULTON, S. R., 1980. Geostrophic adjustment in an axisymmetric vortex. J. Atmos. Sci., 37: 1464-1484. SIEGMUND, P. C., 1990. The effect of doubling atmospheric CO2 on the storm tracks in the climate of a general circulation model. Rep. DM-90-01. KNMI, De Bilt, 12 pp. SILVA DIAS, P. L., SCHUBERT,W. H. and DEMARXA, M., 1983. Large scale response of the tropical atmosphere to transient convection. J. Atmos. Sci., 40: 2689-2707. SIMPSON, J. S., KEENAN, T. D., FERRmR, B., SIMPSON, R. H. and HOLLAND, G. J., 1993. Cumulus merging in the maritime continent region. Meteorol. Atmos. Phys., in press. STEPHENSON, D. and HELD, I. M., 1993. Response of GCM stationary waves and storm tracks to increasing CO2. J. Climate, 6:1859-1870. TANAKA, M., 1992. Intraseasonal oscillation and the onset and retreat dates of the summer monsoon over east, southeast Asia and the western Pacific region using GMS high cloud amount data. J. Meteorol. Soc. Jpn., 70:613-629. TOWNSEND, R. D. and JOHNSON, D. R., 1985. A diagnostic study of the isentropic zonally averaged mass circulation during the First GARP Global Experiment. J. Atmos. Sci., 42:1565-1579. TRENBERTH,K. E., 1991a. Storm tracks in the Southern Hemisphere. J. Atmos. Sci. 48:2159-2178. TRENBERTH, K. E., 1991b. Climate diagnostics from global analyses: conservation of mass in ECMWF analyses. J. Climate, 4: 707-722. TRENBERTH, K. E. (Editor), 1992. Climate System Modeling. Cambridge University Press, Cambridge, UK. TRENBERTH, K. E. and OLSON, J. G., 1988. ECMWF global analyses 1979-1986: circulation statistics and data evaluation. NCAR Technical Note NCAR/TN-3OO+STR, Boulder, CO, 94 pp. TRENBERTH, K. E. and SOLOMON, A., 1993. Implications of global atmospheric spatial spectra for processing and displaying data. J. Climate, 6:531-545. TRENBERTH,K. E. and SOLOMON,A., 1994. The Global heat balance: heat transports in the atmosphere and ocean. Clim. Dynamics, 10:107-134. TRENBERTH, K. E., BERRY, J. C. and BUJA, L. E., 1993. Vertical interpolation and truncation of modelcoordinate data. NCAR Technical Note NCAR/TN-396+STR, Boulder, CO. WALLACE,J. M. and GUTZLER,D. S., 1981. Teleconnections in the geopotential height field during the Northern Hemisphere winter. Mon. Weather Rev., 109:784-812. WALLACE, J. M., LIM, G.-H. and BLACKMON,M. L., 1988. Relationship between cyclone tracks, anticyclone tracks and baroclinic waveguides. J. Atmos. Sci. 45: 439-462. WANG, Y. and HOLLAND,G. J., 1995. On the interaction of tropical-cyclone scale vortices. IV: BarDclinic processes. Q. J. R. Meteorol. Soc., 121: 95-126. WEBSTER, P. J., 1972. Response of the tropical atmosphere to local, steady forcing. Mon. Weather Rev., 100: 518-541.

313

Dynamics of future climates WEBSTER, P. J. and CHANG, H.-R., 1988. Equatorial energy accumulation and emanation regions: impacts of a zonally varying basic state. J. Atmos. Sci., 45: 803-829. WEBSTER, P. J. and HOLTON, J. R., 1982. Cross-equatorial response to middle-latitude forcing in a zonally varying basic state. J. Atmos. Sci., 39: 722-733. YAMASAKI, M., 1984. Dynamics of convective clouds and 'CISK' in vertical shear flow - with application to easterly waves and squall-line systems. J. Meteorol. Soc. Jpn., 62: 833-862. YAMASAKI, M., 1988. Towards an understanding of the interaction between convection and the largerscale in the tropics. Aust. Meteorol. Mag., 36:171-182. YAMASAKI, M., 1989. Numerical experiment of the tropical cyclone formation. J. Meteorol. Soc. Jpn., 67: 529-539. ZHANG, C. and WEBSTER, P. J., 1991. Laterally forced equatorial perturbations in a linear model. Part I: stationary transient forcing. J. Atmos. Sci., 49: 585-607.

314

III. Human Factors Affecting Future Climate

This Page Intentionally Left Blank

Chapter 9

The greenhouse effect of trace gases WEI-CHYUNG WANG, MICHAEL P. DUDEK AND XIN-ZHONG LIANG

Introduction The impact of humans on the atmosphere is now widely recognized: the enhanced greenhouse effect resulting from emissions of CO2 and other gases, such as CH 4, CFCs and N20 and the observed 03 depletion in the lower stratosphere and increase in the upper troposphere in the middle to high latitude of the Northern Hemisphere are two important features. This chapter discusses the enhanced greenhouse effect associated with increasing greenhouse gases with focus on the potential climate change in the next few decades. The first section provides a description of the factors related to the greenhouse effect and the approaches to assess the climatic effects. The following section summarizes the studies of the greenhouse effect due to the projected increases of uniformly mixed gases CO2, CH4, CFCs and N20. Then model simulations of the climatic effect due to observed lower stratospheric 03 depletion during the last few decades are presented and the radiative forcing due to tropospheric 03 increase is described. A brief summary and discussion about future greenhouse effect is given. Extensive documentation about the greenhouse effect can be found in HOUGHTON et al. (1990, 1992), WMO (1991, 1994) and also in this book (e.g. Chapter 2 by BERGER; Chapter 10 by ANDREAE; Chapter 11 by BRASSEUR et al.).

The greenhouse effect The Sun provides the Earth's only external source of heat, the solar radiation in the visible and near-infrared spectra. The Earth also radiates energy back into space and, because of a much colder temperature than the Sun, the energy is in the form of infrared radiation. A balance between solar radiation and infrared radiation yields the mean temperature of the Earth, about-18~ The thermal structure of the atmosphere of the Earth is influenced by the presence of trace gases, which modulate the solar radiation and thermal emission. The principal gaseous absorbers of solar radiation are water vapour (H20) in the troposphere and ozone (03) in the stratosphere. H20 absorbs primarily in the near-infrared spectral region while 03 is most effective in the ultraviolet and visual regions. About 100 W m -2 of the 340 W m -2 incident solar radiation at the top of the atmosphere are reflected back to space mainly by clouds and surface while about 80 W m -2 are absorbed by the atmosphere and the rest absorbed by the surface. In the infrared, H20 effectively blocks thermal emission from the surface except for the "window" region between 7 and 12/~m where CO2, 03, CH4, CFCs and N20 with strong ab-

317

The greenhouse effect of trace gases

G(Wm -z)

vs

SST(K)

(April; Global Ocean) 300

(1987,total)

(1987,clear)

200

I i-~

""

/ ,

100 9

0

:....~

~

.~

9

o

9

,

270

i

,

280

i

290

,

1

,

' 2ao

zTo

300

,

I

zgo

i

300

300 (1988,total)

(1988,clear)

,i. ,.;C

200

..

f

,~...... ~ : ~ " ~9 9 ~.

"~

_ ~.,:- ~..~,,; ..:4~ cooling is dominated by C O 2. The stratospheric cooling shows stronger latitudinal variation during winter. However, the seasonal contrast of the latitudinal variation of tropospheric warming is small. The larger warming calculated in the tropical upper troposphere (>6~ during summer) is due mainly to convective heating (HOUGHTON et al., 1990), although the additional radiative warming of CH 4 and CFCs further augments the temperature increase. Larger surface warming at high latitudes is also calculated, mainly caused by the melting of ice and snow cover; this high latitude amplification is particularly large during winter in the Northern Hemisphere. Fig. 5 shows the model calculated changes of the latitude-altitude distribution of cloud cover for the two seasons. The cloud cover generally increases at upper levels and decreases at middle and lower levels, with the transition closely matching the tropopause height. Larger increases of high level clouds occur in the tropics and at high latitudes in winter. On the 34

34

32

32

3O

3O

28

28

26

28 -

24

24 22 2O

._.22 .Jr 2O

~

18

t

~16 ~I--._..~

18

r--i---.~

14 12

,o

{~..-:-:~,0.85. In a series of measurements at the Mauna Loa observatory, CLARKE and CHARLSON (1985) found a mean to s of 0.93 for the remote mid-tropospheric aerosol over the Pacific. They suggested soot particles at a concentration of some 10 ng m -3 to be the light-absorbing component of this aerosol, levels similar to those measured over the remote oceans by ANDREAE (1983) and ANDREAE et al. (1984). The single scattering albedo of smoke particles from forest fires in North America has been reported to be relatively low (0.83 _+0.11; RADKE et al., 1991),

him 0.005 0.01

0.001

51I.1"~--" ~//r . .

/

'~9'' / /

0.05

Wirming -'"~-~'~'~" ///.i

// //

0.05I:i ~~~ ~~""C~176 [~" Urban ~\ /I: ~"Forests Farmland Snow & Ice \\ / Ocean Deserts Clouds

o. 1 0.20 0.30 0.40 0.50 0.60 0.70 0.80

.--,

0.90

o

0.95

.~

-~ / 0.98 |-t

6

.~

r~

0.01/ i l i i i i i i i / 0.99 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.O Surface albedo, c~

Fig. 8. The relationship between absorption-to-backscattering ratio (a/b) and the critical surface albedo (a) (lower scale), and between a/b and the imaginary index of refraction (nim, upper scale) for a given radius of particle (dashed line). Calculations are for real part of index of refraction equal to 1.5, and for a solar zenith angle of 65 ~ (an average value taking the globe as a whole). Shaded area indicates locus of some critical values when infrared radiation was taken into account along with solar radiation (from KELLOGG,1980).

366

Climatic effects while smoke from biomass burning in Brazil had single scattering albedos around 0.90 (KAuFaVIAN et al., 1992). Remote sensing studies have suggested even higher albedos (near 0.97) for biomass smoke (KAU~AN et al., 1990a,b). Overall, most aged anthropogenic aerosols seem to have ~os values >0.90 (a/b < 0.11). Combinations of aerosol a/b and surface albedo a which plot to the left of the dividing line in Fig. 8 lead to cooling, those on the right to warming. It is evident from this figure that only some very dark urban aerosols can lead to warming, and that anthropogenic aerosols remote from their sources with typical ~os values in the range of 0.85-0.95 will have a cooling effect unless they are over clouds, snow or ice. Soil dust aerosols typically have even higher ~os (0.97-0.99; CLARKE and CHARLSON, 1985), although very low values of 0.86 have been reported for the highly absorbing red Saharan dust (CARLSON and CAVERLY, 1977). Thus, most soil dust aerosols will also result in cooling, especially if they are blown over the oceans or over vegetated areas. Fig. 8 shows that dust aerosols over desert regions are near the dividing line between warming and cooling. The strong size dependence of aerosol optical properties can introduce a large degree of uncertainty into calculations of the radiative effect of such aerosols (GRASSL, 1988). For example, CARLSON and BENJAMIN (1980) suggested that Saharan dust aerosols would have a warming effect on the net local radiation balance, whereas FOUQUART et al. (1987a,b) point out that they could lead to a net energy loss, i.e. cooling, depending on the amount of submicrometre-size particles present. It is obvious from this discussion that there is a great need for additional measurements of the size distributions and size-dependent optical properties of atmospheric aerosols. As we have shown, while the net effect of aerosols on the global radiation balance is, in most cases, in the direction of cooling, there is a significant amount of radiation absorbed by atmospheric aerosols, so that the locus of some of the solar energy absorption is shifted from the Earth surface to the atmosphere. For example, PENNER et al. (1992) estimate that biomass smoke aerosols absorb about 0.5 W m -2 (globally averaged and adjusted for cloud cover). The absorption of radiation leads to reduction of about 0.3 W m -2 in the solar radiation absorbed at the ground (adjusting for ground albedo and cloud cover). This leads to a net increase of 0.2 W m -2 in the heating of the surface/atmosphere system due to the absorbing component of the smoke. At the same time there is a 0.8 W m -2 shift of energy absorbed at the ground to energy absorbed within the atmosphere. Here, it leads to a local warming aloft, which influences the thermal stratification of the troposphere, as noted already by CHARLSON and PILAT (1969). This effect is particularly relevant in situations characterized by optically thick aerosol layers, e.g. dust storms. In general, warmer air aloft results in more stable stratification, less convection and consequently also a reduction in the likelihood of precipitation from afternoon convective clouds. Furthermore, a reduction in the amount of heating at the ground (which results from the presence of any kind of aerosol, regardless of its COs) reduces the amount of water evaporated from the surface due to a reduction in the amount of energy available for latent heating. This in turn results in less cloudiness and precipitation, which, in principle, should act to reduce the cooling effect of clouds. MITCHELL (1971) estimated that this effect could lead to reductions in evaporation by some 1-3%, an amount which, if it resulted in a reduction of cloud cover by a similar percentage, would have a pronounced climatic significance. He concludes: "Further investigations of this and other possible hydrometeorological effects of

367

Climatic effects of changing atmospheric aerosol levels an atmospheric aerosol are clearly warranted", a statement which remains valid after 22 years.

Modelling the direct radiative effect of aerosols While no complete global models incorporating the temporal and spatial distribution of the various aerosol types have yet been assembled, some more sophisticated calculations have been made for specific cases. COAKLEY et al. (1983) used assumed aerosol optical properties and a latitudinally varying distribution of "natural background aerosol" in a latitudedependent climate model. They conclude that this background aerosol leads to a cooling of about 3 K, with the highest cooling rates at high latitudes (up to 5-6 K), and "that the present background aerosol exerts a rather substantial influence on the present climate ... comparable in magnitude but opposite in sign to what would be expected for a doubling of atmospheric CO2". It should be noted that their calculations are based on an assumed "natural background aerosol" which is, however, based on data from anthropogenically influenced regions in the Northern Hemisphere, and that their model actually "is representative of a symmetric planet for which both hemispheres correspond to the Earth's Northern Hemisphere". If we take into account that in the present view the submicrometre aerosol loading in the Northern Hemisphere may be about twice the natural level, it follows that anthropogenic aerosols could be cooling the Earth by about 1-2 K at present by the direct radiative effect alone. First attempts to introduce aerosol effects into three-dimensional models at fixed aerosol loads and unchanged cloud properties also showed cooling (TANRg et al., 1984). GRASSL (1988) discussed the combined direct and cloud-optical effects of anthropogenic aerosol input using 2-D model calculations with assumed aerosol size distributions and refractive indices. Based on calculations by NEWIGER (1985), REHKOPF (1984) and GRASSL and LEVKOV (1984), he concluded that in cloudless areas aerosol can lead to zonal-average negative forcings up to -7 W m -2 in the northern mid-latitudes, while the increase in cloud albedo leads to a forcing up t o - 3 . 0 W m -2, with a global mean forcing of about -(0.7-1.0) W m -2 by the combined effects. Effects were significantly greater in summer than in winter. CHARLSON et al. (1991) obtained a global distribution of sulphate aerosol from the threedimensional chemical-transport model of the sulphur cycle by LANGNER and RODHE (1991). They used the radiation calculations described above to obtain a global distribution of sulphate aerosol radiative forcing by natural and anthropogenic sulphates. They estimated that natural sulphate aerosols produce a global mean forcing of-0.42 W m -2, and that anthropogenic sulphates add 0.11 W m -2 in the Southern and 1.07 W m -2 in the Northern Hemisphere. Peak anthropogenic forcings (up to 4 W m -2) were predicted over the eastern US, southeastern Europe, the Near East, and China (Fig. 9).

The most detailed modelling study on the direct effect of aerosols to date was made by KIEHL and BRIEGLEB (1993), who investigated the influence of greenhouse gases and natural and anthropogenic sulphate aerosols on radiative forcing in a three-dimensional diagnostic climate model. They also obtained the abundance and distribution of the sulphate aerosol from the chemical-transport model of LANGNER and RODHE (1991). While they find a geographical distribution of the anthropogenic radiative forcing similar to that of CHARLSON et

368

Climatic effects , t l l i l l i l l i l t i l l i l , , l , , l , , l , , l , , l , ,

90~

~

,.

1~ ~ . ~

.....

J_,.,

_...

00 ~

30*

'3

,

'

3

'

..

. ...-.~. 9

30* 60*

900S f-';"'~'"i''~ , I ' ' 1800W

w,,

120 ~

I,,

i,,

60 ~

I,' 0~

I,,

I,,

i,,

60 ~

I,' 120 ~

w' ' 1800E

Fig. 9. Calculated increase of reflected flux to space due to tropospheric sulphate aerosols derived from anthropogenic sources (in units of W m-2) (from CHARLSONet al., 1991).

al. (1991), the magnitude is generally less by a factor of 2. About half of this difference is because CHARLSON et al. (1991) did not account for the wavelength dependence of the extinction coefficient for sulphate aerosol. Much of the remaining discrepancy is due to differences in the asymmetry parameter (the ratio between upward and downward scattered radiation) used by the two groups. KIEHL and BRIEGLEB (1993) estimate a global mean anthropogenic sulphate forcing of 0.3 W m -2, and a Northern Hemisphere forcing of 0.43 W m -2. They find that over considerable areas of the Northern Hemisphere the negative forcing from aerosols nearly compensates the greenhouse gas forcing. In some regions of Europe, the eastern United States and China, aerosol negative forcing even exceeds the positive forcing by the gases and results in negative net forcing. The preceding studies only include the effect of sulphate aerosols and do not account for other anthropogenic aerosols, particularly those from biomass burning. Inclusion of these sources can be expected to roughly double the sulphate aerosol effect (PENNER et al., 1992). A 3-D modelling study including aerosols from other sources is highly desirable, and must take into account the differences in the source distributions and optical properties of sulphates, biomass burning smoke, and other materials. Overall, we can estimate that the direct radiative forcing from anthropogenic aerosols falls somewhere between 0.6 and 2.5 W m -2, with a best guess probably near 1 W m -2. Indirect

effects through

perturbation

of cloud properties

A substantial fraction of incident solar radiation (some 27%) is reflected by clouds, which cover on average about 60% of the Earth. Satellite measurements show that the average albedo of clouds is about 0.45, and that most clouds have albedos in the range 0.25-0.65 (COAKt.EY and DAVIES, 1986). The albedo of a cloud depends on a number of properties: its thickness, liquid water content, droplet number concentration and size distribution, and solar elevation. Obviously, other things being equal, the more liquid droplets that lie in the path of the radiation, the higher the likelihood is that photons get scattered backwards and leave the Earth towards space rather than being transmitted to the Earth's surface. Thus, everything else being constant, the albedo of a cloud increases, and its transmittance decreases

369

Climatic effects of changing atmospheric aerosol levels with its thickness. On the other hand, if we keep the cloud thickness constant, but increase the liquid water content (LWC, the amount of liquid water per volume air), albedo also increases and transmittance decreases. Finally, at constant liquid water path (LWP, the vertically integrated liquid water content), cloud albedo also depends on the droplet size distribution. Since the scattering interactions, which result in the backscattering of light by clouds, happen at the liquid-gas interface of the droplets, an increase in the interface area in the light path leads to an increase in the amount of light scattered. At constant liquid water path, this can be effected by a reduction in the average drop size, which results in the presence of more droplets within a given air volume. Liquid water content (L), droplet number concentration (N), and effective radius (re) are related by

L= 4 ~re3N

(5)

The relationship between cloud albedo (ACT), sun angle and cloud optical properties can be summarized by the following equation:

ACT =

fl(#o)6c //~o 1 § fl(~o)Sc / IZo

(6)

where/t o is the cosine of the solar zenith angle, fl(/t0) is the fraction of light scattered in the upward direction for sunlight incident on the cloud at angle/t o (for single scattering), and 6c is the optical depth of the cloud (COAKLEY and CHYLEK, 1975; CHARLSON et al., 1992). The upward fraction fl is a slowly varying function of/~o and 6c (STEPHENS, 1978). Cloud optical depth is dependent on the effective radius (re) of the cloud droplet population, the number concentration of droplets (N), cloud thickness (Zc), and the average extinction efficiency of the droplets (Qext) which for cloud droplets of a radius much larger than the wavelength of visible light can be approximated by a constant value of 2 (CHARLSONet al., 1992): c = ~r2 QextZc

(7)

or, inserting equation (5),

0 C = ~QextZc

N 113

(8)

so that, at constant LWC, the cloud optical depth depends on the cube root of the droplet concentration. The number of cloud droplets in a cloud is regulated by the concentration of cloud condensation nuclei (CCN), i.e. aerosol particles which are soluble enough and large enough to serve as nuclei for the growth of water droplets from supersaturated water vapour. The presence of CCN is critical, since the formation of droplets purely from gaseous water molecules requires that an energy barrier associated with the very small radius of curvature of the in-

370

Climatic effects cipient droplet must be overcome. This would only happen at supersaturations of some 300%, which are never reached in the real world since the condensation of water vapour on a pre-existing particle involves a much lower energy threshold, and thus happens at much lower supersaturations (some tenths of a percent). Consequently, when a rising air mass exceeds the water vapour saturation level (i.e. a relative humidity of 100%) by a certain critical level, condensation of water begins on the largest and most soluble aerosol particles. This critical supersaturation level is a function of the mass of soluble material in a given aerosol particle (Fig. 10). Below this level, the particle remains stable as a hydrated aerosol particle, above it the nucleus grows rapidly into a cloud droplet. The fraction of particles which become involved in the formation of cloud droplets depends on the size spectrum of the aerosol, the chemical composition (solubility) of the particles, the rate of ascent of the air mass, and the total number of CCN present. CCN number concentrations vary over a wide range, from a few tens per cm 3 over the remote oceans, especially during the winter season, to about 100-200 cm -3 over warmer ocean areas, to hundreds and thousands per cm 3 in continental areas, depending on the level

2.0

-

1.5

-

Volume

E

/'N

accumulation m

/

\

coarse

1.0o

>,

0.5 0 0.01

I

I

I

I

I

~

I

0.1

I

I

I

I

I

Ill

I

(a)

500-

exl) O

/ \

400 _

300

'X:l

/

/

200 - /

I

I

I

I

Ill

10

Number

|

E

I

1 Diameter, D (gm)

/

nuclei mode \ \ \ \ ~~

m

o

d

accumulation e

\\,

100 0

i

0.01

i

0.1

i

i

i

i

!

i

i ii

10

Diameter, D (~tm)

I

i

6

0...~5

(b)

1

i

0.2 0._._21

i

I

0.005

0.0002

Critical Supersaturation, So,par (%)

Fig. 10. Typical marine aerosol size distribution. Both (a) and (b) are based on three overlapping, log normally distributed particle populations with volume mean diameters D of 3.0, 0.3 and 0.03/zm; geometric standard deviations of 2.0, 1.6 and 1.5; and total numbers (per cm 3) of 0.4, 100 and 200 for the coarse, accumulation and nuclei modes, respectively. The data are displayed so that the area under each curve is proportional to total particle volume (a) or number (b). Also shown at the bottom of (b) is the critical supersaturation corresponding to each particle size (assuming a composition of pure ammonium sulphate). The vertical lines show the range of typical maximum supersaturations for marine stratiform clouds (from ANDERSONet al., 1992).

371

Climatic effects of changing atmospheric aerosol levels of pollution (WARNERand TWOMEY,1967; HOBBS et al., 1970; BRAHAM, 1974; SCHMIDT, 1974; TWOMEY et al., 1978; HOPPEL, 1979; SAX and HUDSON, 1981; HUDSON, 1983; GRAS, 1990; HUDSON, 1991; HUDSON and FRISBIE, 1991; ANDREAE et al., 1994, 1995; ANDREAE, unpublished data). Sea-salt particles make up only a small fraction of CCN in the marine boundary layer; their number concentration is typically between 1 and 10 cm -3 (RADKE, 1968; HOBBS, 1971; PRUPPACHER and KLETT, 1978). The rest is predominantly made up of sulphate particles, which over pristine ocean areas are produced from the photooxidation of DMS (ANDREAE,1990; CHARLSONet al., 1987). At CCN concentrations of up to a few hundred per cm 3, a large fraction of the CCN present develops into cloud droplets (BIGG, 1986), while at higher concentrations (> 1000 cm -3) the competition for available water vapour limits the number of particles which are activated to grow into cloud droplets (LEAITCH et al., 1986). Unfortunately, there are few studies which provide simultaneous measurements of CCN and droplet concentrations in stratus clouds. Most of the early work (e.g. TWOMEY and SQUIRES, 1959; TWOMEY and WARNER, 1967) was done in convective clouds, where changing droplet number concentrations would have relatively little climatic effect. But where such measurements have been made, they confirm the influence of CCN concentrations on cloud droplet numbers. HUDSON (1983), for example, found good agreement between CCN and cloud droplet concentrations in marine stratus off the California coast. Similarly, HEGG et al. (1991) observed a highly significant correlation between cloud droplet numbers in marine stratus and CCN concentrations in the boundary layer in the range of 25-120 cm -3. PUESCHEL et al. (1986) investigated the presence of cloud droplets and interstitial aerosol particles in relation to air mass origin at Whiteface Mountain, New York State. In clean marine air masses from the North, 50% of a total of 130 particles per cm 3 had been activated, in background continental air 37% of 350 cm -3, and in polluted air masses 17% of 4350 cm -3. (The non-activated particles measured in this experiment were in the size class >0.2/tm diameter, and would probably have been registered as CCN by a CCN counter.) Thus, while there is no simple linear relationship between the CCN concentration present in a given air mass and the number of cloud droplets that will form if it becomes supersaturated, it is obvious that an increase in CCN will lead to higher cloud droplet concentrations, in this case from 64 to 750 cm -3 when going from clean to polluted air masses. Similar conclusions were drawn by LEAITCH et al. (1992), who investigated the relationship between cloud water sulphate concentration (as an indicator of the degree of anthropogenic pollution) and cloud droplet number over Ontario, Canada, and found highly significant, but far from linear correlation. In these relatively polluted air masses with droplet number concentrations typically in the range 110-510 cm -3, droplet number increased approximately with the 1/5 power of cloud water sulphate concentration. Clearly, cloud optical properties are most sensitive to variations in CCN concentrations in relatively clean areas of the world, especially the remote oceans, while in highly polluted areas cloud droplet numbers are nearly saturated and further addition of aerosol will have much less consequence for cloud properties. Besides changing the albedo of clouds, variations in CCN concentrations and the resulting changes in droplet size also influence the probability that raindrops can form from the cloud droplets. A cloud containing a large number of small droplets is likely to persist longer and eventually evaporate again without producing rain than a cloud in which the same amount

372

Climatic effects of water is distributed over a smaller number of larger droplets. Thus, an increase of CCN concentration increases the lifetime of clouds, and consequently the fraction of the Earth that is covered by cloud (BRAHAM, 1974; FITZGERALD and SPYERS-DURAN, 1973; ALBRECHT, 1989; TWOMEY,1991). This would also lead to an increase in the Earth's albedo, and cause cooling. Since the probability of precipitation is also changed, enhanced cloud stability would also influence the distribution of water vapour in the atmosphere. This effect can be illustrated by SIMPSON and WIGGERT's (1969) model calculations for clouds with "virtually identical dynamics" but differing CCN concentrations: At N = 50 cm -3 the cloud produced 1.5 g m -3 of rainwater, whereas at N = 2,000 cm -3 it precipitated only 0.19 g m -3. Such a range in CCN concentrations is typical for the present-day atmosphere, with the lower values being typical for clean marine areas, and the higher values characteristic of mildly polluted continental regions. Water vapour is the most important greenhouse gas, and is responsible for a large fraction of the infrared radiation retained by the atmosphere. Consequently, a large-scale change in the horizontal and vertical distribution of water vapour has the potential for significant changes in the energy budget. These effects are so complex, however, that they are beyond our current modelling capabilities. In the case of marine stratus, an increase in CCN should reduce drizzle production, resulting in higher cloud LWC. Drizzle is very commonly associated with marine stratus clouds, especially under conditions of very low CCN concentrations, as for example during winter at Cape Grim, Tasmania (E. K. BIGG, personal communication, 1987). In addition to regulating the LWC of stratus clouds, drizzle also has an important effect on the diurnal cycle of marine clouds due to the cooling effect of evaporating drizzle droplets underneath the cloud (NICHOLLS, 1987). This evaporation may create a stable layer below the cloud and thus uncouple the cloud from the fluxes of heat and water vapour from the sea surface. As a result, the cloud evaporates more easily due to solar heating, and only forms again in the evening. A reduction of drizzle due to increased CCN would disrupt this cycle and stabilize daytime marine stratocumulus, enhancing cloud cover and consequently albedo. Drizzle is the dominant sink of CCN at low CCN concentrations, since other sink mechanisms (coagulation and coalescence) are not effective at low particle densities. Thus, the CCN concentration in clean marine air may be controlled by the relationship between particle production and removal by drizzle (BAKERand CHARLSON,1990). At higher rates of CCN production and consequently higher CCN levels, drizzle production is suppressed, and CCN concentrations must rise until they are limited by coagulation and coalescence removal. These relationships may explain the tendency for CCN numbers to fall into two distinct ranges, one around 20-200 cm -3 and another around 1000 cm -3. Field evidence that the inhibition of drizzle production by elevated CCN may have an important influence on the LWC in marine stratus comes from an analysis of the enhancement of cloud albedo due to ship's emissions (RADKEet al., 1989; ALBRECHT, 1989). Not only were cloud droplets smaller in the ship tracks, but liquid water concentration was higher by 40-100%. This LWC increase had a two times greater effect on cloud reflectivity than that of the change in droplet size. Clouds interact not only with the incident shortwave radiation, but also with the infrared radiation emitted by the Earth. For this reason, they also have a greenhouse function. Therefore, an increase in cloud density or fractional cloud cover should, in principle, also have a warming component in addition to their cooling effect. However, this warming effect is very

373

Climatic effects of changing atmospheric aerosol levels small, since tropospheric clouds are nearly opaque to infrared radiation already, and a further increase in absorbance has little additional effect. A fractional increase in low clouds caused by increased CCN concentrations would result in a proportional increase in longwave absorption, but for low clouds with temperatures relatively close to that of the Earth's surface, the cooling component dominates.

Estimating indirect radiative forcing The effect of changing CCN concentrations on global climate was assessed quantitatively by CHARLSON et al. (1987) in connection with the discussion of a hypothesis which linked the production of biogenic sulphur gases by marine phytoplankton to the concentration of sulphate aerosol in the atmosphere, and thus to climate (Fig. 11). These authors used the relationships between droplet concentration, effective droplet radius, and cloud thickness given in equations (5) to (8) above to predict the changes in cloud albedo resulting from changes in CCN, and showed that under the conditions of the clean marine atmosphere, modest changes in CCN concentrations would significantly influence cloud albedo. This is summarized in Fig. 12, which shows the change of cloud albedo as a function of the original albedo of the reference cloud and the change in CCN number density. These results indicate that cloud-top albedo is most sensitive to change in cloud droplet number (N) when the original albedo of the cloud is in the range 0.2-0.7, i.e. for clouds of average and below-average al-

Radiation budget

1-

Cloud condensation nuclei

Global temperature

C- -3

Sulfate aerosol

Climate feedbacks

+

S02 DMS

I

DMS

d, + o r - ?

+ or-

Phytoplankton abundance and speciation

~

9 "

e Ocean

Marine ecology

Fig. 11. Climatic feedback loop linking oceanic phytoplankton, emission of DMS, atmospheric sulphur chemistry, cloud albedo and climate. The plus and minus signs indicate the sign of the feedback.

374

Climatic effects

1.0

Relative number-density of cloud droplets N/N o 6 4 3 2 1.5 1.00.8 0.6 0.4 0.3 0.2 0.15 +0.(

II

0.9

+0.02 +0.04

0.8

o

0.7

O O

0.6

~

o.5

0

~ 0

0.4

o

0.3

(D

0.2 O o~,,i ~

0.1 I 0 0.5

'

0.6

Albedo of ocean surface -[ "', . 0.' ' I ' ' ' ' ' '' 07 .8 09 1.0 1.2 1.4 1.6 1.8 2.0

Relative effective radius of droplets reff/reff o Fig. 12. Change (Aa) of visible albedo (0.6-/tm wavelength) at cloud top caused by changing droplet number-density N while holding vertically integrated liquid water content (liquid water path, LWP) constant. Aa is plotted as a function of the albedo of the reference cloud and the effective radius (reff, the surface-area-weighted mean radius) of the dropsize distribution relative to that of the reference cloud (reff0) (from CHARLSONet al., 1987). bedo, such as stratus clouds. For such clouds, an increase in N by 1% would increase cloudtop albedo by about 0.08%. CLAW estimated that a change in CCN concentration over the oceans by +30% would result in a decrease in the effective droplet radius by 10%. This would change the albedo of low stratus and stratocumulus clouds (the type of cloud most sensitive to this effect) by +0.020 (2%) at the top of the cloud and for the wavelength region 500-700 nm. Adjusting for the absorption in the 0.6/tm band of ozone in the layer of air between the top of the cloud and the top of the atmosphere reduces this albedo change to +0.018 (WISCOMBE et al., 1984). Finally, when the entire solar spectrum is considered, the increase in top-ofatmosphere (TOA) albedo over the cloud is reduced to 0.016, since cloud albedo is lower in the near-infrared than in the visible part of the spectrum (SHINE et al., 1984; WISCOMBE et al., 1984). On the basis of a more accurate method of resolution of the radiative transfer equation than the simplified approach used by CLAW, FOUQUART and ISAKA (1992), come to essentially the same estimate of the effect of a 10% change in effective radius on TOA albedo. When the area covered by the type of cloud modelled here is taken into account, an average albedo change for the total surface area of the Earth of +0.005 is calculated. As discussed above for the case of the direct effect of aerosols, there is a linear relationship between a change in global mean albedo and the resulting change in radiative forcing: c~Fc = 1

375

FTOAGM

(9)

Climatic effects of changing atmospheric aerosol levels Consequently, an albedo change of +0.005 would lead to a negative radiative forcing of -1.7 W m -2, or change in global-average surface temperature o f - 1 . 3 K. This assumes, of course, that a change in CCN and N does not alter the average area covered by clouds, i.e. effects on cloud lifetime are ignored here. Following the publication of the CLAW hypothesis, SCHWARTZ (1988) argued that, if this hypothesis were correct, the production of sulphate aerosol from anthropogenically released SO2 would have already led to a pronounced increase in CCN concentrations in the Northern Hemisphere. Based on an assessment of the rather uncertain database presently available on the concentrations of sulphate aerosol over the world oceans, and assuming that number concentration would be proportional to mass concentration, he postulated that CCN in the Northern Hemisphere were greater by 30% than in the relatively unpolluted Southern Hemisphere, and that therefore there should be a radiative forcing o f - 2 W m -2 due to the indirect effect from anthropogenic sulphate aerosols in the Northern Hemisphere. Schwartz's assumption that the enhancement in CCN number concentration is the same as the enhancement in mass concentration is probably not justified, since the particle number density decreases faster during transport than the mass concentration. CHARLSONet al. (1992) assume a 15% global mean enhancement in N, and obtain a global mean forcing o f - 1 W m -2 from cloud effects. They emphasize that there is a great amount of uncertainty (at least a factor of 2; KAUFMAN et al., 1991) related to the enhancement of CCN number concentrations and to the resulting enhancement of cloud-droplet number density. Most authors have considered only the anthropogenic sulphate aerosol as a source of increasing CCN concentrations. However, other aerosol sources, especially biomass burning, also release large amounts of aerosols which may act as CCN (WARNER and TWOMEY, 1967; HOBBS and RADKE, 1969; DI~SALMAND, 1987; RADKE, 1989; CRUTZEN and ANDREAE, 1990; HUDSON et al., 1991; ROGERS et al., 1991; PENNER et al., 1992). Smoke particles from biomass burning appear to be effective CCN in spite of the fact that they consist predominantly of organic material (HALLET et al., 1989; HUDSON et al., 1991). This is because soluble materials, like nitric, ammonia, and short-chain organic acids, are rapidly incorporated from the gas phase of the smoke into the aerosol and thus render it hygroscopic (ANDREAE et al., 1988, TALBOT et al., 1988). RADKE et al. (1991) estimated the global production rate of CCN from biomass burning to be in the range of 102~

particles per sec-

ond, which corresponds to a global average pyrogenic CCN concentration of 102-103 cm -3, assuming a CCN lifetime of 5 x 105 s and a scale height of 1 km. In laboratory combustion experiments with Guinean savanna vegetation, PHAM-VAN-DINH et al. (1994) found CCN (at 0.1% supersaturation) emission factors around 2 x 1011 g-1 dry matter, with the highest emission factors observed during the smoldering stage. This corresponds to a CCN production rate of (3-6) x 1019 s-1 for biomass burning in the African savannas. Based on airborne measurements over savanna fires in southern Africa, the global production of CCN from biomass fires is estimated to be about 2-4 x 102~s-1 (M. O. ANDREAE, unpublished data). Most of these CCN will be present over the continents, and a considerable concentration decrease must be expected during transport over the remote oceans, where they would have the strongest effect on climate. However, maps of the distribution of aerosol optical depth over the oceans clearly show plumes extending over thousands of kilometres from the biomass burning regions (HUSAR and STOWE, 1994). Therefore, PENNER et al. (1992) have estimated that the cooling effect of smoke aerosols from biomass burning via the enhancement

376

Climatic effects of cloud albedo is about 1 W m -2, if a mean pyrogenic CCN concentration of 10 cm -3 prevails over the remote oceans. This forcing due to biomass burning is of a magnitude similar to the cooling due to sulphate CCN. The soot content of smoke and industrial aerosols adds a light-absorbing component to cloud water, which in principle could reduce the albedo of clouds and counteract the whitening effect of larger droplet numbers (TWOMEY, 1977). The resulting increased absorption would have the strongest effect in the case of optically thick clouds, which have a high initial albedo and are not appreciably brightened by increasing droplet number. This prediction is confirmed by satellite observations made on Amazonian cumulus clouds, which showed that in spite of a reduction of droplet size from 15 to 9/zm due to dense smoke from biomass burning, cloud reflectance decreased from 0.71 to 0.68 (KAUFMANand NAKAJIMA, 1993). On the other hand, the albedo of warm, relatively thin clouds with an initial albedo of 0.30.4 increased by 0.08-0.15 in the presence of smoke aerosols (Y. J. KAUFMAN, 1993, personal communication). On a global scale, the effect of increasing albedo is thought to dominate over that of increased absorption (TWOMEYet al., 1984). This is supported by the observation of increased albedo in marine stratus clouds which have incorporated stack effluents from ships at sea (COAKLEY et al., 1987; KING et al., 1993). Radiative transfer calculations based on the measurement of light absorbing substances in cloud droplet nuclei in marine clouds impacted by polluted continental air also suggest that light absorption from soot does not significantly influence the cloud optical properties under normal circumstances (TWOHYet al., 1989). Direct evidence for a net cloud-brightening effect of aerosols is also found in satellite observations, which show higher cloud reflectance in areas with elevated aerosol loadings (DURKEE, 1994). However, in a satellite study of low-level cloud albedo over the North Atlantic, FALKOWSKI et al. (1992) was able to observe increased cloud albedo due to anthropogenic sulphate only over the area immediately adjacent to the North American coast. Marine sulphur emissions appeared to govern the variability in cloud albedo over most of the North Atlantic. As in the case of the direct aerosol effects, the cloud effect is very unevenly distributed in space and time. Obviously, it is limited to cloudy regions, while the aerosol effect is significant only in cloud-free areas. However, an assessment of the regional distribution of the impact of the enhancement of cloud albedo is made extremely difficult by the very strong dependence of the albedo increase on the absolute CCN concentration. Figure 13 shows the dependence of albedo susceptibility (dA/dN, the increase in albedo per unit increase in CCN concentration) on CCN concentration and cloud albedo. This figure shows that the strongest effect would result from an even relatively slight increase in CCN in the cleanest and most remote parts of the world. Since the aerosol or its precursors can travel over considerable distances (thousands of kilometres) from their sources, especially in the free troposphere, the effects could be on a hemispheric scale. While the mass concentration of anthropogenic aerosols at such large distances might be low, an addition of even some 10-20 CCN cm -3 into a clean marine air mass with some 50 CCN cm -3 would have a considerable effect on cloud albedo. At the present time, it is usually suggested that the impact of increasing CCN is felt predominantly in the Northern Hemisphere, although the effect of smoke from biomass burning as well as SO2 emissions from increasingly industrializing regions in the developing world may already significantly impact the Southern Hemisphere. Smoke from

377

Climatic effects of changing atmospheric aerosol levels

0.3

100 1000

Albedo

Fig. 13. Susceptibility of albedo to change in cloud droplet concentration (dA/dN) for different conditions. The vertical unit is per cent reflectance per additional droplet cm -3 (from TWOMEY, 1991). The horizontal axes are droplet concentration (N, in cm-3) and initial cloud albedo. biomass burning is detectable almost everywhere over the remote oceans, especially in the areas downwind from Africa and South America. Evidence for this is the presence of soot aerosols and correlated levels of submicrometre potassium-rich particles over the remote oceans (ANDREAE, 1983; ANDREAE et al., 1984; CLARKE, 1989). Given the low CCN numbers typical of the remote marine atmosphere, an addition of only 10 pyrogenic CCN per cm 3 would be enough to cause a global cooling forcing of 1 W m -2 (PENNER et al., 1992). Cloud optical effects of aerosols through changes in droplet size and increased cloud stability have not yet been modelled in the necessary detail. Some calculations with different versions of the NCAR CCM model suggested strong effects resulting from relatively high increases in droplet concentrations. SLINGO (1990) suggested that a greenhouse forcing of ca. +4 W m -2 could be counterbalanced by an increase of 80-185% in the droplet concentration of low clouds, or by a 40% droplet increase in clouds at all levels. GHAN et al. (1990) increased the droplet concentration of marine clouds from 50 to 200 cm -3 and kept continental clouds at a constant droplet concentration of 200 cm -3, which resulted in a negative forcing o f - 6 W m -2. Their simulation suggests that this cooling effect would be reinforced by a concomitant increase in cloud fraction of 3.5%.

Estimate of total (direct plus cloud) anthropogenic climate forcing by aerosols Our estimate for the present-day direct forcing by anthropogenic aerosols ranges between -0.6 W m -2 (based on the Kiehl and Briegleb approach, with aerosol sources other than sulphate added) to -2.5 W m -2 (Table II, based on CHARLSON et al., 1991, radiation equations). The biomass burning term may have to be discounted by one-third, based on the argument that while at present practically all biomass burning is anthropogenic, there was about 1/3 of the present level of emissions already present around 1850 (ANDREAE, 1993). This reduces the range of estimates for the direct effect to -(0.5-2.0) W m -2. The cloud-optical effects of sulphate and pyrogenic aerosols have each been estimated to be of the order o f - 1 W m -2 (CHARLSON et al., 1991; PENNER et al., 1992). Discounting for pre-1850 biomass burning,

378

Evidence for the climatic impact of anthropogenic aerosols and considering that these estimates are uncertain by at least a factor of two, we obtain a range o f - ( 0 . 8 - 3 . 3 ) W m -2. Finally, adding direct and indirect terms yields a range of -(1.3-5.3) W m -2 for the total aerosol effect. Using a two-dimensional climate model and rather conservative assumptions, KAUFMAN and CHOU (1993) estimate a value o f - 0 . 4 5 W m -2 for the total radiative forcing due to aerosols from anthropogenic SO2 emissions. Given that sulphate accounts for about one-half of the anthropogenic forcing, the total effect based on the KAUFMAN and CHOU (1993) model would be about-0.9 W m -2, slightly below the lower limit of the range proposed here. It is clear, however, that even the lowest estimates are of significance relative to a present-day greenhouse forcing of 2.3 W m -2, especially considering that we have ignored the possibility of a LWC increase in our estimate of cloudoptical forcing, which may enhance this effect considerably.

Evidence for the climatic impact of anthropogenic aerosols Unfortunately, the influence of changing aerosol levels on climate cannot at this time be documented based on direct measurements of trends in global albedo, cloud albedo, or cloud area (CHARLSON, 1988). While there is a clear record of increasing cloud area by as much as 10% in North America (HENDERSON-SELLERS and MCGUFFIE, 1988, 1989), this could be explained by increasing CCN levels as well as by a number of other factors. Planetary albedo has been measured from satellites for three decades, but trends at the required resolution cannot be extracted from these data due to changes in sensor characteristics and calibration. Possibly, detailed analysis of data from the Earth Radiation Budget Experiment may provide some insight into the relationships between clouds, aerosols and climate, but this remains to be accomplished. For the time being, we are limited to more circumstantial evidence for an examination of the linkage between aerosols and climate.

Aerosol trends The growth of anthropogenic aerosol sources over the last 150 years is indisputable, but direct evidence for a global increase in aerosol concentrations in the atmosphere is rather scarce. Due to the large variability of aerosol concentrations in space and time, secular time trends are much more difficult to observe than for the long-lived and relatively evenly distributed greenhouse gases. Measurements of optical depth at remote mountaintop sites over several decades have not shown any significant trends (ELLIS and PUESCHEL, 1971; ROOSEN et al., 1973; CHARLSON, 1988). There is, however, a clear change evident in some statistical analyses of aerosol or haze occurrence in more directly polluted areas (YAMAMOTO et al., 1971; HUSAR et al., 1981). For example, HUSAR and WILSON (1993) document a considerable increase in haze over the US over the last three decades. Satellite measurements of light-scattering over the North Atlantic have shown that this haze can be transported for great distances over the oceans and still produce an easily detectable impact on the scattering properties of the atmosphere (FRASER et al., 1984; HUSAR and STOWE, 1994). One of the most salient predictions from recent model studies of the global sulphur cycle is the pervasive sulphate pollution in the troposphere over the North Atlantic and Pacific (LANGNER et al., 1992; TARRASON and IVERSEN, 1992). Field measurements of aerosol

379

Climatic effects of changing atmospheric aerosol levels chemistry in remote regions confirm the influence of anthropogenic emissions in areas far from continental sources (LAWSON and WINCHESTER, 1979; ANDREAE et al., 1984; WARNECK, 1988; SAVOIE and PROSPERO, 1989; CHURCH et al., 1991; DUCE et al., 1991). Based on aircraft measurements, ANDREAE et al. (1988) showed that mid-tropospheric sulphate concentrations over the North Pacific off the west coast of the United States were dominated by long-range transport from Asia. This was established by air-mass trajectory analyses and supported by radon measurements. Analyses of the interrelationship between ozone, nitrate and sulphate concentrations in North Atlantic air masses sampled at Barbados also showed pervasive influence of anthropogenic sources (SAVOIE et al., 1989, 1992b). Shipboard measurements of aerosol and precipitation chemistry over the Atlantic document widespread enhancement of sulphate levels above the biogenic background (CHURCH et al., 1991). Long-term ground-level measurements of non-sea-salt sulphate concentrations at island stations also show clearly elevated concentrations over the Northern Hemisphere oceans relative to those in the Southern Hemisphere at corresponding latitudes (Table III). This hemispheric difference is also clearly seen in the relatively few airborne measurements available in remote areas (Table IV). The increase in CCN due to anthropogenic emissions is well documented in industrialized regions (WARNER and TWOMEY, 1967; HOBBS et al., 1970; BRAHAM, 1974; SCHMIDT, 1974; TWOMEY et al., 1978). That elevated sulphate levels over the North Atlantic cause an increase in CCN is suggested by the data of HOPPEL (1979), who finds that CCN concentraTABLE III ANNUAL AVERAGESURFACECONCENTRATIONOF Nss-SO4 2- IN AIR (PMOL/MOLAIR) (MODIFIEDFROM

LANGNERANDRODHE, 1991) Location

Latitude

Longitude

76~ 79~ 52~ 32~ 30~ 20~ 13~ 12~ 8~ 3~

87~ 20~ 174~ 65~ 175~ 158~ 60~ 144~ 133~ 160~

210 168 155 455 133 119 175 119 147 154

BARRIE and HOFF(1985) HEINTZENBERGand LARSSEN(1983) PROSPEROet al. (1985) WOLFFet al. (1986) SAVOIEand PROSPERO(1989) SAVOIEand PROSPERO(1989) SAVOIEet al. (1989) SAVOIEand PROSPERO(1989) SAVOIEand PROSPERO(1989) SAVOIEand PROSPERO(1989)

15~ 22~ 30~ 40~

170~ 166~ 168~ 144~

68~

61~

84 98 56 21 b 26c 21 20

SAVOIEand PROSPERO(1989) SAVOIEand PROSPERO(1989) SAVOIEand PROSPERO(1989) AYERSet al. (1991) ANDREAEet al. (1991) SAVOIEet al. (1992a) TUNCELet al. (1989)

nss-SO42-

References

Northern Hemisphere Canadian Arctic a Spitsbergen Shemya Bermuda Midway Oahu Barbados Guam Belau Fanning

Southern Hemisphere Am. Samoa New Caledonia Norfolk Island Cape Grim Cape Grim Mawson South Pole

aMean for stations Alert, Mould Bay and Iglooik. bSubmicrometre fraction. tin fraction 5 m s-l), thermal effects are overwhelmed by friction effects: large downward momentum fluxes, flow deceleration and uplift result (BRYANT et al., 1978). This is consistent with the barrier effect, following an abrupt change in surface roughness in the absence of a thermal influence. The uplifted airmass can form a "plume" of ascending air which will be advected downwind. In the lee of the city, another change in roughness is experienced by the airflow (rough to smooth) sometimes resulting in descending, diverging air and enhanced momentum transfer. Sometimes a local windspeed maximum at ground level is found, as a result of this greater downward transport of horizontal momentum. A decrease in horizontal wind velocities at any height in the UBL is accompanied by an increase in Zg, the altitude at which the geostrophic wind speed is reached. The urban-induced deceleration perturbs the balance between the pressure gradient forces and the Coriolis force leading to cyclonic turning of the flow (BORNSTEIN and JOHNSTON, 1977; BRYANT et al., 1978). This feature was observed by ANGELL et al. (1973) over Oklahoma City, where tetroon balloons turned cyclonically upon entering the built-up area of the city, then turned anticyclonically in the lee of the urban area. An extensive study of turbulence in the urban surface layer was conducted by CLARKE et al. (1982). Turbulent intensities exceeded rural values by 50% as a result of the frictiongenerated shear stress at the rough urban canopy/air interface and the effects of buoyancy. Increased shear stresses over the urban land use were observed both during the day and extending into the night. The instability of the nocturnal UBL maintains turbulent exchange of momentum between the UCL and UBL. Other studies report similar enhancement of turbulent intensities in the UBL plus increased momentum fluxes and friction velocities (e.g. BOWNE and BALL, 1977; HOGSTROM et al., 1982; STEYN 1980, 1982; HILDEBRAND and ACKERMAN, 1984; YERSEL and GOBLE, 1986; ROTH et al., 1989a; ROTH and OKE, 1993). A recent analysis of surface layer turbulence structure by ROTH and OKE (1993) and ROTH (1993) are among the most complete and include the first published urban moisture spectra. Their turbulent wind results reveal many of the features discussed above. The vertical wind speed (w) spectra peak was shifted to lower frequencies, in agreement with other turbulent studies (CLARKE et al., 1982; HOGSTROM et al., 1982; STEYN, 1982). The shape of the ucomponent spectrum showed the influence of building wakes with the peak energy being at higher frequencies (smaller spatial scales) with a relatively faster roll-off. The co-spectra of w'T' and u'w' (e.g. COPPIN, 1979) and the turbulent velocity statistics all conform to MOST

(Monin Obukhov Similarity Theory), in agreement with other studies. The turbulent moisture spectra showed influences from larger scale structures, in particular surface evaporation appeared to be driven by downdrafts of dry air. These findings suggest non-similarity between heat and moisture transport processes. Many of these basic effects on turbulence are propagated into the convectively dominated UBL. GODOWITCH (1986) found that vertical velocities were larger in the UBL compared to

496

Current and future urban climates a rural boundary layer and variances in urban vertical velocities were 50% greater than their rural counterpart. The UBL was characterised by a greater production of vertical turbulent energy across a wide frequency range with the dominant length scales of the updrafts and downdrafts being comparable to the depth of the UBL. All these observations also suggest that the UBL obeys convective similarity scaling. In summary, this overview of urban airflow and energy exchange processes reveals that the combination of an aerodynamically rough surface, the juxtaposition of wet and dry surface elements within the UCL, and an increased active surface area, yields an environment that has the potential for large heat or moisture fluxes. If the urban canopy is wet, then large evaporation rates can prevail, even larger than in well-watered rural environments. Under dry conditions, however, most net radiation is channelled into turbulent or stored sensible heat fluxes. The division into "wet" and "dry" surfaces is linked to anthropogenic control of surface watering, hence sizeable latent heat fluxes can be maintained long after rainfall has ceased. Large turbulent sensible heat and momentum fluxes lead to large vertical velocities and turbulent intensities throughout the UBL. By day, this can lead to enhanced entrainment from the stable atmosphere aloft, which can increase the thermal anomaly in the UBL (see below). Surface turbulent fluxes are therefore strongly influenced by mesoscale advective effects through coupling between the UCL and the UBL. The partitioning of net radiation into sensible heat depends also on canyon geometry, canyon construction materials and solar angle. Limited evidence from tropical energy balance studies suggests that high solar elevation angles may enhance AQ s.

Current and future urban climates In this section we address the way that these fundamental atmospheric processes lead to characteristic urban climates. The focus is deliberately at the local (viz. land use) and UBL (viz. meso-) spatial scale and the diurnal to annual temporal scale. The climate elements include temperature, humidity, rainfall and severe weather. Wind is also an important "climate" variable, but urban effects on airflow were covered in the previous section.

Temperature Temperature anomalies associated with urban areas have long been recognised. Luke Howard first observed elevated air temperatures in cities in 1833. The existence of this "urban heat island" has now been well documented in human settlements ranging from villages to megacities and in geographical locations from the high latitudes (Fairbanks, USA) to the tropics (including, among many others, cities in Nigeria, Mexico, India and Malaysia). Urban thermal anomalies have been observed both during short-term, intensive measurement campaigns (e.g. OKE and EAST, 1971) and in longer term climatological studies (e.g. CAYAN and DOUGLAS, 1984; ACKERMAN, 1985, 1987; BALLING and BRAZEL, 1987a; FENG and

PETZOLD, 1988; KARL et al., 1988; CHOW, 1992). Many of these longer term studies (e.g. CAYAN and DOUGLAS, 1984; KARL et al, 1988) report a decrease in daily maximum tem497

Urban climates peratures in all seasons except winter and increases in daily minimum and mean temperatures as a result of urbanisation. Most urban heat island studies conform to a common methodology: the temperatures are near-surface air temperatures measured in the UCL; the reported heat island strength (ATu_r) refers to urban-rural temperature differences either at a specific time, or in terms of the daily maximum, minimum and mean temperatures. The rural temperatures are measured at a site presumed to be uninfluenced by the nearby city. Intensive studies would involve a mobile temperature survey across the city and out to the rural periphery, and urban heat island strengths would refer to the almost simultaneous difference between urban and rural temperatures. Longer term, climatological studies (such as those mentioned above, e.g. by KARL et al., 1988) may simply use just two sites, one representing rural and one urban, and only use daily or monthly maximum, minimum and mean temperatures. Other studies use just an "urban" site, where such a site may well be located at the nearest airport, raising the question of whether it is urban or rural. For example, BALLING and BRAZEL (1987a) found little evidence of a temperature increase in Tucson (USA) because, they suggest, of the type of land-cover that urban growth is replacing. The so-called "urban" station for this study was in fact a nearby airport, removed from the city of Tucson. The suburban/rural energy balance comparison for Tucson (GRIMMONDand OKE, 1990) indicated that local climate differences should develop between suburban Tucson and its rural periphery. Most studies of the UCL (air temperature) heat island have been conducted in mid-latitude cities with temperate climates. From these, a fairly consistent picture of the UCL heat island has emerged; the reader is referred to Landsberg's text (LANDSBERG,1981) and the review papers by OKE (1982) and OKE (1979a) for details of these studies. The thermal anomaly that exists in the UCL is propagated upwards, leading to a warmer UBL. The UCL heat island is best expressed at night during stagnant synoptic conditions, such as exist with an anticyclone. Thus with light winds and clear skies where radiative cooling is optimal and turbulence is damped, and in flat terrain with minimal mesoscale flows, the spatial and temporal pattern of the UCL heat island would be similar to the hypothetical pattern shown in Fig. 9. Embedded in the general pattern of increasing temperatures from the rural periphery to the urban core are microscale variations associated with differences in land-cover: parks appear as cool "ponds" and commercial districts as warm "islands". Other variations in the temporal and spatial structure of the UCL heat island arise because of heating variations due to topography or the presence of water bodies. Measurements show that the heat island results from differential cooling rates between urban and rural land use, where rural really means well-watered, short vegetation (see below). Most studies agree that the UCL heat island reaches its maximum strength 3-5 h after sunset. Apart from wind speed and cloud cover, heat island strength could be expected to depend on specific features of individual cities, such as their size, density, etc. OKE (1973) used urban population as a surrogate for city size and found a log-linear relationship between population and heat island strength. Subsequent observations from outside North America revealed different log-linear relationships for European, Japanese and tropical cities (Fig. 10) that arise because of differences in building materials, building densities and building heights. Plotting heat island strength versus sky view factor reveals that heat island strength also strongly depends on the urban geometry.

498

Current and future urban climates

(D

E _

A

B -

City core

1

_

I Built-up

I

area

B 1/'"

A ...-"""

v

Built-up

area

Fig. 9. Temperature structure in an urban heat island (after Ord~, 1982). The upper figure shows nearsurface temperatures along transect A-B, while the lower illustrates isotherms of heat island strength. The modelling work by OKE (1981), VOOGT and OKE (1991) and OKE et al. (1991) provides insight into the processes involved in urban heat island development. These studies represent three approaches to simulating urban heat islands: a scale model (OKE, 1981); a model canyon in a real boundary layer (VOOGT and OKE, 1991) and a numerical simulation model (Surface Heat Island Model, SHIM) (JOHNSON et al., 1991; OKE et al., 1991). OKE (1981) showed the role of canyon geometry in determining the UCL heat island by using a scale model of an urban canopy. OKE et al. (1991) used a numerical model (validated by JOHNSON et al., 1991) with various combinations of sky view factor, urban and rural thermal admit-

o 16

o

I

I

~ 14

.~ if) E

~

12

c-

-o 10 c-" ~ 8 ~

I

I

I

-

NorthAmerican J / ~ o o"" e"/'E~r~ .../~/~ f _ . ~ o o'~

-

9

-

6

r-

E E ~

4 2 0

et

Tropical

9

I

1

I

I

9149

I

I

10 100 1000 10000 Urban population ( thousands )

Fig. 10. Relationship between city size and UCL heat island strength (after JAUREGUI, 1986) for North American (closed circles); European (open circles) and tropical (triangles) cities.

499

Urban climates

tances, enhanced downward longwave radiation and anthropogenic heat fluxes to determine which factors are more important for urban heat island development. Their results highlight not only the influence of geometry on urban heat island development, but also the importance of urban thermal admittance, compared to rural, values. They suggest that some of the heat island strengths and temporal patterns observed in arctic and tropical cities can be explained by the contrast in thermal admittance between the surrounding landscape and the city. This is well illustrated by the results of JAUREGUI et al. (1992) where the strength of the urban heat island in Guadalajara (Mexico) had a seasonal dependence on the urban-rural contrast in thermal admittance. These contrasts will be largest in the dry season and smallest in the wet season. These results clearly show that urban thermal anomalies based on urban-rural intercomparisons depend on the nature of the rural surface. Recalling LOWRY (1977), the rural site is a surrogate for the pre-urban landscape. The different interpretations that result when using urban/rural as a surrogate for urban/pre-urban intercomparisons were illustrated by GRIMMOND et al. (1993). Their study compared suburban and rural air temperatures in and around Sacramento (USA), where the farmland surrounding the city is both irrigated and unirrigated. A maximum heat island strength of 4~

was observed in the comparison of the

suburban and irrigated rural site, but the suburban site appeared as a "cold" island by day (i.e. ATu_r was negative) when the unirrigated temperatures were used for "rural". The actual urban climate effect is probably better represented by the suburban-unirrigated comparison. TABLE VI HYPOTHESISED CAUSES OF THE URBANHEAT ISLAND(FROM OKE, 1987)

Feature of urban land use and urban activity

(a) Urban canopy layer heat island Canyon geometry - increased surface area and multiple reflection Air pollution and warmer UBL Canyon geometry - reduced sky view factor Domestic, industrial and traffic heat emissions Urban construction materials - increased thermal admittance Reduction in vegetation and increase in impervious surfaces Canyon geometry - increased aspect ratio and sheltering (b) Urban boundary-layer heat island Chimney and stack releases of heat UCL heat island- increased heat flux from roofs and canopy layer Increased sensible heat flux, effect of heat island and roughness Air pollution- increased aerosol absorption

500

Effect on atmospheric process, leading to a positive thermal anomaly

Increased absorption of shortwave radiation Increased radiative absorption and reemission of longwave radiation Increased longwave radiation from UCL building walls Increased anthropogenic heat source Increased sensible heat storage Reduced latent heat flux Reduced total turbulent heat transport

Anthropogenic heat source Increased sensible heat input from below Increased convective and mechanical turbulence hence entrainment of warm and dry air from above the UBL Increased absorption of shortwave radiation

Current and future urban climates The two "rural" sites yield quite different urban climate anomalies. The factors involved in the genesis of the urban heat island (both at the UCL and UBL scales) are summarised in Table VI, these were initially based on OKE's proposal (1982). As the preceding discussions have shown, many of these hypotheses have since been confirmed. Far fewer studies have been conducted outside of the mid-latitudes, viz. in the tropics or higher latitudes. We expect that urban climate processes will be different outside the midlatitudes because, firstly, the basic climate can be quite different. Seasons are determined primarily on the basis of rainfall, wind direction and cloudiness rather than temperature, hence the seasonal variation in the urban climate will differ from mid-latitude cities. Secondly, some subtropical cities are located in arid landscapes where irrigation of the surrounding lands is frequent (e.g. Phoenix, Sacramento, USA) or the landscape is desert or semi-desert (e.g. Kuwait City, Kuwait; Karachi, Pakistan; Cairo, Egypt). The nature of the urban effect will vary drastically between such locations, especially because of variations in the thermal admittance. In desert areas dust can raise the atmospheric turbidity, and lower K,I, in both rural and urban locations. Thirdly, the solar zenith angle is smaller at tropical latitudes, hence radiation absorption by rooftops becomes important and solar penetration into urban canyons is increased. Finally, artificial heating is not required, however in the wealthier cities or suburbs, artificial cooling demand will be high and anthropogenic heat fluxes potentially large (recall Table IV). Urban climate studies in the tropical latitudes are important because this is the region of future urban expansion. Thermal comfort may already be compromised in many tropical cities where the combination of an already warm and humid climate is exacerbated by excess urban heating. Increased illness and mortality plus excess energy consumption are just two of the potential impacts from the heat island effect in cities whose regional climates already predispose them to be close to the limits of thermal comfort. Unfortunately, the issue of urban climate does not rank particularly highly in many developing, tropical nations where the struggle to provide food, shelter and clean water is of more immediate importance. Nonetheless studies are emerging which describe the urban climate of tropical, urban cities. Some of these are reviewed next. ADEBAYO (1987) compared rural, suburban and urban temperatures in Ibadan (Nigeria), a low latitude, humid city. He found that urbanisation led to increased minimum and mean temperatures, and slightly reduced daytime maximum temperatures. Heat island strengths were larger during the dry season. These results are similar to other tropical locations, including Guadalajara, Delhi, Mexico City, Nairobi and cities in Malaysia. JAUREGUI (1987) compared urban and rural cooling rates for four cities in Mexico (three in inland valleys and one on the coast) using mean monthly temperature data. Although some results were broadly similar to mid-latitude cities, there were some important differences in the urban heat island timing: peak heat island strengths were observed at sunrise or shortly thereafter and cool islands developed in the afternoon. As reported in many studies, maximum heat islands were observed in the dry season. NASRALLAHet al. (1990) reviewed 23 years of annual mean, maximum and minimum temperatures in and around Kuwait City. They found only a "modest" urban heat island; increased minimum temperatures were not confined to sites within Kuwait City and so no significant differences between urban and desert minimum temperatures could be found. Differences in maximum temperatures between subur-

501

Urban climates

ban and desert sites were significant; the authors suggest that this could be the result of a growing urban heat island. In comparison to other arid cities, Kuwait City's heat island is very small. The authors argue that this is a result of the lack of trees; the similarity of thermal properties of construction materials between the city buildings and the desert; lower building heights; and the moderating influence of the Arabian Gulf. These studies show that urban heat islands exist in tropical cities, but they are less pronounced and the timing of the peak heat island strength is altered compared to mid-latitude cities. Heat island strengths and occurrence are greater in the dry season. Changes to surface energy partitioning in urban land use should be reflected in surface temperatures. For this reason, plus the wealth of spatial information provided, several studies (e.g. PRICE, 1979; VUKOVXCH, 1983; ROTH et al., 1989b) have investigated the spatial expression of the urban heat island using surface temperatures measured from satellites. GALLO et al. (1993), for example, found a linear relationship between the normalised difference vegetation index (NDVI) and observed urban-rural minimum air temperature differences for 37 cities in the USA. Satellite-derived temperatures, however, do not always reveal the same spatial and temporal patterns as air temperature surveys. Several common features found in satellite studies were summarised, and confirmed, by ROTH et al. (1989b). They noted, firstly, that heat island intensities were largest by day and smallest by night, almost the reverse of the diurnal cycle of the UCL air temperature heat island. Secondly, the spatial temperature structure was closely linked to the land use patterns (i.e. features such as parks, shopping centres) during the day, but not at night. ROTH et al. found that the warmest urban surfaces (by day) were not in the city core but in areas with large fiat-topped buildings or extensive pavement. Some of these discrepancies may simply be linked to the difference between a radiative surface temperature and near-surface air temperatures where the latter depend on the surface energy balance plus heating and cooling in an air volume. Satellites may also preferentially view horizontal surfaces such as pavements, rooftops and treetops. The amount of the three dimensional active urban surface "viewed" or sampled by a satellite sensor then depends on its view angle and the ratio of active to plan area of the urban landscape being sampled. GOWARD (1981) showed that rooftops have much lower thermal admittance values and thus will show a greater surface temperature amplitude which may explain the satellite thermal patterns. The discrepancy between the UCL air temperature heat island and that viewed from a satellite simply illustrates our limited knowledge of urban processes. But it also means that satellitebased observations of urban climate processes must be interpreted with care.

Humidity Much less is known about urban effects on humidity. Humidity is expected to differ between urban and rural locations because of the following factors: 9

smaller input of vapour as a result of a reduction in vegetation in the UCL and hence smaller transpiration rates (but note that the reduced atmospheric stability at night may allow transpiration and evaporation to continue into the night);

9 9

emissions of water vapour by industry and transport; intercepted water on impervious surfaces may lead to enhanced evaporation following rainfall;

502

Current and future urban climates * 9

anthropogenic control of water application; juxtaposition of wet and dry surfaces means microscale advection can enhance evapo-

ration. Like the thermal anomaly, urban humidity effects need to be separated into the UCL and UBL effects. Cities are often asserted to have lower relative humidities because of the misconception that urban areas are devoid of water. Most measurements of relative humidity in the UCL support these assertions but only because the relative humidity index depends both on atmospheric water vapour content and temperature. In fact, relative humidity is more sensitive to temperature changes than to humidity changes, so lower relative humidities often simply reflect the extra warmth of cities. Studies comparing actual humidity show that cities are often drier than their surroundings but there is a diurnal and seasonal dependence. HAGE'S (1975) study found, in summer, that urban areas were drier than rural by day and more moist by night. Seasonally, HAGE found that the vapour density was always greater in winter. Similar results (see OKE, 1979a) have been reported from other cities, e.g. in Chicago, ACKERMAN (1987) found lower vapour pressures and dew point temperatures in the forenoon and on spring afternoons. Urban-rural humidity differences, like temperatures, were found to be very sensitive to wind speed and cloud cover. For tropical Ibadan, ADEBAYO (1991) observed slightly reduced vapour pressures, especially in the afternoon, based on 2 years of data.

The UBL climate These thermal and moisture anomalies are not restricted to the UCL. The UCL warm, dry "island" extends upwards into the UBL and can be advected downwind. The temperature and humidity in the UBL also result from entrainment, advective influences and the effects of radiative heating and cooling. Measurements conducted in St. Louis in summer daytime conditions also showed a thermal excess and a specific humidity deficit in the UBL temperature profiles (ACKERMAN and MANSELL, 1978; TAPPER, 1990) which were displaced downwind. These thermodynamic anomalies varied inversely with wind speed

(SHEAand

AUER, 1978) and extended to heights of 500-1000 m (AUER, 1981); they also resulted in higher convective condensation levels. The deeper UBL slightly dilutes and reduces the heating impact. These observations were in summertime, anticyclonic conditions. In wintertime, or nocturnally, the UBL can be more moist. The drier UBL is due to both a decreased surface evaporation flux and enhanced entrainment of dry air from the stable inversion layer above the UBL, as demonstrated by HILDEBRAND and ACKERMAN (1984). They found that QH decreased with height in both urban and rural boundary-layers, but urban QH fluxes were always larger. Note that these large fluxes, accompanied by large vertical velocities in the UBL, enhance the entrainment process. Their urban/rural comparison found large rural QE fluxes at the surface which decreased with height in contrast to the large QE flux at the top of the UBL. Evaporation in the UBL thus increases with height and leads to a drying of the UBL.

Precipitation Two of the basic ingredients for raindrop formation; condensation nuclei and strong up-

503

Urban climates

drafts, are prevalent above and downwind of cities. It is therefore no surprise that cities have long been suspected of creating positive rainfall anomalies. CHANGNON (1968) identified the La Porte rainfall anomaly; HARNACK and LANDSBERG (1975) also observed enhanced summer rainfall over Washington, USA; HUFF and CHANGNON (1973) showed anomalies in precipitation and storm conditions for cities with populations greater than 1 million. Furthermore, the results of AYERS et al. (1982) show the potential for urban modification of rainfall. They found that the fluxes of cloud condensation nuclei (CCN) from Australian cities were an order of magnitude greater than the CCN fluxes generated by the continental landmass. The METROMEX study in St. Louis aimed to identify and explain links between summer rainfall anomalies and urbanisation. Summer rainfall was observed to increase from west to east across St. Louis, reaching a maximum northeast of the Mississippi River and the locus of the St. Louis urban heat island. Climate studies showed that although there was a peak in rainfall totals at this location, there was no observed shift in the diurnal pattern of rainfall. The rainfall maximum occurred at the site with the largest incidence of severe weather phenomena such as thunder and lightening. The major finding of the METROMEX project was that a precipitation anomaly arose from altered UBL dynamics that resulted from a range of urban influences, the most important including: 9 modified inputs of latent and sensible heat from the urban surface into the UBL; 9 input of cloud condensation nuclei from urban activities; 9 increased convergence and updrafts throughout the UBL due to perturbations in momentum and heat exchange across the city; 9 enhanced moisture input from industrial sources (BRAHAM, 1981, cited in CHANGNON, 1991). These combined influences enable cumulus clouds to grow taller and shift the droplet size spectrum to smaller median droplet diameters. The input of giant nuclei from industry, plus stronger updrafts, enhance the collisioncoalescence process with a resulting increase in rainfall. Nocturnal rainfall increases in St. Louis (CHANGNONand HUFF, 1986) were linked to urban influences on already heavy rainfall events. There are other studies that report the existence of "rain islands" in association with urban areas, including several from tropical cities. BALLING and BRAZEL (1987b) observed both increased rainfall amounts from, and frequency of, late afternoon and early evening storms in Phoenix (USA). They suggest rainfall anomalies may be linked to the urban heat island whose existence in Phoenix is well documented. However the failure to find any statistical significance in these observed differences highlights the difficulties in ascribing anomalies in rainfall occurrence and amount to an urban influence. Nonetheless, there is agreement that in some cities, positive rainfall anomalies exist. ELSOM and MEADEN (1982) examined the distribution of tornadoes in the Greater London district in the period 1830-1980 and found that the inner city shows a much smaller occurrence of tornadoes. They argue that the increased roughness of cities and excess heating combine to either dissipate weak tornadoes or suppress their formation altogether. They believe that the coincidence of an area of reduced tornado activity with the most densely populated area of London is significant. Similar tornado-free zones have been observed in Chicago and Tokyo (FUJITA, 1973, cited by ELSOM and MEADEN). Increased thunder activity was found over London (ATKINSON, 1969, 1971), St. Louis and other cities (CHANGNON

504

Future urban climates et al., 1977). ELSOM and MEADEN argue that the same factors that lead to enhanced convective precipitation and thunderstorm activity could also lead to a reduction in tornado activity and cite FUJITA's (1973) studies that show urban heating and frictional effects prevent tornadoes from forming. The thermodynamic anomalies in the UBL which influence convective and severe storms also have the potential to modify the passage of fronts. A classic study in New York City (LOOSE and BORNSTEIN, 1977) found that frontal movement was accelerated with a well developed urban heat island and retarded by the city's frictional influence if an urban heat island was absent. A series of observational and modelling studies in Tokyo have revealed that the movement of the seabreeze front through the Central Business District and suburbs is modified. The elevated air temperatures lead to local low pressures over Tokyo's Central Business District. The increase in pressure gradient found between the city core and the outer suburbs prevented the seabreeze from advancing through the city (YOSHIKADO, 1992).

Future urban climates

The previous sections have provided a general picture of urban effects on those atmospheric processes (radiation, heat, mass, momentum exchanges) that determine weather and climate. The urban influence on the atmosphere can extend vertically to the top of the PBL and, by day, can enhance the entrainment of air into the UBL. The regional wind advects this "plume" of warm, dry and polluted air downwind, enabling cities to "export" their atmospheric impacts to neighbouring areas. There is thus an urban "footprint" extending ca. 100 km downwind of a city. This does not include long range transport of atmospheric pollutants (which may extend to many thousands of kilometres). At the local scale, the urban temperature regime is characterised by higher daily minimum temperatures with no consistent change in maximum temperatures, thus daily mean temperatures are also elevated. The local scale climate tends to be warmer and drier, especially at night. There is a slight to medium reduction in the quantity and quality of solar radiation received at the top of the UCL, however these reductions are offset by enhanced longwave radiation. Within the UCL shading and specular radiation will dominate the microscale radiation regime. Wind speeds and directions are modified both within the urban canopy, where perturbed flow around buildings dominates, and in the UBL depending on the regional wind speed. Turbulence, and hence turbulent exchanges, are modified because of the altered thermal and mechanical properties of the urban surface. Cities tend to store more heat within their fabric, and in dry conditions heat is preferentially channelled into sensible rather than latent forms. These climate effects are primarily local to mesoscale in extent. The extra friction and heating generated by the UCL also modifies the occurrence and amount of rainfall, thunder and lightening, tornadoes and frontal movements. These effects may extend to regional scales. What will be the nature of future urban climates? Some have queried whether ongoing urban expansion will begin to modify continental and global scale climates. The total area of land in "urban use" was estimated as 1 million km 2 in 1980, growing at a rate of 2 • 104 km 2 year -1 (OKE, 1980). Thus in 1993 the total land area occupied by cities will be 505

Urban climates (very approximately) only 0.25% of the total surface area of Earth. It seems very unlikely that global-scale climate change could result from the modifications to albedo, roughness and moisture availability as a result of such a small fraction of urban land use. It is instructive here to review our current knowledge of the effect of cities on climates at this larger scale. The search for a greenhouse-induced signal in the near-surface air temperature record for the globe has focused considerable attention on the issue of urban warming and its large scale impacts. The US has the longest and most spatially comprehensive measurement network and so most understanding relates to urbanisation in the US. In 1986 KUKLA et al. analysed the US temperature record from 1941 to 1980 using paired urban/ rural meteorological sites and found that urban stations showed an average temperature increase of order 0.11 ~

Eliminating those station pairs suspected of introducing er-

rors made little difference to this urban warming figure, except for the small subset of stations that had no change in instrument or station location over the data record. These showed a warming trend of 0.34~ This range (i.e. 0.11-0.34~ encompassed that found in their survey of other studies of long-term urban warming. KARL and JONES (1989) suggested that the urban warming bias present in the entire US climate record was between 0.1 and 0.4~ in the 84 years from 1901 to 1984; JONES et al. (1989) estimated a similar-sized urban bias for the US which they extrapolated to 0.01-0.1 ~ in 84 years for the global temperature record. These studies show that in the period of rapid urban expansion in the US (1901-1984), urban areas may have warmed by between 1 and 3 ~ been estimated to have yielded a 0.1-0.4~ entire US.

This has

rise in the near surface air temperature for the

JONES et al. (1989) extrapolated their estimates of urban warming for the US to the globe, yielding a 0.01-0.1 ~ warming in 84 years for the global temperature record. This is an order of magnitude less than the observed warming, suggesting that urban impacts on the global temperature are not important. In summary, there is clearly some suggestion that a portion of the increased temperatures observed in the US is attributable to urban warming, however there is no evidence of a similar affect at the global scale. At the local scale, however, urban dwellers in the US have undergone "climate change" of a similar order to that being predicted by current GCMs to result from enhanced greenhouse warming. People living in cities in developed nations probably are exposed to a climate that has elevated minimum and mean temperatures; is potentially less humid; the sunlight they receive may be more diffuse, contain less UV radiation and be less intense; nocturnal temperatures are warmer; and rainfall and storm patterns may be altered. CHANGNON(1992) notes, interestingly, that this climate change has occurred by and large without comment. A social attitude assessment conducted in St. Louis (FARHAR, 1979, cited in CHANGNON, 1992), found that residents in that part of the city shown to have an altered climate were either unaware of the changed climate, or were unconcerned. While urban-induced climate change does not appear to be directly important to urban dwellers, it does affect their lives and environment in many indirect ways. Cities located in cold regions are slightly warmer during winter which saves heating costs at both the personal and environmental level; conversely more energy will be consumed in those cities already located in warm climate regimes. Microclimate effects within the UCL influence the level of pollutants, radiation and turbulent gusts to which pedestrians are exposed. In warm climates the thermal stress resulting from urban heating superimposed on pre-existing heat-wave conditions (which may be

506

Future urban climates

heightened even further due to global warming) can contribute to illness and premature mortality (MATZARAKIS and MAYER, 1991). It is unlikely that there will be any further, sizeable change in the urban climates of cities in developed nations because the period of rapid urbanisation is over. These cities will continue to grow and encroach on the surrounding landscape but they are unlikely to continue increasing their population and, if anything, urban consolidation may slow down this encroachment. There will therefore be some expansion of the urban footprint, but this is unlikely to greatly increase urban climate effects. SEAMAN et al. (1989) simulated the effects on mesoscale airflows of doubling the size of St. Louis. Surprisingly, they found an amelioration of urban effects, namely a reduction in the urban generated convergence and vertical velocities. Future changes in local and meso scale urban climates may arise from efforts to reduce urban sprawl and energy consumption in developed nations. Calls by governing bodies to increase building densities in urban areas through urban consolidation will lead to larger aspect ratios, greater sheltering and reduced SVF in the UCL. Using the relationship developed by OKE (1987), a doubling in H/W ratio from 1 (common for European, Australasian and North American cities, OKE, 1987) to 2 (towards the upper end for some North American cities) leads to a decrease in SVF from ca. 0.5 to 0.3. This would be accompanied by a 3~ increase in the maximum heat island strength. As Fig. 8 illustrates, an increase in aspect ratio may move an urban block or land use zone from a well-ventilated "wake interference flow" landscape to a "skimming flow" landscape. Urban temperatures will be further increased if urban consolidation is accompanied by reductions in greenspace; external water use and pervious surfaces. In cold climates, the resulting warmer temperatures will further reduce heating demand, but urban consolidation will have a real cost in warmer climates, unless other steps are taken to reduce the excess thermal stress. Improved architecture and construction of buildings in cities will enhance passive energy sources and further reduce energy consumption and anthropogenic heating. Reductions in QF potentially mean a reduction in the energy inputs and hence a potential reduction in warming, however these simple considerations ignore the importance of the site and height of the energy release (recall equation (2) and discussion). Future reductions in photochemical smog may arise from improved technology which will increase the receipt of shortwave radiation in those cities where photochemical smog is prevalent. Such increases are more likely, however, in cities such as Shanghai where coal burning is currently the major fuel source. Switching to alternative fuels, likely within the next 20 years, will significantly increase shortwave radiation receipts in such cities. In short, cities in developed nations without a heavy industrial base are unlikely to show any further significant warming at the continental-global scale. As explained in the Introduction, most urban growth will occur in cities in poorer nations, often located in the tropical latitudes. Already, most of the world's megacities are located within either the poor, or rapidly industrialising (e.g. Korea, Thailand), nations. Unlike most developed or industrialised nations, urbanisation is still increasing in many of the poorer nations. A United Nations report in 1989 found that in less developed nations, the urban population has been increasing at a rate of 3.6% per year and this was projected to continue until 2000. By way of comparison, urban populations in the developed world were increasing at or below 1% per year. Urban populations in Asia and Africa were projected to in507

Urban climates

crease by 2.3 billion between 1990 and 2025. To put this in perspective, this is equivalent to the total urban population for the globe in 1990, and represents 82% of the global increase in urban dwellers for the period 1990-2025. Recalling the findings of Karl et al. on the size of the urban warming signal for the US, it is of interest to note that this increase in urban population for Asia alone (1990-2025) is about ten times the increase in urban population in the US from pre-1900 to 1985. Increasing urbanisation in the African and Asian regions is not only concentrated into megacities. The same 1989 United Nations report found that the fastest growing cities (1985-2000) in the less developed nations (and for the globe) were projected to be those with populations less than 7.5 million (in 1985). Small towns will grow into medium-sized towns; middle-sized cities will grow to become megacities. If the urban contribution to the total warming observed in the US from 1901 to 1984 could be generalised (viz. 0.1--0.4~ then these large increases in urban populations and cities have the potential to make a significant contribution to the global temperature. Similarly, if the population-urban heat island relationships developed by KARL et al. (1988) are valid globally, then a large amount of the urban heating will result from this increase in population and size of medium-sized towns. JAUREGUI's (1986) graph of (log) population versus urban heat island strength shows a smaller slope for tropical, compared to mid-latitude, cities (recall Fig. 10). Furthermore KARL and JONES (1989) note that the urban warming apparent in the US record was not evident in the global temperature record. Clearly the observational evidence of global warming due to the urban growth during the 20th century is equivocal. It is thus very difficult to quantify the impact of the enormous increases in urban dwellers and city sizes projected into the 21st century. Regional and local scale climate changes will occur in those regions undergoing large-scale urbanisation. The IPCC (1991) predicted that with doubled CO2, tropical regions will show only small temperature increases but possibly increased rainfall. Urban effects will be superimposed on this backdrop of greenhouse-induced climate change. The projected increases in urban populations discussed above are almost beyond our comprehension, if they are correct then many areas in the Asian and African regions which will have a much different climate than exists today. Many of these changes will be similar to those already experienced by urban dwellers in industrialised and developed nations, and described above. Minimum and mean temperatures could be at least 5~ warmer, using the US data as an example. There will also be some differences: the amount of heating due to the urban heat island may be roughly similar in magnitude, but its strength will depend on the season. Changes in the timing and length of the wet and dry seasons in the tropics will therefore impact on the duration of urban-induced warming. Many cities in the poor and rapidly industrialising nations will have a markedly different radiation climate. Unless alternative fuel technologies are developed rapidly, these cities will experience reductions in the incoming solar radiation; UBL warming may result from radiation absorption or alternatively there will be a large increase in diffuse radiation. Improvements in fuel technology should mean that these cities pass through this "stage" of solar dimming (STANHILLand KALMA,1993) more rapidly than industrialised nations. If current patterns of urban growth in less developed nations prevail, much of the urban land-use will be informal squatter settlements. The effect of this altered land use depends, of course, on the pre-urban land-use. There is insufficient knowledge to quantify the change in energy balance partitioning that would arise from a transition from a

508

References grassland or a forested landscape into a squatter settlement. The results from OKE et al. (1992) would suggest that the convective sensible heat flux is the smallest heat flux in a tropical city energy balance. If so, possibly energy balances will not be altered greatly. The model simulations of SEAMAN et al. (1989) show that rainfall modifications are more likely to occur in those small to medium sized towns that will increase in population and extent. Although these results cannot be generalised, such modifications to rainfall patterns, storms and possible flooding will be of far greater significance. These local scale climate impacts are important, especially in terms of energy consumption, human health and their potential for causing larger scale climate change. Their importance, however, pales in comparison to the likelihood of increased disease and mortality arising from limited air and water quality plus hazards of increased flooding and landslides that will occur in the very poor cities. It can only be concluded that the combined effects of weather and climate changes, possible sea level rise (more than 20% of the world's cities are sited on coasts) and increased climate variability predicted to result from the enhanced greenhouse effect plus urban-induced local climate change and worsening air and water quality will severely limit the quality and sustainability of the urban environment in less developed and poor nations.

References ACKERMAN, B., 1985. Temporal march of the urban heat island. J. Clim. Appl. Meteorol., 24: 547554. ACKERMAN, B., 1987. Climatology of Chicago area urban-rural differences in humidity. J. Clim. Appl. Meteorol., 26: 427-430. ACKERMAN, B. and MANSELL, J. W., 1978. Thermodynamic structure of the PBL at midday. In: Summary of METROMEX. Volume 2: Causes of Precipitation Anomalies. Illinois State Water Survey, Urbana, Bulletin 63, pp. 129-164. ADEBAYO, Y. R., 1987. A note on the effect of urbanization on temperature in Ibadan. J. Climatol., 7: 185-192. ADEBAYO, Y. R., 1991. Day-time effects of urbanization on relative humidity and vapour pressure in a tropical city. Theor. Appl. Climatol., 43: 17-30. AIDA, M., 1982. Urban albedo as a function of the urban structure - a model experiment. BoundaryLayer Meteorol., 23: 405-413. ANGELL, J. K., HOECKER, W. H., DICKSON, C. R. and PACK, D. H., 1973. Urban influence on strong daytime air flow as determined from tetroon flights. J. Appl. Meteorol., 12: 924-936. ARNFIELD, J. A., 1982. An approach of the estimation of the surface radiative properties and radiation budgets of cities. Phys. Geog., 3: 97-122. ATKINSON, B. W., 1969. A further examination of the possible effect of London's urban area on the distribution of thunder rainfall. Trans. Inst. Br. Geog., 48:97-117. ATKINSON, B. W., 1971. The effect of an urban area on the precipitation from a moving thunderstorm. J. Appl. Meteorol., 10: 47-55. AUER, A. H., 1981. Urban boundary-layer, in METROMEX: a review and summary. Meteorol. Monogr., 18, No. 40: 41-62. AYERS, G. P., BIGG, E. K., TURVEY, D. E. and MANTON, M. J., 1982. Urban influence of condensation nuclei over a continent. Atmos. Environ., 16:951-954. BALLING JR., R. C. and BRAZEL, S. W., 1987a. Temporal variations in Tucson, Arizona summertime atmospheric moisture, temperature and weather stress levels. J. Clim. Appl. Meteorol., 26: 995999. BALLINGJR., R. C. and BRAZEL, S. W., 1987b. Recent changes in Phoenix, Arizona summertime diurnal precipitation patterns. Theor. Appl. Climatol., 38: 50-54.

509

Urban climates BORNSTEIN, R. D., 1986. Urban climate models: nature, limitations, and applications. Urban Climatology and its Applications with Special Regard to Tropical Areas, WMO No. 652, World Meteorol. Organization, Geneva, Switzerland, pp. 237-276. BORNSTEIN, R. O., 1989. Nature, limitations and applications of urban climate models. Proceedings of

the Workshop on Saving Energy and Reducing Atmospheric Pollution by Controlling Summer Heat Islands. Berkeley, CA, pp. 125-149. BORNSTEIN, R. D. and JOHNSTON, D. S., 1977. Urban-rural wind velocity differences. Atmos. Environ., 11: 597-604. BOWNE, N. E. and BALL, J. T., 1977. Observational comparison of rural and urban boundary-layer turbulence. J. Appl. Meteorol., 9: 862-873. BRAHAM JR., R. R., 1981. Urban precipitation processes. METROMEX: a review and summary. Meteorol. Monogr., 18, No. 40: 141-152. BRYANT, n. K., GLANCY, R. T. and AUER, A. n., 1978. METROMEX urban and rural wind velocity profiles. Proceedings of the Third U.S. National Council on Wind Engineering Research. Wind Engineering Research Council, Gainesville, 11-5-1. CARLSON, T. N., DODD, J. K., BENJAMIN, S. G. and COOPER, J. N., 1981. Satellite estimation of the surface energy balance, moisture availability and thermal inertia. J. Appl. Meteorol., 20: 67-87. CAYAN, D. R. and DOUGLAS, A. V., 1984. Urban influences on the surface temperatures in the southwestern United States during recent decades. J. Appl. Meteorol., 23: 1520-1530. CHANGNON JR., S. A., 1968. The La Porte anomaly - Fact of fiction? Bull. Am. Meteorol. Soc., 49: 165-168. CHANGNONJR., S. A., 1981. METROMEX: a review and summary. Meteorol. Monogr., 18:181 pp. CHANGNON JR., S. A., 1992. Inadvertent weather modification in urban areas: lessons for global climate change. Bull. Am. Meteorol. Soc., 73: 619-628. CHANGNON JR., S. A. and HUFF, F. A., 1986. The urban-related nocturnal rainfall anomaly at St. Louis. J. Appl. Meteorol., 25: 1985-1995. CHANGNON JR., S. A., HUFF, F. A., SHICKEDANZ, P. T. and VOGEL, J. L., 1977. In: Summary of METROMEX. Vol. 1, Weather Anomalies and Impacts, Bulletin 62. Illinois State Water Survey Division, Urbana, IL, 258 pp. CUING, J. K. S., CLARKE, J. F. and GODOWITCH, J. M., 1983. Modulation of heat flux to different scales of advection in an urban environment. Boundary-Layer Meteorol., 25:171-191. CHOW, S. D., 1992. The urban climate of Shanghai. Atmos. Environ., 26B: 9-15. CHOW, S. D. and CHANG, C. C., 1981. On some features of the Shanghai urban climate. Paper presented to IAMAP Symposium. Hamburg, Germany. CLARKE, J. F., CUING, J. K. S. and GODOWITCH, J. M., 1982. An experimental study of turbulence in an urban environment. NTIS P8 226085 ESRL. U.S. Env. Protect. Agency, Research Triangle Park, NC, 150 pp. CLEUGH, n. A. and OKE, T. R., 1986. Suburban-rural energy balance comparisons in summer for Vancouver, BC. Boundary-Layer Meteorol., 36:351-369. COPPIN, P. A., 1979. Turbulent fluxes over a uniform urban surface. Ph. D. Thesis, Flinders University, South Australia. DABBERDT, W. F. and DAVIS, P. A., 1978. Determination of energetic characteristics of urban-rural surfaces in the greater St. Louis area. Boundary-Layer Meteorol., 14: 105-121. ELSOM, D. M. and MEADEN, G. T., 1982. Suppression and dissipation of weak tornadoes in metropolitan areas: a case study of Greater London. Mon. Weather Rev., 110: 745-756. ESTOURNEL, C., VEHIL, R., GUEDALIA, O., FONTAN, J. and DRUILHET, A., 1983. Observations and modeling of downward radiative fluxes (solar and infrared) in urban/rural areas. J. Clim. Appl. Meteorol., 22: 134-142. FENG, J. Z. and PETZOLD, D. E., 1988. Temperature trends through urbanization in metropolitan Washington, DC, 1945-1979. Meteorol. Atmos. Phys., 38: 195-201. GALLO, K. P., MCNAB, A. L., KARL, T. R., BROWN, J. F., HOOD, J. J. and TARPLEY, J. D., 1993. The use of NOAA AVHRR data for assessment of the urban heat island effect. J. Appl. Meteorol., 32: 899-908. GODOWITCH, J. M., 1986. Characteristics of vertical turbulent velocities in the urban convective boundary layer. Boundary-Layer Meteorol., 35: 387-407. GOWARD, S. N., 1981. Thermal behaviour of urban landscapes and the urban heat island. Phys. Geog., 2: 19-33.

510

References GRIMMOND, C. S. B., 1988. An evaporation-interception model for urban areas. Ph. D. Thesis, The University of British Columbia, Vancouver, BC, 206 pp. GRIMMOND, C. S. B., 1992. The suburban energy balance: methodological considerations and results for a mid-latitude west coast city under winter and spring conditions. Int. J. Climat., 12:481-497. GRIMMOND, C. S. B. and OKE, T. R., 1986. Urban water balance II: results from a suburb of Vancouver, BC. Water Resour. Res., 22: 1404-1412. GRIMMOND, C. S. B. and OKE, T. R., 1990. Preliminary comparisons of measured summer suburban and rural energy balances for a hot dry city, Tucson, AZ. Proc. Int. Symp. Urban Climatol., Air Poll. and Planning in Trop. Cities. Guadalajara, Mexico.. GRIMMOND, C. S. B., OKE, T. R. and CLEUGH, H. A., 1993. The role of "rural" in comparisons of observed suburban and rural flux differences. Proceedings of IAHS Conference. Yokohama, Pub. No. 212, pp. 165-174. HAGE, K. D., 1975. Urban-rural humidity differences. J. Appl. Meteorol., 14: 1277-1283. HAMMOND, A. L. (Editor), 1992. World Resources (1992-1993). In: A Guide to the Global Environment. Oxford University Press, Oxford, UK. HARNACK, R. P. and LANDSBERG, H. E., 1975. Selected cases of convective precipitation caused by the Metropolitan area of Washington, DC. J. Appl. Meteorol., 14: 1050-1060. HAY, J. E., 1984. An assessment of the mesoscale variability of solar radiation at the Earth's surface. Sol. Energy, 32: 425-434. HESS, M. et al., 1978. Opracowanie bilansu radiacyjnego miasta Krak6wa i terenow pozamiejskich. lnstit. Geografii Uniwersytetu Jagiellon-skiego 90, Krak6w. HILDEBRAND, P. H. and ACKERMAN,B., 1984. Urban effects on the convective boundary-layer. J. Atmos. Sci., 41: 76-91. HJELMFELT, M. R., 1982. Numerical simulation of the effects of St. Louis on mesoscale boundarylayer airflow and vertical air motion: simulations of urban vs. non-urban effects. J. Appl. Meteorol., 21: 1239-1257. HOGSTROM, U., BERGSTROM,H. and ALEXANDERSSON,H., 1982. Turbulence characteristics in a near neutrally stratified atmosphere. Boundary-Layer Meteorol., 23: 449-472. HOWARD, L., 1833. The Climate of London, Vols. I-III. London. HUFF, F. A. and CHANGNON,S., 1973. Precipitation modification by major urban areas. Bull. Am. Meteorol. Soc., 70: 265-270. JAUREGUI, E., 1986. Tropical urban climates: review and assessment. In: T. R. OKE (Editor), Urban Climatol. and its Applic. with Special Regard to Trop. Areas. WMO No. 652, World Meteorol. Organization, Geneva, pp. 63-86. JAUREGUI, E., 1987. Urban heat island development in medium and large urban areas in Mexico. Erdkunde, 41:48-51. JAUREGUI, E., GODINEZ, L. and CRUZ, F., 1992. Aspects of heat-island development in Guadalajara, Mexico. Atmos. Environ., 26B: 391-396. JOHNSON, G. T., OKE, T. R., LYONS, T. J., STEYN, D. G., WATSON, I. D. and VOOGT, J. A., 1991. Simulation of surface urban heat islands under 'ideal' conditions at night. Part I: Theory and tests against field data. Boundary-Layer Meteorol., 56: 275-294. JONES, P. D., KELLY, P. M., GOODESS,C. M. and KARL, T., 1989. The effect of urban warming on the Northern Hemisphere temperature average. J. Climate, 2: 285-290. KALANDA, B. D., OKE, T. R. and SPrrrLEHOUSE, D. L., 1980. Suburban energy balance estimates for Vancouver, BC., using the Bowen ratio-energy balance approach. J. Appl. Meteorol., 19:791-802. KALMA, J. D., 1974. An advective boundary-layer model applied to Sydney, Australia. BoundaryLayer Meteorol., 6:351-361. KARL, T. R. and JONES, P. D., 1989. Urban bias in area-averaged surface air temperature trends. Bull. Am. Meteorol. Soc., 70: 265-270. KARL, T. R., DIAZ, H. F. and KUKLA, G., 1988. Urbanization: its detection and effect in the United States climate record. J. Climate, 1:1099-1123. KUKLA, G., GAVIN, J. and KARL, T. R., 1986. Urban warming. J. Clim. Appl. Meteorol., 25: 12651270. LANDsBERG, H., 1981. The Urban Climate. Academic Press, London, 275 pp. LOOSE, T. and BORNSTEIN, R. D., 1977. Observations of mesoscale effects on frontal movements through an urban area. Mon. Weather Rev., 105: 563-571.

511

Urban climates LOWRY, W. P., 1977. Empirical estimation of urban effects on climate: a problem analysis. J. Appl. Meteorol., 16: 129-135. LYONS, T. J. and FORGAN, B. W., 1975. Atmospheric attenuation of solar radiation at Adelaide. Q. J. R. Meteorol. Soc., 101: 1013-1017. MATZARrdS, A. and MAYER, H., 1991. The extreme heatwave in Athens in July 1987 from the point of view of human biometeorology. Atmos. Environ., 25B: 203-211. MERONEY, R. N., 1982. Turbulent diffusion near buildings. In: E. J. PLATE (Editor), Engineering Meteorology. Ch. 11, pp. 481-521. METHOD, T. and CARLSON, T. N., 1982. Radiative heating rates and some optical properties of the St. Louis aerosol, as inferred from aircraft measurements. Atmos. Environ., 16: 53-66. MILLS, G. M. and ARNFIELO, A. J., 1993. Simulation of the energy budget of an urban c a n y o n - I. model structure and sensitivity test. Atmos. Environ., 27B: 157-170. NASRALLAH, H. A., BRAZEL, A. J. and BALLINGJR., R. C., 1990. Analysis of the Kuwait City urban heat island. Int. J. Climatol., 10:401--405. NUNEZ, M. and Or~, T. R., 1977. The energy balance of an urban canyon. J. Appl. Meteorol., 16:1119. O i l , T. R., 1973. City size and the urban heat island. Atmos. Environ., 7: 768-779. OKE, T. R., 1974. Review of Urban Climatology 1968-73. World Meteorological Organization Technical Note 134, Geneva. O I l , T. R., 1976. The distinction between canopy and boundary-layer urban heat islands. Atmosphere, 14: 268-277. OKE, T. R., 1979a. Review of Urban Climatology 1973-76. World Meteorological Organization Technical Note 169, Geneva. Or,E, T. R., 1979b. Advectively-assisted evapotranspiration from irrigated urban vegetation. Boundary-Layer Meteorol., 17:163-173. O I l , T. R., 1980. Climatic impacts of urbanization. In: W. BACH, J. PANKRATH and J. WILLIAMS (Editors), Interactions of Energy and Climate. pp. 339-356. O I l , T. R., 1981. Canyon geometry and the nocturnal urban heat island: comparison of scale model and field observations. J. Climatol., 1: 237-254. OKE, T. R., 1982. The energetic basis of the urban heat island. Q. J. R. Meteorol. Soc., 108: 1-24. Or~, T. R., 1984. Methods in urban climatology. In: W. KIRCHHOFER, A. OHMURA and H. WANNER (Editors), Applied Climatology (Ziircher Geog. Schriften, 14). Eidgen6ssiche Technische Hochschule Geographische Institut, Zurich, pp. 19-29. OKE, T. R., 1986. Urban Climatology and its Applications with Special Regard to Tropical Areas, WMO No. 652. World Meteorological Organization, Geneva, pp. 63-86. O i l , T. R., 1987. Boundary Layer Climates, 2nd edition. Methuen, London, 448 pp. O ~ , T. R., 1988. The surface energy budgets of urban areas. Prog. Phys. Geog., 12:471-508. O I l , T. R. and CLEUGH, H. A., 1986. Urban heat storage derived as energy balance residuals. Boundary-Layer Meteorol., 39: 233-245. Or~, T. R. and EAST, C., 1971. The urban boundary-layer in Montreal. Boundary-Layer Meteorol., 1: 411-437. Or,E, T. R. and FUGGLE, R. F., 1972. Comparison of urban/rural counter and net radiation at night. Boundary-Layer Meteorol., 2: 290-309. O i l , T. R. and MCCAUGHEY, J. H., 1983. Suburban-rural energy balance comparisons for Vancouver, BC: an extreme case? Boundary-Layer Meteorol., 26: 337-354. OKE, T. R., TAESLER, R. and OLSSON, L. E., 1990. The tropical urban climate. Energy Buildings, 1516, pp. 67-73. Or~, T. R., JOHNSON,G. T., STEYN, D. G. and WATSON, I. D., 1991. Simulation of surface urban heat islands under 'ideal' conditions at night. Part 2: Diagnosis of causation. Boundary-Layer Meteorol., 56: 339-358. O I l , T. R., ZEUNEI~, G. and JAUREGUI, E., 1992. The surface energy balance in Mexico City. Atmos. Environ., 26B: 433-444. PASZYNSrd, J., 1972. Studies on the heat balance and on evaporation. Geog. Pol., 22. PETERSON, J. T. and STOFFEL, T. L., 1980. Analysis of urban-rural solar radiation data from St. Louis, Missouri. J. Appl. Meteorol., 19" 275-283. PETERSON, J. T., FLOWERS, E. C. and RUDISILL, J. H., 1978. Urban-rural solar radiation and atmospheric turbidity measurements in the Los Angeles basin. J. Appl. Meteorol., 17: 1595-1609.

512

References PRICE, J. C., 1979. Assessment of the urban heat island effect through the use of satellite data. Mon. Weather Rev., 107: 1554-1557. PROBALD, F., 1974. Air pollution and the urban climate of Budapest. In: M. P~scI and F. PROBALD, (Editors), Man and Environment. Akademiai Kiad6, Budapest, pp. 191-196. Ross, S. L. and Or,E, T. R., 1988. Tests of three urban energy balance models. Boundary-LayerMeteorol., 44: 73-96. ROTH, M., 1993. Turbulent transfer relationships over an urban surface. II: integral Statistics. Q. J. R. Meteorol. Soc., 119:1105-1120. ROTH, M. and Or,E, T. R., 1993. Turbulent transfer relationships over an urban surface. I: Spectral characteristics. Q. J. R. Meteorol. Soc., 119:1071-1104. ROTH, M., OKE, T. R. and STEYN, D. G., 1989a. Velocity and temperature spectra and cospectra in an unstable suburban atmosphere. Boundary-Layer Meteorol., 47: 309-320. ROTH, M., OKE, T. R. and EMERY, W. J., 1989b. Satellite-derived urban heat islands from three coastal cities and the utilization of such data in urban climatology. Int. J. Remote Sensing, 10:1699-1720. ROUSE, W. R., NOAD, D. and MCCUTCHEON, J., 1973. Radiation, temperature and atmospheric emissivities in a polluted urban atmosphere at Hamilton, Ontario. J. Appl. Meteorol., 12: 798-807. SANDERSON,M., KUMANAN, I., TANGUAY, T. and SCHERTZER,W., 1973. Three aspects of the urban climate of Detroit-Windsor. J. Appl. Meteorol., 12: 629-638. SCHMID, H. P., 1988. Spatial scales of sensible heat flux variability: representativeness of flux measurements and surface layer structure over suburban terrain. Ph.D. Thesis, University of British Columbia, 299 pp. SCHMID, H. P. and OKE, T. R., 1992. Scaling North American urban climates by lines, lanes and rows. In: D. G. JANELLE(Editor), Geographical Snapshots of North America. Guilford Press, New York, pp. 395-399. SCHMID, H. P., CLEUGH, H. A., GRIMMOND, C. S. B. and Or,E, T. R., 1991. Spatial variability of energy fluxes in suburban terrain. Boundary-Layer Meteorol., 54: 249-276. SEAMAN, N. L., LUDWIG, F. L., DONALL, E. G., WARNER, T. T. and BHRUMRALKAR,C. M., 1989. Numerical studies of urban planetary boundary-layer structure under realistic synoptic conditions. J. Appl. Meteorol., 28:760-781. SEKIHARA, K., 1973. Some aspects of radiation climate of Tokyo as observed by the eleven year measurement of spectral global radiation from 1961 to 1971. J. Meteorol. Soc. Jpn., 51:119132. SHEA, D. M. and AUER, A. H., 1978. Thermodynamic properties and aerosol patterns in the plume downwind of St. Louis. J. Appl. Meteorol., 17: 689-699. SHREFFLER, J. H., 1979. Heat island convergence in St. Louis during calm periods. J. Appl. Meteorol., 18: 1512-1520. STANHILL, G. and KALMA, J. D., 1993. Solar dimming and urban heating at Hung Kong: a test of a negative feedback hypothesis. Personal communication. STANHILL, G. and MORESHET, S., 1992. Global radiation climate changes in Israel. Clim. Change, 22: 121-138. STEYN, D. G., 1980. Turbulence, diffusion and the daytime mixed layer depth over a coastal city. Ph. D. Thesis, University of British Columbia, Vancouver, BC. STEYN, D. G., 1982. Turbulence in an unstable surface layer over suburban terrain. Boundary-Layer Meteorol., 22:183-191. STEYN, D. G. and Or,E, T. R., 1980. Effects of a small scrub fire on the surface radiation budget. Weather, 35: 212-215. SUCKLING, P. W., 1981. Nocturnal observations of incoming longwave radiation and the urban heat island for a small prairie city. Arch. Meteorol., Geophys. Bioklimatol., Ser. B, 29: 23-27. SWAID, H. and HOFFMAN, M. E., 1990. Thermal effects of artificial heat sources and shaded ground areas in the urban canopy layer. Energy Buildings, 15-16: 253-261. TAKAMURA, T., 1992. Spectral reflectance in an urban area: a case study for Tokyo. Boundary-Layer Meteorol., 59: 67-82. TAPPER, N. J., 1990. Urban influences on boundary layer temperature and humidity: results from Christchurch, New Zealand. Atmos. Environ., 24: 19-27. UNITED NATIONS, 1989. Prospects of World Urbanization, 1988. Population Studies, 112. UNITED NATIONS ENVIRONMENT PROGRAM, 1992. Urban Air Pollution in Megacities of the World. UNEP and WHO Report.

513

Urban climates VOOGT, J. and Or~, T. R., 1991. Validation of an urban canyon radiation model for noctumal longwave fluxes. Boundary-Layer Meteorol., 54:347-361. VUKOVICH, F. M., 1983. An analysis of the ground temperature and reflectivity pattern about St. Louis, Missouri, using HCMM satellite data. J. Appl. Meteorol., 22:560-571. WANG, C. and LIU, J., 1982. The climate of the city of Hangzhou. Acta Geog. Sinica, 37: 164-173. WHITE, J. M., EATON, F. D. and AUERJR., A. H., 1978. The net radiation budget of the St. Louis metropolitan area. J. Appl. Meteorol., 17: 593-99. YAMASH1TA, S., 1979. On the attenuation of solar radiation in Tokyo. Bull. Tokyo Gakuei Univ., Ser. III, 30:119-130. YAMASHITA, S. and SEKINE, K., 1990/91. Some studies of the Earth's conditions relating to the urban heat island. Energy Buildings, 15-16: 279-288. YAK D., 1973. Sensible heat fluxes in and near Vancouver, BC. Ph.D. Thesis, University of British Columbia, 177 pp. YERSEL, M. and GOBLE, R., 1986. Roughness effects on urban turbulence parameters. Boundary-Layer Meteorol., 37: 271-284. YOSHIDA, A., 1991. Two-dimensional numerical simulation of thermal structure of urban polluted atmosphere effects of aerosol characteristics. Atmos. Environ., 25B" 17-23. YOSHIDA, A., TOMINAGA, K. and WATATANI, S., 1991. Field measurements on energy balance of an urban canyon in the summer season. Energy Buildings, 15-16:417-423. YOSHIKADO, H., 1992. Numerical study of the daytime urban effect and its interaction with the sea breeze. J. Appl. Meteorol., 31:1146-1164.

514

IV. Geophysiology and Future Climates

This Page Intentionally Left Blank

Chapter 14

Future climate surprises TSUNG-HUNG PENG

Introduction

The release of CO 2 and other greenhouse gases to the atmosphere since the industrial revolution two centuries ago has caused a significant increase in the concentration of these gases in the atmosphere. Continuous monitoring of atmospheric CO2 at Mauna Loa Observatory, Hawaii (KEELING et al., 1989) indicates that the concentration of CO2 gas has increased by about 26% over the pre-industrial level. Prediction of the consequences of the build-up of these gases is a major subject of study and debate among scientists. In general, a global warming is expected if the trend of build-up continues. However, because of our lack of basic knowledge regarding the operation of the Earth's atmosphere-hydrosphere-biospherecryosphere system, it is difficult to make a reasonable and useful, detailed prediction. Usually, a large range of uncertainty with respect to possible greenhouse effects is associated with these predictions. In addition, these sophisticated climate model simulations all point to a gradual warming in response to a gradual build-up of greenhouse gases. Can we trust these simulations and feel comfortable with the possibility of coping with the gradual climate change in the next 100 years? We need only to examine how well the climate models incorporate the key interactions of the Earth system into the model structure to find out that, as yet, we do not have these real interactions figured into models, making any predictions questionable. Indeed, the future climate change is uncertain. As pointed out by BROECKER (1987), there may be unpleasant surprises in the greenhouse. Nature herself has conducted large-scale climatic experiments in the past. The climate in the past million years has been characterized by a series of cyclic glaciations (see Chapter 2 by BERGER). The response of the Earth system to these natural experiments is recorded in ocean sediments, peat bogs and polar ice. The results of analyzing the most recent record over the past 100,000 years indicate that the Earth's climate does not respond to forcing in a smooth and gradual way (small change over thousands of years). Instead, it responds with sharp changes (the response time usually is less than a century). For example, results from central Greenland ice cores show large and abrupt climate changes during the late stages of the last glaciation, suggesting that climate is able to reorganize itself rapidly, perhaps within a few decades (DANsGAARD et al., 1993). The warming at the end of the last glaciation was interrupted with a few abrupt returns to glacial climate. The best example is the Younger Dryas event. Based on oxygen isotope data from Greenland ice cores (DANsGAARD et al., 1989; JOHNSEN et al., 1992), the Younger Dryas ended abruptly over a period of about 50 years (ALLEY et al., 1993). If this is correct, the lesson we have learned from the past is that the main responses of the Earth system to greenhouse gas build-up could come in jumps

517

Future climate surprises

whose timing and magnitude are unpredictable. A study by BROECKER and DENTON (1990) indicates that massive reorganizations of the ocean-atmosphere system are key events that link cyclic changes in the Earth's orbit to the advance and retreat of ice sheets. What are the relationships between climate change and the mode of operation for ocean circulation systems? What sorts of surprises might be in store for us if the greenhouse warming is effective in changing the mode of operation of the ocean-atmosphere system? In this chapter, we review lessons learned from the study of climatic changes in the past, the corresponding large-scale reorganization of the mode of operation for the ocean-atmosphere system, and the possible linkage between climate change and ocean circulation. Understanding the mechanism of global ocean circulation is crucial in seeking such linkages. The great ocean conveyor circulation of BROECKER (1991) is reviewed for this purpose. Based on analysis of air bubbles in polar ice cores, the CO2 content in the atmosphere varied significantly with climate changes over the last 100,000 years. We summarize in this chapter the history of atmospheric CO2 variations and review possible causes of atmospheric CO2 variations in relation to the last glacial-to-interglacial climate change.

Changes in climate and ocean circulation in the past

The cyclic climatic changes between warm and cold periods are recorded by oxygen isotope (~180) in foraminifera shells deposited in the deep sea sediments. Fig. 1 shows the marine 6180 record obtained from the benthic foraminifera from a deep sea core raised from the east equatorial Pacific (SHACKLETONet al., 1983). Sharp changes in dl80 128,000 and 12,000 years ago denote the end of the glacial period (Terminations II and I). Slower

S ~80 (%.)

O~ 12,ooo I

(.~ '-

,

,

,

.--

TERM I

r) .~

r"

75,000 o _

(3" 0

t./1

5

,(-.

TERM ~E

128'0001

mi-i~'~-

-

-

(yrs BP)

_ I

I

!

Fig. 1. Oxygen isotope record of deep sea core V19-30 from equatorial East Pacific (SHAcKLETONet al., 1983). The marine isotope stages of Emiliani and the terminations are shown. The generally accepted ages for the major climate changes are also given.

518

Changes in climate and ocean circulation in the past

MARl NE STAGES 3 I I so 15bls Isdlsel

2 I

+,.o

,,,

6

+0.5 ~-~ 0.0 -0.5 -I.0

!

0

!

5O

!

I

IO0

150

+0.5~

-3 .