B. 8 Global Warming Potential (GWP) The Global Warming Potential is an attempt to provide a simple measulre of the relative radiative effects of the emissions of various greenhouse gases. The index is defined as the cumulative radiative forcing between the present and some chosen tinie horizon caused by a unit mass of gas emitted now. esprcssed relative to thal: for some reference gas (hcre CO, is used). The future global warming commitment of a greenhouse gas over a chosen tihe horizon can be estimated by multiplying the appropriate GWP by the amount of gas emitted. For. example. GWPs could be used to compare the effects, of reductions in C02 emissions relative to reductions in methane emissions, for a specified time horizon. Derivation of GWPs requires knoyledge of the fate of the emitted gas and the radiative forcing due to the amount remaining in the atmosphere. Although the G'IVPs are quoted as single values, the typical uncertainty is +35%, not including the uncertainty in the carbon dioxide reference. Because CWPs are based on the radiative forcing concept, they are difficult t o a p p l y t o r a d i a t i v e l y i m p o r t a n t constituents that are unevenly distributed in the atmosphere. No attempt is made to define a GWP for aerosols. Additionally the (choice of time horizon !vi11 depend on policy considerations. G'IVPs need to take account of any indirect effects of the emitted greenhouse gas if they are to reflect correctly future warming potential. The net G\\ 'Ps for the ozone depleting gases. which include the representation of the carbon cycle. Revised lifetimes for gases destroyed by chemical reactions in the lower atmosphere (particularly methane, HCFCs and HFCs ) have resulted in GWPs that are slightly lower (typically by 10 15%) than those cited in IPCC (1994). The IPCC definition of GN'P is based on calculating the relative radiative impact of a release  of a trace gas over a time horizon in a constant background atmosphere. In a future atmosphere with larger CO, concentrations, such as occur in all of the IPCC emission scenarios (see Figure 5b). we would calculate slightly larger CWP values than those given in Table 4. '

9.9 Emissions and concentrations of greenhouse gases and aerosols in the future 5.9. I The 1592 emission scenarios The projection of future anthropogenic climate c h a n g e d e p e n d s , a m o n g o t h e r t h i n g s ,' o n  assumptions made ,about future emissions of greenhouse gases and aerosol precursors and the proportion of e m i s s i o n s r e m a i n i n g i n t h e atmosphere. Here we consider the IS92 emission scenarios (IS92a to 0 which were first discussed in IPCC (1992). The IS92 emission scenarios estend to the year 2100 and include emissions of CO,. CH,, NzO, the halocarbons (CFCs and their substitute HCFCs and HFCs). precursors of tropospllerii ozone and sulphate aerosols and aerosols from biomass burning. A wide range of assumptions regarding future economic, demographic and policy factors are encompassed (IPCC. 1992). In this report. the emissions of chlorine and bromine containing 25 ..

.. i Table 4: Global Marming Potential referenced to the updated decay response for the Bern carbon cycle model ancljicture CO, atmospheric concentrations held constant at current levels.

S p e c i e s C h e m i c a l GIobal ll'arming Potential Formula (Time Horizon) (years) 20 years 100 years 500 years 1 1 variables co2 co2 1 hlethane*

Nitrous oxide CH4 N20 120 280 1223 56 21 6.5 3 1 0 1 7 0 .

halocarbons listed in IS92 a r e a s s u m e d t o b e phased out under the Montreal Protocol and its Adjustments and Amendments and so a single * revised future emission scenario for these gases is incorporated in all of the IS92 scenarios. ,

Emissjons of individual HFCs are based on the original IS92 scenarios, although they do not reflect current markets. C02 emissions for the six scenarios are shown in Figure 5a. The calculation of future concentrations of greenhouse gases, given certain emissions, entails modelling the processes that transform and remove the different gases from the atmosphere. For example. future concentrations of COZ are calculated using models of the carbon cycle which model the exchanges of COz b e t w e e n . t h e a t m o s p h e r e a n d t h e o c e a n s a n d t e r r e s t r i a l biosphere (see Section B. 1); atmospheric chemistry models are used to simulate the removal of chemically active gases such as methane. All the IS92 emission scenarios, even IS92c. imply increases in greenhouse gas concentrations from 1990 to 2100 (e. g., C02 increases range from 35 to 170% (Figure 5b); CH, from 22 to 175%; and NzO from 26 to 40%).

dependent on 'the concentration of the gas and the. strength with which it absorbs and re emits long wave radiation. For sulphate aerosol, the direct and indirect radiative forcings were calculated on the basis of sulphur emissions contained in the IS92 scenarios. The radiative forcing due to aerosol from biomass burning was assumed to remain constant a t 0.2 iVm 2 after 1990. The contribution from aerosols is probably the most uncertain part of future radiative forcing. Figure 6a shows a single "best estimate" of historical radiative forcing from 1765 to 1990 (including the effects of aerosols), followed by radiative forcing for Scenarios ,IS92 a to f. Figures 6b and c shoiv the contribution to future radiative forcing from various components of the 'IS92a Scenario; the largest contribution comes from COz, with a radiative forcing of almost 3 6 Wm 2 by 2100. The negative forcing due to tropospheric aerosols, in a globally averaged sense, offsets some of the greenhouse gas positive forcing. Hotvever, becaust. tropospheric aerosols are highly variable regionally, their globally averaged radiative forcing will not adequately describe their possible climatic impact. Future projections of temperature and sea level b a s e d o n t h e I S 9 2 e m i s s i o n s s c e n a r i o s a r e discussed in Section F. .For greenhouse gases, radiative forcing is ,

..

I .

. IS92f IS92a IS92b IS92d IS92C . .. .  :2000 ,2020 2040 2060 2080 2100 ..  .. b ' .. Year ..." ' , ., ,

IS92e IS92f IS92a IS92b IS92d IS92c .*.3 0 0 [ ' ' ' ' ' ' I ' ' ' 1 . . . ,2000 : 2020 2040 2C60 2080 2100 .. .. . ' . .. .. . . . ,.

, .. . Year [Figure 5: (a) Total anthropogenic CO, emissions under the IS92 emission scenarios and (6) the resulting atmospheric ,C02 concentrations 'calculated using the 'Bern'carbon 'cycle model and the carbon budget for the 1980s shown in Table 2. . ... .

.. ... . . .. .. . .. ... . .. .

:;%.:"  :'. ..  . . . .

27 . .. 8 Figure 6: (a) Total'globally and annually averaged historical radiative forcing from 1765 to 1990 due to changes in greenhouse gas concentrations and tropospheric aerosol emissions and projected radiative forcing ualues to 2100 derived from the IS92 emissions scenarios. (b) Radiative forcing components resulting from the IS92a emission scenario for 1990 to 2100. The 'Total non C02 trace gases' curue includes the radiatiue forcing from methane (including methane related increases in stratospheric water uapour). nitrous oxide. tropospheric ozone and the halocarbons (including the negative forcing effect of stratospheric ozone depletion). Halocarbon emissions have been modged to take account of the Montreal Protocol nd its Adjustments and Amendments. The t d e aerosol components are: direct sulphate, indirect sulphate and direct biomass burning. (c) Non C02 trace gas radiative forcing components. 'WBr direct' is the direct radiative forcing resulting from the chlorine and bromine containing halocarbons: emissions are assumed to be controlled under the Montreal Protocol and its Adjustments and Amendments. The indirect forcing from these comp. otrnds (through stratospheric ozone depletion) is shown separately (Stra;. 03. All other emissions follow the IS9Za Scenario. The tropospheric ozone forcing (Trop. 63 takes account of concentration changes due only to the indirect effect of methane.

2s . .

.. . . .. Year " "" "" I""

__" "" " """ " " " " "  Sulphate aerosol  indirect 2 2000 2020 2,040 20602080. 2100 Year l .2 ,b l l l l l l l I I Year .. ..

.. 8.9.2 Stabilisation of greenhouse gas and 1000 4 aerosol concentrations 950 (a) ' e.  ' SlijOO i / /.

* An important question to consider is: how might 900 greenhouse gas concentrations be stabilised in the 850 ; future? g 800 I f global C02 emissions were maintained at near ' g 750 ; current (1994) levels, they would lead to a nearly 3 700 ; c o n s t a n t r a t e of i n c r e a s e i n a t m o s p h e r i c 5 650 : S 650 .

concentrations for at least two centuries. reaching 2 6oo : about 500 ppmv (approaching twice the pre ' 5j0 1 industrial concentration of 280 ppmv) by the end of 8 500 t h e S 450.

In IPCC (19941, carbon cycle models were used to calculate the emissions of C02 which would lead to stabilisation of different concentration Year levels from 350 to 750 ppmv. The assumed concentration profiles leading to stabilisation are shown in Figure i a (excluding 350 ppmv). Many different stabilisation levels, .time scales for achieving these levels. and routes to stabilisation could have b'een chosen. The choices made are not intended to have policy implications: the exercise is illustrative of the relationship between CO? emissions and concentrations. Those in Figure 7a ' assume a smooth transition from the current average ratc of CO, c o n c e n t r a t i o n i n c r e a s e . t9, ' . , C02 emissions stabilisation. To a first approsimntion. the stabiliscd concentration level d e p e n d s m o r e u p o n t h e accumulated amount of carbon emitted up to the tirnc of s t a b i l i s a t i o n , t h a n u p o n t h e e x a c t c o n c e n t r a t i o n p a t h followed en route t o s t a b i l i s a t i o n . Year  / /

/ /

 /

v /

.

h: ew results have bcen produced to take account of the revised. carbon budget for the 1980s. (Table 21, b, ut the main conclusion, that stabilisation of concentration requires emissions eventually to drop well below current levels, remains unchanged from IPCC (1994) (Figure ibl. Because the new budgct implies a reduced terrestrial sink. the allowable emissions to achieve stabilisation are up to 10% lower than those in IPCC (2994). In addition, these calculations have been estended to include alternative pathways towards stabilisation (Figure 7a) and a higher stabilisation level (1000 ppmv). The alternative pathways assume higher emissions in the early years, but require steeper reductions in en1ission. i i i 1 latcr years (Figure 7bl. The 1000 ppxn; stabilisation case allows higher maximum emissions, but still requires a decline to. current levels by about 2 10 years from now and further reductions thereafter (Figure 7b). Figure' 7: (a) CO, concentration profiles leading to stlnbilisation at 450. 550. 650 and 750 ppmc following the pathways defined in IPCC (1994) (solid curtes) and for . pathways that allow emissions to follow IS92a until at least 2000 (dashed curves). A single profile that stabilises at a CO, concentration of 1000 ppmv and follow IS92a emissions until at'least 2000 has also been defined. (b) C02 emissions leading to stabilisation at concentrations of 450. 550, 650, 750 and 1000 ppmv following the profiles shown in (a). Current anthropogenic CO, emissions and those for IS92a are shown for comparison. The calculations use the Bern' carbon cycle model and the carbon budget for the 7980s shown in Table 2.

.. .

29 halocarbons listed in IS92 are assumed t o b e phased out under the Montreal Protocol and its 4 Adjustments and "hendrnents and so a single revised future emission scenario for these gases is incorporated in all ot'the IS92 scenarios. Emissiors of individual HFCs are based on the original 1592 scenarios, although they do not reflect current merkets. CO, emissions for the six scenarios a r e shoivn in Figure 5a. The calculation of future calncentrations of greenhousd ga. ses. given certain emissions, entails m o d e h g the processes that transrorm and remove the d. ifferent gases from the atmosphere. For example, future concentrations of COP a r e calculared using models of the carbon cycle which model the exchanges of CO? b e t v i e e n t h e a t m o s p h e r e a n d t h e o c e a n s a n d t e r r e s ,t r i a l biosphere (see Section B. 1); atmospheric chemistry m o d e l s a r e u s e d to simulate the remova. 1 of chemically active gases such as methane. All the IS92 emission scenarios. even IS92c. imply increases in greenhouse gas concentrations from 1990 to 3100 (e g . C02 increases range from 35 to 170% (Figure 5b); CII, from 22 to 175%; and N20 from 26 to 10%).

dependent on the Concentration of the gas and tiii strength ivith which it absorbs and re emits long wave radiation.. For sulphate aerosol. the direct and indirect radiative forcings were c; alculated on th. e basis of sulphur emissions contained in ithe IS92 scenarios. The radiative forcing due to aerosol from biomass burning was assumed to remain constant a t 0.2 Zvmn* after 1990. The cantribulion from aerosois is probably the most uncertain part of future radiatiwe forcing. Figure Ga shows a single "best estimate" of historical radiative forcing from 1765 to 1990 (including ,the effects of aerosols). folllavcved by radiative forcing for Scenarios IS912 a to 1. Figures 6b anld c show the Contribution to future radiative forcing from variow components of. the IS92a Scenario; the largest contribution comes from C02, with a radiative forcing of almost +15 Wm 2 !by 2100. The negative forcing due to tropospheric aerosols, ir. R z1? 5?.!! y everF. yd serse. n f f w t y snmp of t h ~ greenhouse gas positive forcing. However, because tropospheric aerosols are highly veriable regionally, their glo5ally averaged radiative forcing will not adequetely describe their possible climatic impact. Future projections of temperature and sea level based on the IS92 emissions scenarios are discussed in Section F. For greenhouse gases. radiative forcing is ' 4 0 ,,,,,,,,,,

lS92t IS92a IS92b IS92d IS92C .. . 2000 , 2020 2040 2060 2080 2100 t ': .. d. Year .. . .. . . . .. . . IS92e IS92f .. 2000 2020 2040 2060 2080 2100 .. .. . . .. Year Figure 5: (a] Taltal anthropogenic CO, emissions under the 15912 emission scenarios and (a) the resulting atmospheric GO2 concentrlutio ns calculated using the 'Bern' carbon ' cycle model hnd the carbon budget for the 1980s shown in Table 2. .. . . . ,. _   3 ." x,,.:: . . .. .^ . .. . .. B. 6 Tropospheric and stratospheric aerosols  Aerosol is a term used for particles and very Small droplets of natural and human origin that occur in the atmosphere; they include dust and other particlcs which can be made up of many different chemicals. Aerosols are produced by a variety of processes. both natural (including dust storms and volcanic activity) and anthropogenic (including fossil fuel and biomass burning). Aerosols contribute to visible haze and can cause a diminution of the intensity of sunlight at the ground. .4erosols in the atmosphere influence the radiation balance of the Earth in two ways: (i) by scattering and absorbing radiation  the direct e p c t , and'( ii1 by modifying the optical properties, amount and lifetime of clouds  the irdirrct effect. .Uthough some aerosols.. such as soot, tend to warm the surface. the net clinmtic effect of anthropogetlic aerosols is belicved to be a negative radiativc forcing. tending to cool the surface (see Sectidn B. 7 and Figure 21. Most aerosols tvith anthropogenic sources are found in the lower troposphcrc (below 2 km). A e r o s o l s i n d e r g o c h e r n i c a l a n d p h y s i c a l transformations in the atmosphcre: cspecially within clouds. and'are removed largriy by prccipitation. Consoqucntly aerosols in the lower troposphcrc typically have rcsidcncc times of a few days. Bccausc of thcir short lifetime. aerosols in thc lowdr troposphcrc are distributrd inhomogcncously ivith maxima closc to the natural (especiaiiy desert) and anthropogenic (cspccinlly industrial and biomass combustion) source regions. Aerosol particlcs resulting from volcanic activity .can reach thc . stratosphrrc where they are transportrd around thc globc over many months or years. The radiative forcing due to aerosols depends on the size. shapc and chcmical composition of thc particles and the spatial distribution of the aerosol. LVhile thcsc factors are comparatively well knoun for ' stratospheric aerosols, there remain many uncertainties concerning trop& pIleric aerosols. Since IPCC (1994). there have been scveral a d v a n c c s i n u n d e r s t a n d i n g t h e i m p a c t of 'tropospheric aerosols on clim5te. These include: li) new calculations of the spatial distribution of sulphnx neroso! Inr221y resulting from fossil fud combustion and (ii) the first calculation of the spatial distributiop of soot aerosol. The impact of thcse developmcnts on the calculation of aerosol radiative forcing is discussed in Section f3.7.

..

'4 .. .. .. . ..  .. , B. 7 Summary of radiative forcing Globally averaged radiative forcing is a useful concept for giving a first order estimate of the potential climatic importance pf various forcing mechanisms. However. as was emphasised in IPCC (1994), there are limits to its utility. In particular. :he spatial patterns of forcing differ betiyeen the globally tvell mixed greenhouse gases, the regionally \wying tropospheric ozone, and the even more regionally concentrated tropospheric aerosols. and so a comparison of the global mean radiative forcings docs not give a complete picture of their possible elinlatic impact. Estimates of the radiative forcings due to changes in greenhouse gas concentrations since pre industrial times remain unchanged from IPCC. (19341 (see Figure 2). These are +2.45 W m 2 (range: t2.1 to +2.8 iVm 2) for the direct effect of the main \veli mixed greenhouse gases (C02. CH+ N20 and the halocarbonsj. +0.4 Ct'm 2 (range: 0.2 to 0.6 V h 2 ) for tropospheric ozone and 0.1 iVm 2 (range: 0.03 to .

0.2 Wm 2) for stratospheric ozone. The total direct forcing due to anthropogenic aerosol (sulphatcs, fossil fuel soot and organic aerosols From biomass burning) is estimated to be 0.5 iVm 2 (range: 0.25 to 1.0 iVm 2). This estimate is smallcr than that givenjn IPCC (1994) owing to a reasscssnlent of the model results used to derive the geographic distribution of aerosol particles and the inclusion of anthropogenic soot aerosol for the first time. Thc direct forcing due to sulphate aerosols rcsulting from fossil fuel emissions and smelting is estimated 10 be 0.4 CVm 2 (range: 0.2 to 0.8 tvnr'). The first estimates 0. f the impact of soot in aerosols from fossil fuel sources have been made: significant uncertainty rcmains but an estimate of +0.1 C t ' m ~ (range: 0.03 to 0.3 iVm 2) is made. The direct radiative forc'ing since 1850 of particles associated with ,biomass burning is estimated to be 0.2 LVm 2 (range: 0.07 to 0. G \. 'dm 21. unchanged from IPCC (1994). I t has recently' been suggested that a significant fraction of the tropospheric dust aerosol is influeked by human activities but the radiative forcing of this component has not yet been quantified. Tho rerlyr nf v +t T r t w fnr thn rac! i?.! i. r fc:~!::: due to changes in cloud properties caused b). aerusois arising from human acrivitj (tho indirccr effect) is unchanged from IPCC (1994) at between 0 and 1.5 !Ym?. Several new studies confirm that, the indirect effect of aerosol tnai have caused 2 substantial negative radiative forcing since pre

..

.. . .. .. .. .

' , .. .. .. . . .. .. ," ~ I direct 'warming" and indirect "cdoling" effects. have now been estimated. In IPCC (1994), only the direct GLYPs were presented for these gases. The indirect effect reduces their GLVPs, but each ozone dep! eting gas must be considered indiridually. The net G\ VPs of the chlorofluorocarbons (CFCs) tend to be positive. while those of the halons tend to be negative. The calculation of indirect effects for a number of other gases (e. g.. NO,. CO) is not currpntly possible because of inadequate characterisation of many of the atmospheric processes involved. Updates or new CWPs are given for a number of key species (Table 4). based on improved or new estimates of atmospheric lifetimes, molecular ~a d i a t i v e ' f o r c i n g f a c t o r s , a n d i m p r o v e d representation of the carb'on cycle. Revised lifetimes for gases destroyed by chemical reactions in the lon. er atmosphere (particularly methane, HCFCs and HFCs ) have resulted in GWPs that are slightly lolver (typically by 10 15%/ 0) than those cited in IPCC (1994). The' IPCC definition of GWP is based on calculating the relative radiative impact of a release of a trace gas over a time horizon in a constant background atmosphere. In a future atmosphere with larger COz concentrations. such as occur in all .

of the IPCC emission scenarios (see Figure 5b), we \vould calculate slightly larger GWP values than ,

those given in Table 4.

o! J Emissions and concentrations of reenhouse gases and aerosols in the future B. 9. I The IS92 emissr" on. scenarios . '

The projection of, future anthropogenic climate change depends. among other things. on assumptions made about future emissions of greenhouse gases and aerosol precursors and the proportion of e m i s s i o n s r e m a i n i n g i n t h e atmosphere. Here we consider the IS92 emission scenarios'( IS92a to 0 which were first discussed in

IPCC (1992). The IS92 emission scenarios extend to the year 2100 and include emissions o f C02, CH,, N, O, the halorarhons KFCs and their sz~ hstitute HCFCs and HFCsl. precursors of tropo, spheric ozone and a. 2> lLL.: c aerosols and aerosols fro:, biomass burning. A wide range of assumptions regaiding dutijre economic, demographic and policy factors ;! re encompassed (IPCC. 1992). In this report, the emissions of chlorine and !bromine containing 23 .. .. ,. , .. .. . .. .. > .. .. ,"

Global average methane concentrations increased by 6% over the decade starting in 1984 (Figure 3). Its concentration in 199 1 was about 1720 ppbv,

 14504 greater than the pre industrial concentration of 700 ppbv (Table 1, Figure 3). Over the last 20 years, there has been a decline in tlie methane gronrh rate: in the late 1970s the concentration was increasing by about 20 ppbvlyr, during the 1980s the groivth rate dropped to 9 13 ppb::/ yr. Around the middle of 1992, merhane concentrations briefly stopped growing. but since 1993 the global growth rate has returned to about 8 ppbvlyr. I n d i v i d u a l m e t h a n e s o u r c e s a r e n o t w e l l quantified. Carbon isotope measurements indicate t h a t a b o u t 20% of the total annual methane emissions are related to the production and use of foskil fuel. In total. anthropogdic activities are responsible for about 6040% of current methane emissions (Table 3). Nethane emissions from natural wetlands appear to contribute about 20% to the gtobal methane emissions to the atmosphere. Such emissions will probably increase with global warming as a result of greater microbial activity. In 1932 the direct radiative forcing due to the increase in methane concentration since pre industrial times was about +0.47 Wm (Figure 2). Changes in the concentratim of methane have clearly identified che~ nical feedbacks. The main removal process for methane is reaction with the hydroxyl radical (OIH). .Addition of methane to the atmosphere reduces the . Concentration of tropospheric Oll which can in turn feed back and roduco the rate of methane removal. The adjustment time for a pulse of methane added ., to the atmosphere has been revised to 12 (+ 3) years (compared with 14.5 (c2.5) years in IPCC (1991)).

. Two factors are responsible fof the change: (a) a new estimate for the chemical removal rate (11% faster); and (b) inclusion of the uptake of methane by soils. The revised global sink strength is 560 (* loo) Tg( CH$/ year, higher than the 1991 estimate, but still consistent with the previous range of global . source strength.

B. 3 Nitrous oxide (N, O) There are many small sources of nitrous oxide, ' . both nntrlra! and zc:!; r2p;;::::;, ;; i1iiil arc jiiiicult to quantify. The main' anthropogenic. so. ilrces are fruit1 agricuirure and a number of industrial processes' (e. g.. adipic acid and nitric acid productionl. '4 best estimate of the current (1980s) anthropogenic emission of nitrous oxide is 3 to 8 /.

... .. . ,. 22 i .I .. .. ..

. .I ..  ..

". . . . . _.. .

, I: . 1. . , .. I .. Tg( N)/ yr. Natural sources are poorly q 1 .1 ,~r r i f i p c l ~ bur are probably twice as large as allthropoge; lic sources. Nitrous oxide is removed n; alnl!, by photolysis (breakdowx by sunlight) in the stratqsphere and conseyuer. d~ has a ion; lifetirr:? (about 120 years). Although sources cannot be well cjr. antiEet\, atmospheric measurements and evidencc from ice cores shoiv that the atmospheric abunciance of nitrous oside has incroased since the pre inciustrial era, most likely otving to human activities. In 199: armospheric levels of nitrous oxide were abour 312 ppbv: pre industrial levcls \rere about 275 ppbv (Table 1). The 1993 grow. th rate Eapprosimarely 0.5 ppbvlyr) was lower than that observed in the late 1980s and early 1990s (approximately 0.5 pphI,/ yr). but these short term changes i n gronth rate are within the range of variabilig. seen on decadal time scates. The radiatir. e forcing due t o thc chnnge in nitrous oxide since pre industrial times is about 1 0.14 1Vm. Z (Figure 2).

B. 4 Halocarbons and other

halogenated campounds Halocarbons are carbon compounds conraining fluorine. chlorine. bromine or iodine.. Many of these are efrective greenhouse gases. For: most of these cdrnpounds. human activities am the sole source. Halocarbons that contain chlarinc (CFCs ancl HCFCs) and bromine (hnlonsd I:: BLIS@ ozone depletion, and thcir e~ nissions are contrullcci under the Montreal Protocol and its Adjustments and Amendrncnts. As a result. growth! ratcs i n the conccncraeirjns of many of these compounds have already fallen (Fig& 4) and tPlc radiative impact of these compounds w i l l slowly decline oyer the nest century. The contribution to dirccl radinti1. e forcing due to Concentration increases of these CFCs and HCFCs since pre industrial timcs is itbout +0.25 \l'm?. Halocarbons can also exert an indirect negative radiative forcing through their depletion of stratospheric ozone (see Section R. 5.2). Perfluorocarbons (PFCs. e. g.. CF,, C2F6) and sulphur hexafluoride iSF6) are removed very slo\ vly from the atmosphere \\ lth estimated lifetimes greater than 1000 years. As a result, dh hv> l!: ?1' P Y !:::: :: accumulate in the atmosphere and will continue to influrnre rlirnntc fo~ r thsis: tdz uljears. Although the radiative forcing due to concentration increases of these compounds since pre industrial times is sn? all (about 1.0.01 IVtn. '). it may b e r m e significant in the future if concentrations continste to increase. HydrofZuororarbons IIHFCS) ;! Ye being used ?I )

replace ozone deplehg substances in some applications; rheir concentI ations and radiative impacts are currently small I f emiss: ons increase as envisaged in Scenario IS92a, the! 5r. ould contributc about 3% of the tora. 1 radiati\, e forcing from all greenhouse ga'es bq. the year 2 11.) 5.

9.5 Ozone (0,) Ozone is an .;:: portan: grccnho,~ l; e gas pslescnt i! l both the stratosphere and troposphere. Changes in ozone ca:: se rndintit 2 forciw: b) influencing both solar anti terrestrial radiation. The net radiative forcing is strongi) dependent on the vertical distriburicn of ozone chang  and is panticularli. sensitive to cI: anges around the tropopause levell, where trends are diriit: tdt to estimate due to a 18ack 01'

reliable observations 2nd tho \: xry large , natural variability ThcL patterns o f both troposphcric anti stratospheric ozone changes are spatially variable. Estimation of t! le radi#: ltive forcing d u ~ to changes in ozone is thus more comples than for the well mixed greenhouse gases.

8.5. I Tropcspheric Ozone In the troposphcrc. ozone is produced during the I oxidation of tfxthnnc and from vnri'ous st lort livcdw precursor gases (mainly c; srbon monoxide (CO), nitrogen oxides CNQ,) and non nlcthane hydrocarbons (NMIiC1). Ozone is a. lso transported into the troposphcre from the stratosphrrc IChangcs in ,troposphc~ ic ozonc conccntratilJru, are spatially variable, both regiocnlly and rsticall,: matking assessment of global ilnr1q term trends difficult. I n the Northern iiemisphere. there is some evidence that tropospheric ozone concentrat~ t~ nS have increascd since 1900, with str1on; g widcnlcc that Ih'k has 13ccurrcd in many locations sino: the 19GOs. However, the obsen? ations of h e most recent dccade stlow that the upward trl:> nd has SIO! AWI significantly or stopped. Model simulations and the limited observations together suggest tha. t ozone conccntrmntions thrloughout the troposphere may have douhicc! in the Piortlhern Hemisphcrr:, since pre industrial times, an increase of about 25 ppbv. In the Southern Hemisphere, therc! are insuffcient ldata to determine if tropospheric 0ron. e has has been objervi: d sincc the mid 1 9:: Os. Likiigcd, e. '\~ t. pi <.; LlL dLluiil i c,; u .+ iit. lc ~~c :~~~l ~~~,.j t .

Changes in tropospheric ozone hake potentially important ~cOnsfquence.; for radiative fcrcirrg. The calculated giobal averq? radiati. tt Iurcing dule to the increased cnncer?. tratioil since pre industrial 1: imes is +0.4 (50.2) SVm 2.

..

. . Year ', li Figure 4: Global CFC I1 concentrations (pptcr. ifar 1978 to 1994. As one ofthe ozone depletinggases. rhc emissiohs of CFC I1 are' conrrolled under the Aforzrrecl! Protocol and its Adjustments and Amendrnerlts. Observations at some individual measurement sites are also shown!.

. 8.5.2 Stratiospheric Ozone Decreases in stratosphcric ozone haw occurred since the 14170s. principally in the iowt: r stratosphere. The most ob7; ious featurc i s t h e annual appcarnnce of thc Antarctic "ozonc hole" in Scptonibcr and Octobw. The October average totill w o n 0 values elver Antarctica are 50 70% IOWCI. t h a n those observed in the 1960s. Statistically siznificant losses in tomtni ozone have also been observed in the mi$ latitudes of both hemispheres, Little 01 n'ca dorvnw, ard trcnd in ozone has beet? observeld in the tropics (2QPN 20" S). The iveight 08f recent scientific evidence strengthens thr pre~ ious conclusion that ozonle loss is due lacgely to anthroplogenic chlorine and bromine compounds. .

S i l :c e the stratospheric ahundances of chlorine and bromine arb expected to continue to grot$ for a reit; more years before they decline (see Section B. 41, s, rrntosplheric ozone losses are expected to peak near t h f end of the century, with a gradual recover) r L! . .1 .,.. .x.. ~ ,." t,.. ,ddi. vk. u >.L :. 'aL !iu; T u. ' ?i> L CGLLL;,.

The loss of ozone in tho loiver StratosptlPre over '

the past 1: s to 20 years has led to a globallJr a: cragecl radiative forcing of about 0.1 \Vn1.2. This nepntive radiative famin:: represents an indirect effr. ct o f anthropogcqic chlorine and brom! ne coxpounds:

.. .. .. ., ,\ 1 .. .. .. .. .. B. 6 Tropospheric and stratospheric aerosols  Aerosol is a term used for particles and very small droplets. of natural and human origin that occur in the atmosphere; they include dust and other particles which can be made up of many clifferent chemicals. Aerosols are produced by a variety of processes, both natural (including dust storms and volcanic actiLVity) and anthropogenic (including fossil fuel and biomass burning). Aerosols contribute to visible haze and can cause a diminution of the intensity of sunlight at rhe ground. Aerosols in the atmosphere influence the radiation balance of the Earth in two ways: (i) by scattering and absorbing radiation  the dfrect effect, and (ii) by modifying the optical propertds. amount and lifetinle of clouds  the irrdircct effect. Although some aerosols. such a s soot, tend to warm the surface, the net climatic effect of anthropogenic aerosols is believed to be a negative radiative forcing, tending to cool the surface (see Section 13.7 and Figure 2). Most aerosols with anthropogenic sources are found in thc lower troposphere (bclotv 2 km). Aerosols u n c l c r g o c h e m i c a l a n d p h y s i c a l transformations in the atmosphcp. especially within clouds, and are removed largcly by precipitation. Conscqucntly acrosols in thc loivcr troposphere typically have resiclcncc times o r a f c k days. Bccausc. of thcir short lifctimc. aerosols in the lower troposphere nr! distributcd inhomogcncorrsi~ with masinla close to the natural (especially desert) .and anthropogenic (especially industrial and biomass combustion) source regions. Aerosol particles resulting from volcanic activity can reach the stratosphcre where they are transported around the globe over many months or years. . '

The radiative forcing due to aerosols depcnds on the size. shape and chemical composition of the particlcs and the spatial distribution of the aerosol.

. \iVllilc 'these factors are comparatively well known for stratospheric aerosols, there remain many uncertainties concerning tropospheric aerosols. Since IPCC (1994). there have been several a d v a n c e s i n u n d e r s t a n d i n g t h e i m p a c t o f tropospheric aerosols on cliinate. These include: (i)

sulphate aerosol largely resulting from fossil f w l combustion and (ii) the first calculation of the spatial distribution of soot aerosol. The .impact of these developments on the calculation of aerosol radiative forcing is discussed in Section 13.;. ..

II": '. ' C:.!: L! s:": 5 "T Li, C 5prliidl u l ~l r l u u r i u l l ~ u i 24 9.7 Summary of radiative forcing Globally averaged radiative forcing is a usefill concept for giving a first order estimate of the potential climatic importance of various forcing mechanisms. Hojvever, as was emphasised in IPCC (199111, there are limits to its utility. In particular. the spatial patterns of forcing differ betireen the globally well mised greenhouse gases, the regionally varying tropospheric ozone, and the even more regionally concentrated tropospheric aerosols. and so a comparison of the global mean radiative forcings does EO: give a complete picture of their possible cliniatic impact. Estimates of the radiative forcings due to changes in greenhouse gas concentrations since pre 'industrial tinles remain unchanged from IPCC (19.94) (see Figure 2). These are +2.45 it. 111 2 (range: +2.1 to +2.8 Wm 2) for the direct effectof thc main well rnised grcenhouse gases (CO,. CH,, K20 and the halocarbons), +0.4 IVm 2 (range: 0.2 to 0. G \V d ) for tropospheric ozone and 0.1 iVm 2 (range: 0.05 to 0.2 VVm 2) for stratospheric ozone. The total direct forcing due to anthropogenic aerosol (sulphates. fossil fuel soot and organic aerosols from biomass burning) is cstimated to be 0.5 iVm 2 (range: 0.25 to 1.0 IYm 2). This estimate is smaller than that given in IPCC (1994) oLving to a reassessment of the model results used to dcrivc the geographic distribution of aerosol particlcs and the inclusion of anthropogenic soot aerosol for the first time. Thc dirrtcr forcing due to sulphate aerosols resulting from fossil fuel emissions and smelting is . estimated to be 0.4 \l'm z (range: 0.2 to 0. S LVm. '). Thc'first cstimates of the impact of soot in aerosols from fossil fuel sources have bccn made: significant uncertainty remains but an estimate of +0.1 \t'm z (range: 0.03 to 0.3 Wm 2) is made. The direct ,

.radiative forcing since 1850 of particles associated with biomass burning is estimatdd to be 0.2 \Vm n (range: 0.07 to O h IVrn 2). unchanged from IPCC (1994). I t has recently been suggested that a significant fraction of the tropospheric dust aerosol is' influenced by human activities but the radiative forcing of this component has not yet been quantified: The range of estimates for thc rartistivr fnrri? due t o changes in cloud properties caused by aerosols arising f: G: x iiutnan activity (the indiruct effect) is unchanged from IPCC (1994) at between 0 and 1.5 iVm*~. Sel era1 ~e t y wclics confirm that the indirect effect of aerosol may have caused a substantial negative radiatiw forcing since pre .. I industrial times, but it remains very dilTicu! t to quantify, more so than the direct effect. \Vhile no best estimate of theindirect forcing can 'currently be made, the central vaiue of 0.8 i\ 'rn z, has been used in some of the scenario cAculntions dmcribcd i n Sections B. 9.2 and F 2. There are no significant alterations since iPCC (1994) in the assessment of radiative forcing caused by changes in solar radiatiw outuut or stratospheric aerosol loading resuicing from volcanic eruptrons. The estimate of radiative forcing due to changes in solar radiative output since 1550 is +0.3 il'm. 7 (range: +0.1 to +0.5). I3ndiatiL. e forcing due to volcanic aerosols resultiEe from an individual eruption can be large (the maximum global mean effect from the eruption of h l r . Pinatubo \vas 3 to 4 Wm 2). but lasts for only a f e w ). ears. HIDivever. the transient variations in both these forcings may be important in explaining sornc of tile obsen. ed climate variations on decadal time scales.

B. 8 Global Warming Potential (GWP) The Global Warming Potential is an al: tempo to provide B simple measure o f t h e relative radiative effects of the emissions of various greenh'ouse ga. ses. The indes is defined as the cumulative radiative *

forcing between the present and some chosen tido horizon caused by ;ii unit mass of gas emitted nown expressed relative to that Far some reference gas (here CO, is used). The hlture global warming; commitment of a greenhousc gas over a chosen timi horizon can be estimated by multiplying the appropriate GWPby the arnolu~ nt ofgas emitted. For exan~ plc. GI'VPs could be used t o compare! the efTects of reductions in C02 (emissions lrelative to reductions in methane emissions, for a specified time horizon. Derivation of GWPs requires knowledge of the fate of the emitted gas and the rndiarive forcing due to the amount remaining in thc atmosphere. .4lthough the GWPs are quoted ,as 'single values. the typical uncertainty is r35%. not inciuding the uncertajnty in the carbon dioside. ieference. Because GVVPs are based on the radiative forcing lconcept. tlhey are difficult to apply to radiatively important constituents that are unevenly distributed in the atmosphere. No attem'ot is madle to define R GWP for aerosols. Additionally the clnolce of time ho'rizon >will cicpcnd on policy cor,; idera:. cns. GWPs need to take accour! t of any indirect effects of the emitted greenk. ou& 6 ~s i f they are to re! lect correctly future warming potentlial. The net GZ'v'Ps for the ozone depleting gases. ruhich inlrlude the ..

.. . .. direct "warming" and indirect "cooling" effects, have now been estimated. In IPCC (1991), only the direct GWPs were presented for these gases. The indirect effect reduces their GWPs. but each ozone depicting gas must be considered individually. The net GWPs of the chlorofluorocarbons (CFCs) tend to be positive. while those of the halons tend to be negative. The calculation of indirect effects for a number of other gases (e g , NO,, CO) is not c u r r e n t l y p o s s i b l e b e c a u s e of inadequate characterisation of many of the atmospheric, processes involved. Updates or new GWPs are given for a number of kel, species (Table 4). based on improved or new estimates of atmospheric lifetimes, molecular r a d i a t i v e f o r c i n g f a c t o r s . a n d i m p r o v e d representation of the carbon cycle. Revised lifetimes for gases destroyed by chemical reactions in the Eotver atmosphere (particularly methane, HCFCs and HFCs ) have resulted in GiVPs that are slightly 1ov. w (typically by 10 15% 0) than those cited in IPCC (19941. The IPCC definition of GWP is based on calculating the relative radiative impact 'of a release of a. trace gas over a time horizon in a constant background atmosphere. In a future atmosphere with larger CO2 concentrations, such as occur in all of the IPCC emission scenarios (see Figure 5b). we wduid calculate slightly larger GWP values than those given in Table 4. id B. 9 Emissions and concentrations of greenhouse gases and aerosols in the future B. 9.1 The IS92 emission scenarios The projection of future anthropogenic climate c h a n g e d e p e n d s , a m o n g o t h e r t h i n g s , o n assumptions made about future emissions of greenhouse gases and aerosol precursors and the proportion of emissions remaining in the atmosphere. Here we consider the 1592 emission scenarios (IS92a to D which were first discussed in IPCC (1992). The IS92 emission scenarios extend to the year 21100 and include emissions of COz. CH4, NzO, the halocarbons (CFCs and their substitute HCFCs and HFCs), precursors of tropospheric ozone and bul! pt!: Itc aerosols and z~ rosols from biomass burning. A wide range of assumptions regarding f u t w p economic, demographic. and policy factors are encompassed (IPCC. 1992). In this report, the emissions of chlorine sild bromine containing . .. . .. .~

..

~ . ~ '. " ..

25 Table 4: Globul Wurming Potential referenced to the updated decay response for the Bern carbor: cgclc  and future C02 atmospheric concentrations held constant at current iecels.

Species Chemical Lifetime Global \! 'arming Potential Formula (T h e Horizon) (years) 20 years 100 years 500 years Methane* CH4 1253 56 21 6.5 Kitrous oxide X20 120 280 310 170 26 halocarbons listed :II IS92 u r e assumed to be phased ou& under :he blontreal Protocol and it5 Adjustments and ;kmendnnenI: s and so a single revised future emissioa scenario for these gases is incorporated in all oi' :he 1592 xenarios. Emissions of indi<, idual HFCs are based on the original IS92 scenari. iis, aichourh they do n o t reflecr current markets. CO, emission:; for the si4 scenarios are shown in Figure 3?.. The calculation 3 ; future concentrations of greenhouse gases. g: ven certain emissions, entaiis modelling the processes tilet transform and remove the different. gases from the atmosphere. For esample, future concentrations of C 0 2 are calculated using rnodt. 1~ of the carbon cycle which m o d e l t h e e x c h a q e s of COz bletiveen t h e a t m o s p h e r e a n d t h e o c e a n s a n d t e r r s s t r l a l biosphere (see Section B. 1); atmospheric chemistry m o d e l s a r e u s e d to sirrulatle the removal of chemically ,active ga. sss such 85 methalne. All the IS92 emission scenarios, even IS92c. imply increases' in greenhouse gas concentrations from 1990 to 2100 (e. g., CO? increases range from. 35 to 170% (Figure 5b); CH2 From 2'2 to 175%; and N 2 0 from 26 to ,40%).

dependent on the tconcentration of the gas and ti,; strength with which i t absorbs and ne emits long 'wave radiation. For suiphnte aerosol. the direct a :d indirect radiative forclngs ~e :e calculated on th. 8 basis of sulphur emissions contained in the IS92 scenarios. The radiarlT. e Forcing due to aerosol from biomass burning was assumed to remain constant at 0.2 Wm 7 after 1990. The contribution from aerosols is probabl! the tnosc uncertain part of future radiative forcing. Figure 61, a shows a single "best estimate" of historical radiative forcing from 1765, to 1990 (including the effects of aerosols). foilobved by radiative forcing for Scenarios IS92 a to f. Figures 6b and c show the contribution to future mdiative . forcing from variolus cornpolxnts of the iS9'a Scenario; the largest contribution comes From C02, with a radiative forung of ~Imost c 6 Wm. 2 by 2100. The negative forcing due tu tropospheric aerosols. ir. 7. g?!~ bzl! '! l. averay 2 501's~ pffsets scmo of t h p greenhouse gas positive forcing. However. because tropospheric aerosols are h~ ghiy variablle reglonalL)., their globa. li!. averapcl radiative forcing will not adequately ldescribe thclir possible climatic impact. Future projections o f tenlperature and sea level based on the IS92 ernlssims sclenarios a r e discussed in Sectil? n f:. For greenhouse gases,, radiative forcing is '

I .' ' ..

. .. , . . . I.

.. i" _,

IS92t IS92a IS92b IS92d IS92c 3 0 J "J "' J I f 2000 , 2020 21340 2060 2080 2100 ..

u a Year __. .. IS92eIS92f . 2000 ~

2020 2040 2060 2080 2100 .. .. Year Figure 5: [a) Total anthrbpogenic C02 emissions under the . IS92 emission scenarios and (3) the resulting atmospheric CQz concentrations calculateld using the 'Bern' carbon cycle modelhnd the carbon bludget for the 1980s shown in Tdlbie2. ', : 28 .. . . _. B. 9.2 Stabilisation o f greenhouse gas and aerosol concentrations An important questi; to consider is: how might greenhouse gas concentrations be' stabilised in the future? If global CO, emissions were maintained at near current (1991) levels, they would lead to a nearly c o n s t a n t r a t e  of increase in atmospheric concentrations for at least two centuries, reaching about 500 ppmv (approaching twice the pre industrial concentration of 280 ppmv) by the end of the 21st century. In IPCC (19941, carbon cycle models were used to calculate the emissions of COz which would lead to stabilisation at a number of different .concentration levels from 350 to 750 ppmv. The assumed concentration profiles leading to stabilisation are shown in Figure 7a (excluding 350 ppmv). Many different stabilisation levels. time scales for achieving these levels. and routes to stabilisation could have been chosen. The choices made are not intended to have policy implications; the exercise is illustrative of the relationship between COz emissions and concentrations. Those in Figure i a assume' a smooth 'transition from the current average rate of CO, concentration increase to . stabilisation. To a first approximation, the stabilised concentration level depends more upon the accumulated alnount of carbon emitted up to the time of stabilisation. than upon the exact Concentration path followed en route to stabilisation. New results have been produced to take account of the revised carbon budget for the 1980s (Table 2). but the main conclusion. that stabilisation of concentration requires emissions eventually to drop .

well below current levels. remains unchanged from IPCC (1994) (Figure ibl. Because the new budget implies a reduced terrestrial sink. the allowable emissions to achieve stabilisation are up to 10% lower than those in IPCC (1994). In addition, these calculations have been extended to include alternative pathways towards stabilisation (Figure 7a) and a higher stabilisation level (1000 ppmv). The alternative pathways assume higher emissions in the early years, but require steeper reductions in stabilisation case allo~ vs higher maximum emissions, but still requires a decline to current levels .by about 240 years from now and further reductions thereafter (Figure 7b). cL.,& idL; a ILt; u~ )c a t s (1.1gule 7bj. Tile ibbd ppr11c  .,. .. Year F i p m 7: (a) CO, concentration profiles 1radii; g t o stabihsation at 450, 550. 650 and 750 pprnvJol! iou~ ing the pathuays defined in IPCC( 1994) (solid curms) and for p a t h w y s that allow emissions to follow lS32a plntil at least 2000 (dashed curves). A single projile that srabiiises at a Ccl, concentration of IO00 ppmv and folious 1592a emissiwu until at least 2000 has also been deflned. (b! cO, emissims Leading to stabilisation at concentrarions ofi'30. 550, 6.70. 750 and 1000 ppmv following rhe p r a J k s shorrn I in [a). Current anthropogenic CO, emissions and those f o r IS92a 'Ire shown for comparison. The ca! cn! c! ions ltse the ' B e n ' ' carbon cycle model and the carboll budgetybr the 1980s shown in Table 2. . Table 5: Total anthropogenic CO, emissions accumulatedfrom 1997 to 2100 inclusive (Gt&). All values were calculated using the  carbon budget for 1980s shown in Table 2 and the Bern carbon 3 _. cycle model.

Case Accumulated C 0 2 emissions 1991 to 2100 [GtC)

IS92 scenarios c 770 d 980 .. b 1430 . a 1500 .. f 1830 . . . e 2190 ... .. .... .. .. ........ .. .. .~ ~ ............ ........................... ... ... ....... . .' ..... ,.   .. .  _.; .. .. ..

". The accumulated anthropogenic COz emissions from 1991 to 2100 inclusive are shown in Table 5 for the profiles leading to stabilisation at 450, 550, 650, 750 and 1000 ppmv via the profiles shown in Figure 7 a a n d , for comparison, the ' IS92 emission scenarios. These values are calculated using the ." Bern" carbon cycle model. Based on the results in IPCC (1994) it is estimated that values calculated with different carbon cycle models could be up to approximately 15% higher .or, lower than those presented here. If methane emissions were to remain constant at 1984 1994 levels (i. e., those sustaining an ..

atmospheric trend of +10 ppbv/ yr), the methane concentration would rise to about 1850 ppbv over the next 40 years. If methane emissions were to remain constant at their current (1994) levels (Le., those sustaining an atmospheric trend of Bppbv/ yr), the methane concentratinn wnqlr! rise fn rhntit !c? n .. ppbv over the next 40 years. If emissions were cut by about 30 Tg( CH4)/ yr (about 8% of current anthropogenic emissionls), CH4 concentrations would remain at today's levels. These estimates are lower than those in IPCC (1994). If emissions of N20 were held constant at toda! 's level. the concentration would climb from 312 ppbv to about 400 ppbv over several hundred years. In order for the concentration to be stabilised near current levels. anthropogenic sources would need to be reduced by more than 50%. Stabilisation of PFCs and SF6 concentrations can only be achieved effectively by stopping emissions. Because of their short lifetime, future tropospheric aerosol concentrations would respond almost immediately to changes in emissions. For esample. control of sulphur emissions would immediately reduce the amount of sulphate aerosol in the qt 7ncnhn n ...... ........

. .... ". .  .. I ' . ,. .. I C. Observed Trends and Patterns general and desertification could have contributed only a small part (a few hundredths of a degree) of I Section B demonstrated that human activities have the overall global warming, although urbanisation , influences may have been important in sonle regions. Indirect indicators, such ' as borehole temperatures and glacier shrinkage, provide independent support for the observed warming. Recent years have been among the warmest since 1860, Le.. in the period of instrumental record. The warming has not been globally uniform. The r e c e n t w a r m t h h a s b e e n g r e a t e s t o v e r t h e 1 continents between 40% ' and 70" N. A fen. areas.

C. l Has the climate warmed? such as the North Atlantic Ocean north of 30's. and Global average surface air temperature, excluding some surrounding land areas, have cooled in recent Antarctica, is about 15° C. Year to year temperature ' decades (Figure 9). changes can be computed with much more As predicted in !Pee (1992) and discussed in confidence than the absolute global average IPCC (19941, relatively cooler global surface aut1 temperature. tropospheric temperatures (by about 0.5" C) and a T h e m e a n global s u r f a c e t e m p e r a t u r e h a s relatively warmer lower stratosphere (by about increased by about 0.3" to 0.6" C since the late 19th 1.5OC) were observed in 1992 and eariy 1003, century, and by about 0.2" to 0 .X Over the last 30 following the 1991 eruption of &It. Pinatubo. years, the period with most credible data (see Figure it'armer temperatures at the surface and i n the 8 which shows data up to the end of 1994). The ' l o w e r . t r o p o s p h e r e . a n d a cooler lower 1 in Climate and Sea Level changed the concentrations and distributions 'Of greenhouse gases and aerosols over the 20th 1 century; this section discusses the changes in temperature, precipitation (and related hydrological variables), climate variability and sea level that have been observed over the same period. Whether the observed changes are in part induced by human 1 i activities is considered in Section E.

, j I warming occurrcd largely ,during two periods, stratosphere, reappeared in 1994 following the i .

I.

changed since the IPCC (1990) and IPCC (1992J.; The general tendency toward reduced daily Warming is evident in both sea surface and land temperature range over land, at least since the based surface air temperatures. Urbanisation in middle of the 20th century, noted in IPCC (1992),

 temperatiirc> s 1% '1 IS6 I IO 1994. re1utir. r tu 1961 10 ' . 1990. The solid c~~ rct' represehts snluothircg $IIW annrcul I? alws slronx 69 t l w bars to suppress srtb tir~ ac! nl time scale rariutiwn, The dashed sncoutirecl ~1 1 1 1 ~ is thL, corresponclirry w w l t J i m c 1860 1880 1900 1920 . 1940 1960 1980 2000 Year :: I .. 2 1.5 1 0.5 0 0.5 1 1.5 2 Figure 9: Change (fro112 1955T4 to 1975 9 1) of arlnrtal land snrface air temperature and sen strrfnce ternper~ trrre.

has been confirmed with more data (which have now been analysed for more than 40% of the global land area). The range has decreased in many areas because nights have warmed more than days. Minimum temperatures have typically increased twice as much as maximum temperatures over the last 40 years. A likely esplanation, in addition to the effects of enhanced greenhouse gases. is an increase in cloud cover which has been observed in many of the areas with reduced diurnal temperature range. An increase in cloud reduces diurnal temperature range both by obstructing daytime sunshine, and by preventing the escape of terrestrial radiation at night. .Anthropogenic aerosols may also have an influence on daily temperature range. Temperature trends in the free atmosphere are more difficult to determine than at the surface as there are fewer data and the records .are much shorter. Radiosonde data which are available since the 1950s show warming trends of around 0.1" C per decade. as at the surface, but since 19T9 when satellite data of global tropospheric temperatures became available, there appears. to have been a J ..~.,L 1: !. 4 <u>: iilg (abdui 0. ciO'i per decade), rvtlereas . . surface measurements still show a warming. Thpsp apparently contradictory trends can be reconciled if the diverse response of the troposphere and surface to short term events such as volcanic eruptions and' El NiAo are taken into account. After adjustment for ..

32 * .. these transient effects, both tropospheric and surface data show slight warming (about 0.1" C per decade for the troposphere and nearly 0.2'C per decade at the surface) sincc 1973. Cooling of the lowcr stratosphere has been observed since 1979 both bq' satcllitcs and weather balloons. as noted in IPCC (1992) and IPCC (1994). Current global mean stratospheric tcnlpcrnturcs are the coldest obsen ecl in thc relatively short period of the record. Reduced stratospheric tcmpcrature has been projected to accompany both ozonc losses in the lower stratosphere and atmospheric increases of carbon dioside.

C. 2 Is the 20th century. warming unusual? In order to .establish whether the 20th century warming 'is part of the natural variability of the climate system or a response to anthropogenic forcing. information is needed on climate variability on relevant time scales. As an average over the Northe'rn Hemisphere for summer. recent decades. appear to be the warmest since a t least 1400 from tilt. illiiiieu avauade evidence (Figure 10). The warminn over t h P past century began d1.1ri~ g one qT the colder periods of the last 600 years. Prior to 1400 data are insufficient t o provide hemispheric temperature esTirnates. Ice core data fron: se,< e: al sites suggest that the 20th century is at least as .. warm a s any century in the past 600 years, although the recent warming is not exceptional everywhere. Large and rapid climatic changes occurred during the last glacial period (around 20,000 'to 100.000 years ago) and during the transition period towards the present interglacial (the last 10,000 years, known as the Holocene). Changes in annual mean temperature of about 5° C occurred over a few decades, at least in Greenland and the Korth Atlantic, and were probably linked to changes in oceanic circulation. These rapid changes suggest that climate may be quite sensitive to internal or external climate forcings and feedbacks. The possible relevance of these rapid climate changes to future climate is discussed in Section F. 5. 3 Temperatures have been less variable during the last 10,000 years. Based on the incomplete evidence available, it 'is unlikely that global mean temperatures have varied by more than 1'C in a century during this period.

C. 3 Has the climate become wetter? As noted in IPCC (1992). precipitation has increased over land in high latitudes of the Northern Hemisphere. especially during the cold season. A step like decrease of precipitation occurred a h r the 19GOs over the subtropics and tropics from Africa to Indonesia. These changes are consistent with changes in streamflow. lake levels and soil moisture (whcre data analyses are available). Precipitation. averaged over global land areas. increased from the start o f the century LIP to about 19GO. Since about 1980 precipitation over land has decreased (Figure 11). There is evidence to suggest increased precipitation over the central equatorial Pacific Ocean in recent decades, with decreases to the north and south. Lack of data prevents us from reaching firm conclusions about other precipitation changes over the ocean. Estimates suggest that evaporation may have increased over the tropic. al oceans (although not everywhere) but decreased over large portions of Asia and North America. There has also been an observed increase in atmospheric water vapour in the tropics, at leasr since 1Yi3. 'Cloudiness appears' to have increased sinc'e the 1950s over the oceans. In many land areas where the daily temperature range has decreased (see Section C. 1). cloudiness increased from the 1950s to at least the 1970s. I .. Northern Hemisphere Summer (JJA)

o_ 0.2 Figure 10: Decadal summer (June to August) temperature index for the Northern Hemisphere (to 19i0 19iqJ based on 16 proxy records (tree rings. ice cores, docwnentnrzJ records) from North America. Europe and East Asia. The thin line is a smoothing of the same data. . lnomnlies are relative to 1961 to 1990.

a 1900 * 1920 1940 1950 . 1980 Year Figure 1 I: Changes in land surface precipitatiou arernged over regions between 55's and 8SN. Annual precipitation depurtures from the. 1961 90 period are depicted by the hollow bars. The continuous curue is a smoothing of (he sclr71c'

data.

..

Snow cover extent over the Northern Hemisphere land surface has been consistently below the 21 pear average (1974 to 19943 since 19SS. Snow radiation feedback has amplified springtitlie warming over mid to high latitude' Northern Hemisphere land areas. A summary of observed climate trends is sholvn in Figure 12. ..

3 3 (a) Temperature indicators (0.6 "C decrease 1979 94) *** stratosphere *' 0.3 "C increase 1958 94 troposphere *" little change 1979 94 .. 0 . 4 .

*** N. H. spilng snow cover I (1 0% decrease1973 94) '\, .

"""""""" *** near surface air C"""""""

.......... ;** land night time air .; .. '.!. '. ' ....... .:. ...  . ...... *** near surface ocean (0.3 0.6 "C increase since temperatures rising faster .: ** sea ice (1 973 94) . ~

..

late 19th century) than daytimetemperatures . during 1990s ' ' N. H. below avg. (1951 90) : .:.,: 1, ... ~

...... ... ................ .. .......... ,_. .: ': Tf groundtemperaturesS. H. avg. ..... ... ..... .. (mostlywarming)

.. 33 .. ... _.. t .. , I *. , " .. , .

.... .... ... ... ... ........ ...... ... .. . .... ...  . .. .. '. . : 1 .. .. .. . .. C. 4 Has sea level risen? Over the last 100 years global sea level has risen by about 10 to 25 cm, based on analyses of tide gauge , records. A major source of uncertainty in estimating the rate of rise is the influence of 'vertical land movements, which are included in sea level measurements made by tide gauges. Since IPCC (1990), improved methods for filtering out the effects of long term vertical land movements, as well as greater reliance on the longest tide gauge records for estimating trends, have provided greater confidence that the volume of ocean water has, in fact, been increaSing and causing sea level to rise within the indicated range. It is likely 'that much of the rise in sea level has C. 5 Has the climate become more variable and/ or extreme? Many of the impacts of climate change may result from changes in climate variability or estreme weather events. Some reports have already suggested an increase in variability or extremes h a s t a k e n place in recent decades. Do meteorological records support this? There are inadequate data to determine whether consistent global changes in climate variability or extremes have occurred over the 20th century. On regional scales there is clear evidence of changes in some extremes and climate variability in$ icators (e. g., fewer frosts in several widespread areas: an'increase in the proportion of scale, the warming and consequent espansion of the oceans may account for about 2 to 7 cm of the observed rise in sea level, while the observed retreat of glaciers and ice caps may account for about 2 to 5 cm. Other factors are more difficult to quantify. Changes in surface and ground water storage may have caused a small change in sea level over the last 100 years. The rate of observed sea level rise. suggests that there has been a net positiv? contribution from the huge ice sheets of Greenhpd and Antarctica, but observations of the ice sheets do not yet allow meaningful quantitative estimates of their separate contributions. The ice sheets remain a major source of uncertainty in accounting for past changes in sea level, because there are insufficient data'about these ice sheets over the last 100 years. .. States of the USA). Some of these changes have been toward greater variability: some have been toward lower variability. There have been relatively frequent El Niiio Southern Oscillation warm phase episodes, with only rare excursions into the other estreme of the phenomenon since 1977, as noted in IPCC (1990). This behaviour, and especially the persistent warm phase from 1990 to mid 1995. is unusual in the last 120 years (i. e.. since instrumental records began). The relatively low rainfall. over the subtropical. land areas in the last two decades is related to this behaviour.

.. BOX 2: What took are used to project future climate and how  are they used?

3 G D. Modelling Climate and (a) Obse. wed Climate Change surface air temperature (OC) December Februari * Climate models which incorporate, in various degrees of complexity. mathematical descriptions of t h e a t m o s p h e r e , o c e a n , l a n d , b i o s p h e r e a n d cryosphere. are important tools for understanding climate and climate change of the past, the present cl A \J +:. r ;& 'j '. : LyF : ,:, d;

and the future. These models. which use primarily physical lalvs and physically based empirical relations, are very much more complete than. for example. models based on statistical relationships used in less quantitative disciplines. Detailed (b) Model averaae surface air ternDerature projections of future climate change rely heavily on coupled atmosphere ocean models (see Box 2). How much confidence should we have i n predictions from such models?

D. l The basis for confidence in climate models As discusscd in Section B, changes in the radiatively active trace gases in the atmosphere produce radiative forcing. For cquivalcnt CO, concentrations equal to twice the pre industrial concentration, the positive radiative forcing is about +4 \Vm '. To restore the radiative balance other changes in climate must occur. The initial reaction is for the. locvcr atmosphcrc (the troposphere) and the Earth's surface to warm: in thc absrncc of othcr changcs. the warming would be about 1.2" C. I fon. cvrtr. hcating not only changes tcmpcraturc's. but also altcrs other aspects of the cliunate systcnl and various feedbacks arc invokcd (see Section D. 2). The key role of climate moclcls is to quantify those fecdbacks and determine the ovcrall climate response. Further, .the warming and othor clilnatc effects will not be uniform over thc Earth's surfacc; an important role of models is, to simulate possiblc continental and regional scalc climate responscs. Climate models include, bascd on our current undcrstanding. the most important large scalc .physical processes govcrning the climate system. Climate models have improved since IPCC (1990). but so too has our unclcrstanding of thc complcsity of the climate system and thc recognition of the need to include additional processes. In uruec to assess tile i ;t I u c ui' 't 1il" ici hi projections of future climate. its simulated climate can be compared with known features of the observed current climate and. to a less satisfactory degree. with the morc limited information from significantly different past climate states. I t is , ..

..

.. .. . . .. .. I (4 Observed precipitation rate (rndday) June August (d) Model average precipitationrate 3i important to realise that even though a model may have deficiencies, it can still be of value in quantifying the climate response to anthropogenic climate forcing (see also Bos 2). Several factors give us some confidence in the ability of climate models t o simulate important aspects of anthropogenic climate change in response to anticipated changes in atmospheric composition:

(i) The most successful climate models are able to simulate the important large scale features of the components of the climate system well, including seasonal, geographical and vertical variations which are a consequence of the variation of forcing and dynamics in space and time. For example, Figure 13 'shows the geographical distribution of December to February surface temperature a@ June to August precipitation .simulated by comprehensive coupled atmosphere ocean models of the type used for climate prediction. compared nith observations. The large scale features are reasonably well captured by the models, although at regional scales more discrepancies can be seen. Other seasons are similarly well simulated. indicating the ability of models to reproduce the seasonal cycle in response to changes in solar forcing. The improvement since IPCC (1990) is that this level of a, ccuracy is achieved in models with a fully interactive ocean as compared to the majority of models that employed simpler schemes used in 1990.

(ii) Many climate changes are consistently projected by different models in response to greenhouse gases and aerosols and can be explained in terms of physical processes which are kn0Lr. n to be operating in the real world. for example. the maximum ivarming in high northern latitudes in winter (see Section F).

(iii) The models reproduce with reasonable fidelity other less obvious variations in climate due to changes in forcing:

* Some atmospheric .models when forced' with observed sea surface temperature variations can reproduce with moderate to good skill several regional climate variations, especially in parts of the tropics and sub tropics. For 'esample. a s p e c t s o f the l a r g e s c a l e i n t e r a n n u a l ..?.:;; Lc: c ELL~ G~ L~ U, A> UI. GI tile iruplcai Packtic, r e l a t i n g t o t h e El Nifio Southern Oscillation phenomenon are captured, as are interannual Variations in rainfall in north east Brazil and to some extent decadal variations in rainfall over the Sahel. dl 3s * AS discussed in IPCC (199.11, stratospheric aerosols resulting from the eruption of hit. Pinatubo in Jtine 1991 gave rise to a short lived negative globel mean radiative forcing of rhe troposphere which peaked at 3 to 4 IVm. 2 a few months after the eruption and had v i r t ~~;~i ~ disappeared by about the end of 1994. .A clin:;: e model was used to predict global tempera: ure variations betwen the time of the eruption a2d the end of 1994 and the results agreed c l ~j ;.!~ with observations (Figure 11). Such a, Ureemenr increases confidence in the ability ofcIin1e: e models to respond in a realistic way to transienr. planetary scale radiative forcings of i a r g magnitude.

0 Previous IPCC reports demonstrated the abilj;;: . of models to simulate some known features of palaeoclimate. Only modest progress has been made in this area, mainly because of the pauciv of reliable data for comparison.

0 Currently available model simulations of global mean surface temperature trend m r the past half century show closer agreement wirh observations when the simulations include the likely effect of aerosol in addition to greenhouse gases (Figure 15).

(iv) The model results eshibit "natural" variability on a wide range of rime and space scales which is broadly comparable to that observed. This "natural" .Variability arises from the internal processes at work in the climate system and not from changes in esternal forcing. Variab'ility is a very important aspect of the behaviour of the climate system and has important implications for the detection of climate change (see Section E). The year to year variations of surface air temperature for the current climate are moderately realistic in model simulations at the larger space scales. For example the smaller variability over the oceans compared with continental interiors is captured. Too low interannual variability of the tropical east and central Pacific Ocean temperatures associated with the .El Nifio Southern Oscillation (ENSO) phenomenon is one deficiency. No current coupled atmosphere ocean model simulates a l l aspects of LI'sbu events, bul some of the 0bserve. d interannual variations in thP almwphere nssbcintccl rT: ith t h e s events are captured. ..

Climate models are c a l i b r a t e d , i n p a r t , by * the systematic adjustments (the so called flux m  adjustments) that are used in some models at the ' atmosphere ocean interface in order to bring the $ '''r simulated climate closer to. the observed state. E 1 These adjustments are used to compensate for models, are 'used to ensure that the simulated present day climate is realistic and hence that c l i m a t e f e e d b a c k p r o c e s s e s o p e r a t e i n t h e Year generally be traced to deficiencies in the simulation of current climate in the unadjusted models. for example, systematic errors in sea ice. The main unknown regarding the use of adjustments in models is the extent to which they allow important against known climate variations including the.; 5 seasonal cycle and the perturbations mentivncd above. This provides some conficlence in their use 3 0.5 activities.

improvements. e. g.. the successful incorporation of a d d i t i o n a l p h y s i c a l p r o c e s s e s (s u c h . a s c l o u d .. . microphysics and the radiative effects of sulphate Figure 15: Simulated global annual mean a arming from aerosols) into global coupled models. and the 1860 to 1990 allowing for increases in greenhouse gases only (dashed curve) and greenhouse gases and sulphate improvement in such models' simulation of the wrosols (solid curve), compared with observed changes observed changes in climate over recent decades. over the same period. Further confidence will be gained as models continue to improve.

D. 2 Climate model feedbacks and uncertainties 2.5" C. The range of the estimate arises from the initial response. The likely equilibrium response have emerged to change Lhese estim& es of th5 of global surface temperature to a doubling of climate sensitivity. The present activities regarding equivalent carbon dioxide concentration (the incorporation of these feedback processes in models "climate sensitivity") was estimated in 1990 to be in are described below.

.. .. Water vapour feedback An increase in the temperature of the atmosphere increases its water holding capacity and is expected to be accompanied by an increase in the amount of water vapour. Since water vapour is a powerful greenhouse gas, the increased water vapour would in turn lead to a further enhancement of the greenhouse effect (a positive feedback). About half of this feedback depends on water vapour in the upper troposphere. whose origin and response to surface temperature increase is not fully understood. Feedback by water vapour in the lower troposplwe is unquestionably positi1. e and the preponclcrence of evidence points to the same conclusion. for upper tropospheric water vapour.

' Feedbacks resulting from changes$ the decrease of temperature with height can partially compensate the water vapour feedback. I Cloud/ radiative feedback Several processes are involved in cloud/ radiative feedback. Clouds can both absorb and reflect solar radiation (which cools the surface) and absorb and emit long wave radiation (which warms the surface). depending 0. n cloud height. thickness and cloud radiative properties. The radiative properties of clouds clcpcnd on the evolution of atmospheric watcr in its vapour, liquid and kii phases and upon atmosphc! ric aerosols. The processes are complcs and. although considerable progress has been made since II'CC (1990) in describing and modelling those cloud pr~ c: c! sses thnt arc most important for dctermilling radiative and t1rnce temprrntrtrc changcs. thcir uncertainty represents a significant source of potcntial error in climate simulation. This potrntinl error can be estimated by first noting that if clouds and sea ice are kept k e d accorcling to thcir observed distributions and propertic% climate models would a l l report climate sensitivitic. s i n the range of 2 to 3° C. &Ioclellers have shown foi. various assumptions that physically plausible changes in, cloud distribution could either as much as double the warming expected 'for fixed clouds or. on the other hand. reduce it by up to 1° C. The range in estimated climate sensitivity of 1.5 to 3.5" C is Iarg~ ly dictated by this uncertainty.

Ocean circn. ' ': '5,2 Oceans play an important role in climate because they c a q large amounts of heat from the tropics to the poles. They also store large amounts of heat. carbon and CO, and are a major source of water to osphrre (through evaporation). Coupling of I ' .... . '. .  .. . ..~.. . . . ,. . atmospheric and oceanic GCMs (see B~~ 2) improves the physical realism of models used for projections Of fUtUre dimate change, particularly the timing and regional distribution of the changes. Several models show a decrease or only margilyal increase of sea surface temperatures in the northel E North Atlantic in response to increasing greenhouse g a s e s , r e l a t e d to a s1on. ing down of the thermohaline circulation as the climate warms. This represents a local negative temperature feedback. although changes in cloud cover might be an important factor. The main influence of the oceans on simulations of climate change occurs because of their large heat capacity, ivhich introduces a dela? in warming that is not uniform spatially.

Ice and snow albedo feedback An ice o r snow covered surface strongly reflecrs solar radiation (i. e.. it has a high "albedo"). As solne ice melts at the warmer surface, less solar radiation is reflected leading to further warming (a positive feedback). but this is complicated by clouds, leads (areas of open water in sea ice) and snowcover. The realism of simulated sea ice cover varies considerably between models. although sea ice models that include ice dynamics are showing increased accuracy.

. Land surface/ atrnosphere interactions Anthropogenic climate changes, e. g., increased temperature. changes in precipitation, changes in net radiative heating and the direct effects of CO,. will influence the state of the lend surface (soil moisture. albedo. roughness. vegetation). In turn. the altered land surface can feed back and alter the overlying atmosphere (precipitation, water vapour. clouds). Changes in the composition and structure of ecosystems can alter not only physical climate. but also the biogeochemical cycles ' [see Section B). Although land surface schemes . used in current GChls may be more sophisticated than in IPCC (1990). the disparity between models in their simulation of soil moisture and surface heat and moisture fluses has not been reduced. Conhdence in calculation of regional projections of soil moisture changes in response to greenhouse gas and aerosol forcing remains low.

C ;i L k i L d c > 11; ~C ~;C L ~;L U I ~ call powaailg tuctner modify climate locally and regionally by altering thP exchange of water and energy between. the lard surface and atmosphere. For esample, forests spreading into tundra in a w r m e r worlci tvould absorb a greater proportion of solar 'energy and *

. ... .  .. ~. feedbacks may have important effects on regional natural in origin. Such natural fluctuations occur on climate change projection. a variety of s p a c e  a n d t i m e s c a l e s , a n d c a n b e purely internal (due to complex interactions E. Detection Of `Iimate Change betiveen individual components of the climate , and Attribution of Causes system, such as t h e a t m o s p h e r e a n d o c e a n ) or respect to the increase in global mean temperature natural variability, in the observations. This is over the last 100 years, IPCC (1990) concluded. that b e c a u s e t h e r e a r e l a r g e u n c e r t a i n t i e s in the the observed warming was "broadly consistent lvirh . evaJ, 2tion and magnitude of both human and natural predictions of climate models, but it is also of the forcings, and in the characteristics of natural same magnitude as natural climate variability". The internal variability, which translate to uncertainties . report went on to esplain that "the observed in the relative magnitudes of signal and noise. increase. could be largely due to this natural In the modelled world, however. it is possible to variability; alternatively this variability and other perform multi century control experiments with no. human factors could have offset a still larger human induced changes in greenhouse gases. human induced greenhouse warming". sulphate aerosols or other anthropogenic forcings. Since IPCC (1990). considerable progress has Since 1990, a number of such control esperiments been made in the. search for an identifiable human have been performed with coupled atmosphere induced effect on climate. osean models, These yield important information on ., the patterns, . time scales, and magnitude of `the "internally generated" component of natural climate Espcrimcnts with GChIs are now starting to ' observed changes can be plausibly incorporate some of the forcing due to huInnn . explained by internal climatic fluctuations. but i n d u c e d c h a n g e s i n s u l p h a t e a e r o s o l s a n d only dne part of the "total llatural ' stratospheric ozonc. Thc inclusion o f . these variability ofclimate (since such control runs do not induced climate change from a large (> 18) number of transient experiments in which coupled well misecl greenhouse gases; see "equivalent C02" only. These investigations compared observed in the Glossary). some of these experiments have changes over the past 10 100 years with estimates been repeated with identical forcing but starting of internal or total natural variability noise derived i. ii,. i &. .... ...... L... igiti:].! r l ;v ?.t P ~+o t n . Sl.! rh frnm palneodata. climate models, or statistical repetitions help to better define the espected models fitted to observations. Most. bur nor all ot r . ... .. ,.I :.,L +I ,. cl: r? n ,..,,.

natural fluctuations of the climate system.

." 1850 1880 1910 1940 1970 2000 Time (years)

 0.75 E 5 2 0.50 m c 0.25 m p! . x 0.00 m Ql p 0.25 L.

5 I 0.50 1850 1880 1910 1940 1970 2000 Time (years) Figure 16: Observed changesin global mean temperature over 1861 to 1994 compared with those silldated using an upwelling diffusion energy balance climate model. The model was runprst with forcing due to greenhohe gases alone (a) and then with greenhouse gases and aerosols (3).

These global mean results cannot establish a clear cause and effect link between observed changes in atmospheric greenhouse gas concentrations ,and changes in the Earth's surface temperature. This is the attribution issue. Attribution is. difficult using global mean changes only because of uncertainties '. in the histories and magnitudes .of natural and human induced forcings: there are many possible combinations of these forcings that could yield the same curve of observed global mean temperature change. Some combinations are more plausible than others, but relatively few data exist to constrain the range of possible solutions. Nevertheless, model based estimates of global temperature increase over the last 130 years are more easily reconciled with observations when the likely cooling effect of s,? lgh, t" 7prr7CT, 11 is +",." :. ^,. ^..... . .., I . . .'  qualitative support for an estimated range of clim'ate sensitivity consistent with that given in IPCC (1990) (Figure 16). I 4 <._. i_,___ .._.. ".. ..,., C I r i r V r ..., u.. LA *"""" L E. 4 Studies of patterns of change To better address the attribution problem, a nu[ l, ber of recent studies have compared observations [virh model predicted patterns of temperature change i n response to anthropogenic forcing. The argu[ l; enr underlying pattern based approaches is til>..

different forcing mechanisms (" causes") may ha,,. d i f f e r e n t p a t t e r n s of response (" eflectj..,, particularly if one considers the full three or four dimensional structure of the response, e. p.. temperature change as a function of latitude. longitude. height and time. Thus a good march between modelled and observed multi dimensional patterns of climate change would be dimcult TO achieve for "causes" other than those actually use(! in the model experiment. Several studies have compared observed patterns of temperature change with modei patterns of change', from simulations with changes in both greenhouse gases and anthropogenic sulphate aerosols. These comparisons hare been made at the Earth's surface and in vertical sections through the atmosphere. While there are concerns rearding the . relatively simple treatment of aerosol effects in model experiments. a n d . the neglect of other potentially significant contributions 'to the radiative forcing, all such pattern comparison studies show significant correspondence. ' between the observations and model predictions (an example is shown in Figure 17). Much of the model observed corrcspondencc .in these experi~ ncnts occurs at the largest spatial scales  for csamplc. temperature differences betwccll. hemispheres., land and ocean; or the troposphere 'and strt~ tosphcre. Model . predictions are more reliable at these spatial scales than at the regional scale. Increasing confidence in the identification of a human induced effect on climate comes primarily from such pattern based work. For those seasons during which aerosol effects should be most pronotu~ ced the pattern correspondence is generally higher than that achieved if model predictions are based on changes in greenhouse gases alone (Figure 1 T). As in the global mean studies, pattern oriented detection work relies on model estimates of internal natural variability as the primary yardstick for temperature patterns could be due to natural causes. Concerns remain regarding the reliability of this yardstick. ..

cictlckLiLLg b,, ilciiicr U ~J S ~K V ~U . c[ l; t[ lges 111 42 I .. 60's aws "  180" 135'W SO'W 45" W 0" 45'E 90'E 135'E 180'

W 60's 905 t" s ' 180" 135'W 90" W 45'W 0' 45'E 90'E 135'E 180" Figure 17: Annual mean near srlrjace air temperature changes (" 0 from eqtlilibrirlm response experiments uith an atmospheric GCM with a mixed layer ocean coupled to a tropospheric chemistry model. forced mith present day atmosplreric concentrations of COY (a) and by the combined e/ fects of present day CO, lecels and sulphw emissions (b). Obserced ternperailire changes from 1955P4 to 1975 94.  show1 in Figure 9. are qualitatirely more similar to the changes in the combined forcing experirnefzt than ir~ the CO, only experiment.

arcas o f q u a l i t a t i v e a g r e e m e n t b e t w e e n obscrvations and those model predictions that either i~ icludc aerosol effects or do not depend critically on their inclusion. As in the quantitative studies. one m u s t b e c a u t i o u s i n a s s e s s i n g consistency because the expected climate change signal due to human activities is still uncertain, and has changed as our ability to model the climate system has improved. In addition to the surface warming. the model and obsorved commonalities in which we have most confidence include stratospheric cooling. reduction in diurnal temperature range, sea level rise: high latitude precipitation increases .and water vapour and evaporation increase over tropical oceans. In summary. the most important results related to the issues of detection and attribution are:

0 The limited available evidence from prosy climate indicators suggests that the 20th century . global mean temperature is at least as ivarm as any other century since at least 1400 AD. Data prior to 1400 arc too sparse to allow the reliable estimation of global mean temperature (sec Section .C. 2).

0 Assessments of the statistical significance. of the observed global mean temperature trend over ,.

the iast century have used a variety of new estimates of natural internal and esternally forced variability. These are derived from instrumental data. palaeodata, simple and , complex climate modeis, a11~ siatisilci ihbuL; C: 3 fitted to observations. Most of these studies have detected a significant change. and show that the observed warming trend is unlikely to be entirely natural in origin.

43 ..

1 : . .. . . . . .. .. .. 0 More convincing recent evidence for the (compounded). For comparison the IS92a Scenario. attribution of a human effect on climate. is neglecting the effect of aerosols, is equivalent to a emerging from pattern based studies, in which compounded rate of increase varying from 0.77 to the modelled climate response to combined 0.84%/ yr during the 21st century. forcing by greenhouse gases and anthropogenic The projections of global mean temperattlre sulphate aerosols is compared with observed sea level changes do not come directly from coupleci geographical, seasonal and vertical patterns of atmosphere ocean models. Though these are t h e atmospheric temperature change. These studies most sophisticated tools available for making s h o w t h a t s u c h p a t t e r n c o r r e s p o n d e n c e s projections of future climate change they are increase with time, as one would expect as an computationally expensive, making i t unfeasible to anthropogenic signa! increases in strength. produce results based on a large number of Furthermore, the probability is very low that emission scenarios. In order to assess global these correspondences could occur by chance as temperature and sea level projections for the fuii a result of natural internal variability only. The range of Is92 emission scenarios. simple upnelling vertical patterns of change are also inconsistent diffusion energ). balance models (see Box 2) can be with those expected for s d r and volcanic employed to interpolate and extrapolate the couplecl forcing. model results. These models, used for similar tasks in IPCC (1990) and IPCC (1992). are calibrated to give the same globally averaged temperature response as the coupled atmosphere ocean models. T h e c l i m a t e s i m u l a t i o n s h e r e are called projections instead of predictions to e q h a s i s e that they do not represent attempts to forecast the most likely (or "best estimate") evolution of climate i n the future. The projections are aimed at estimating and undcrstanding responses of the climate system to possible forcing scenarios.

I 0 Our ability to quantify the human influence on global climate is currently limited because the expected signal is still emerging from the noise of n a t u r a l v a r i a b i l i t y ; a n d b e c a u s e t h e r e a r e uncertainties in key factors. These include the magnitude and patterns of long term natural variability and the time evolving pattern of forcing by, a n d r e s p o n s f t o , c h a n g e s . i n concentrations of gree! house g a s e s a n d a e r o s o l s , a n d , land surface changes. I Nevertheless. the balance of evidence suggests that there is a discernible human influence on F* 2 Projections of climate change ,

global climate.

climate change emissions of both grecnho__; .......

F. l Forcing scenarios precursors (Section B. 9.11 projected global mean Projections of future anthropogenic climate change tempcrature challges relative to 1990 were depend; amongst other things, on the scenario used calculated for the 21st century. Temperature to force the 'climate model. The IS92 emission projections assuming the "best estimate" value of scenarios are used here for projections of changes in climate sensitivity, 2.5" C. (sce Section D. 2) are global mean temperature and sea level. The IS92 shown for the full set of IS92 scenarios in Figure 18. scenarios include emissions of both greenhouse. For IS92a the temperature increase by 2100 is 2° C. gases and aerosol precursors (see Section B. 9.1) and Taking account of the range in thc estimate of for the first time both factors have been taken into climate sensitivity (1.5 to 4.5'C) and the full set of account in, the global mean temperature and sea IS92 emission scenarios, the models project an level projections (Section F. 2). increase in global mean temperature of between 0.9 In many coupled model experiments the forcing and 3.5" c (Figure 19). In all cases the al'PralnP raw scenario is simplified by summing the radiative of warming would probably be greater than an) forcings of all the trace gescs (co,, CU;, 0,. ctc.) c:; d scfn in the !as: IC. 3CO cars. hi;: iht: ;Cts; ii ztiiiiu'.:

treating the total forcing as if it came from a n to decadal changes would include considerable "equivalent" concentration of c02. The rate of natural variabilitx. Because of the thermal inertia of increase of "equivalent CO," in these experiments is the oceans. global mean temperarure ~r o u l d o f t e n a s s u m e d t o b e a constant +10/ o/ vr c o n t i n u e t o . i n c r e a s e b e v o n d 2100 even. if .. .. concentrations of greenhouse gases were stabilised a t t h a t t i m e . Only 50 90% of the eventual temperature changes are Tealised at the time of greenhouse gas' stabilisation. All scenarios show substantial climate warming, even when the. negative aerosol radiative forcing is accounted for. Although CO2 is the most important anthropogenic greenhouse gas, the other greenhouse gases contribute significantly (about 30%) to the projected global warming. To allow closer comparison with the projections presented in IPCC (1990) and IPCC (1992) and to illustrate the sensitivity of future global temperature to changes in aerosol' concentrations, the same series of calculations were performed neglecting future aerosol changes, i. e. aerosol concentrations were held constant at 1990 levels. These lead to higher projections of temperature change. Taking account of the range in the estimate of climate sensitivity and the full set of IS92 emission scenarios, the models project an increase in global mean temperature of between 0.8 and 4.5" C. For IS92a, assuming the "best estimate" of. climate sensitivity. the temperature increase by 2100 is 24° C. For comparison. the corresponding temperature increase for IS92a presented in IPCC (1992) was 2. S" C. The projections in IPCC (1990) were based OR an earlier set, of emission scenarios, the "best estimate" for the increase in global temperature by 2100 (relative to 1990) was 3.3" C.

E2.2 Global mean sea level response to . 1592 emissionscenarios .

Using the IS92 emission scenarios. including gre. enhouse .gas and ae'rosol precursors, projected global mean sea level increases relative to 1990 were calculated for the 21st century. Sea level projections assuming the, "best estimate" values for climate sensitivity and ice melt are shown for the full set of IS92 scenarios in Figure 20. For IS92a. the sea level rise by 2100 is 49 cm. For comparison. the "best estimate" of global sea level rise by 2100 given in IPCC (1990) was 66 cm. Also taking account of the ranges in the estimate of climate sensitivity and ice melt parameters, and the full set of 1592 ern ission scenarios, the models project'an increase in global mean sea level 01 betxyeen 13 anu 94 cm (Figure 21). During the first haif of the nesi century, the choice of emission scenario .has relatively little effect on the projected sea level rise due to the large thermal inertia of the ocean ice atmosphel e climate system. but has increasingly Year Figure 18: Projected global mean surface temperature changes from 1990 to 2100 for the full set of IS92 emission scenarios. A climate sensitivity of 2.5% is assumed.

..

"

20002020 2040 2060 2080 2100 . . Year Figure 19: Projected global mean surface temperature change extremes from 1990'to 2100. The highest temperature changes assume n climate sensitiuity of 4.5% 'and the IS92e emission scenario; the lowest a climate sensitivity of 1.5% a, nd the IS92c emission scenario and the mid range curves a climate Sensitivity of 2.5% and the IS92a Scenario. The solid curves include the effect of changing aerosol: the dashed curves assume aerosol emissions remain constant at their 1990 lereis.

' L 45 ... ... ...... .. .. Year 4 Figure 20: Projected global mean sea level rise from 1990 to .. 2100 for the fidl set of IS92 emission scenarios. A climate sensitirity of2.5'C and mid value ice melt parameters are assumed.

a, u) . L 40 I c Year In these projections, the combined contribLltioll of the Greenland and Antarctic ice sheers are projected to be relatively minor over the nest century. However, the possibility of large chal1gej i l l the volumes of these ice sheets (and. consequeiltl!., in sea level) cannot be ruled out, although tlIr likelihood is considered to be low. Changes in future sea level will not O C C ~I ~

uniformly around the globe. Recent couplec\ atmosphere ocean model esperimeuts suggest thal the regional responses could differ substantially. o\ ving to regional differences in heating ant!

circulation changes. In addition, geological ~I I C ~

geophysical processes cause vertical lancl movements and thus affect relative sen` levels 011 local and regional scales. Tides. waves and storm surges couId be ilffected by regional climate changes. but future projections are. at present, highly uncertain.

152.3 Temperature and sea level projections compared with /K C (7990) The global average tempcraturc! aiid sea level projections presented here for 1990 to 2100. both e s c l u d i n g a n d i n c l u d i n g c h a n g i n g a e r o s o l emissions, are lower than the corresponding projections presented in IPCC 11990). Taking into account the negative radiative forcing of aerosols reduces projections of temperature and sca level rise. T ~O S C projections which e s c l d e the ct't'ectof changing aerosol emissions are lonw than IPCC (1990) for a number of reasons, mainly:

0 The revised (IS921 cmission scenarios have been u s e d f o r all grccnhousc gases. This is particularly important for CO, a n d CFCs.

Revised treatment of thc carbon cvcle. The Figure 21: Projected global mean sea leuel rise extremes from 1390 to 2100. The highest sea level rise curue assumes carbon cycle rnoclcl used to calculate future a climnte semiticity of 4.5%. high ice melt parameters and temperature and sea level rise in IPCC (19901 the IS92e emission scenario: the lowest a climate sensitivit. v and II'CC (1992) did not incorporate the effect of of 1.5%. low ice melt parameters and the IS9Zc emission carbon uptake . through C 0 2 fertilisation. scenario nnd the middle curues a climate sensitivity of2.5. c. ,resulting in highkr futurc CO, concentrations for mid rnlrte ice melt parameters and the IS92a Scenario. given emissions in IPCC (1990).

larger effects in the latter part of the nest century. In addition. because of the thermal inertia of the .. oceans, sea level would continue to rise for many centuries bcyo1: cl 2100 even if concrntrations of greenhouse gases were stabilised at that time. The projected rise in sea level is primarily due to thermal espansion as the ocean Ivaters tvarrn. but also due to increased melting of glaciers. 0 The inclusion of aerosol effects in the pre 1990 radiative forcing history. The estimated. historical thanncs of radintivo fnrcinr up to 1990. used in this report for global mean .a component due to aerosols. This particularly affects projections of sea level rise. tvhich are strongly influenced by the history of radiative forcing ocer the last century. ternpcrature ?.: id S E ~ I c \~; ~FU; L~ 'L;,,;~>. iii<!:~<!;< 0 Revised (and more realistic) parameters in the simple upwelling diffusion energy balance c l i m a t e m o d e l .  4 0 The inclusion in the model of spatial variations in the climate sensitivity and the effect of changing strength ,of t h e t h e r m o h a l i n e circulation, to accord with coupled atmosphere ocean general circulation models.

@ Thc usc of improved models for the ice melt component of sea level rise.

F. 3 Spatial patterns of projected climate change Although in global mean terms, the effect of including aerosols is to reduce the projected warming (scc Section F. 2). it can be rnislcading to consider only the global mean surface temperature. which does not give an effective indication of climate .change at smaller spatial scalcs. Because aerosols arc short lived. they are uncvcnly distributed across the globe, being concentrated ncar regions where they are emitted. As a result. thc spatial pattern of aerosol forcing is ve& different 'to that produced by thc long lived w +m i s e d g r e c n h o u s e g a s e s 2 n d . w h c n simulations that include the effects of both aerosols and co2, neither of which have yet been thoroughly analysed. LVe have concentrated on those changes which show most consistency between models, and for which plausible physical mechanisms have been identified. .

Temperature and Precipitation All model simulations. whether they are forced with increased concentrations of greenhouse gases and aerosols, or with increased greenhouse . gas concentrations alone, show the following features:

0 generally greater surface warming of the land than of the oceans in winter, as in equilibrium simulations (Figures 22 and 231;

*. a minimum warming around Antarctica and in &the northern North Atlantic which is associated with deep oceanic mixing in those areas:

0 maximum warming in high northern latitudes in late autumn and winter associated with reduced sea ice and snow cover:

0 little warming over the 'Arctic in summer:

little seasonal variation. of the warming in low  latitudes or over the southern circumpolar ocean: considering pattcrns of climntc? change. thcir cooling , ' .. 0 a reduction in diurnal temperature range ober effcct is not a sirllplc offsct to the warming effcct of greenhouse gases. as might be implied from the global mcarresults. Aerosols are likcly to have a significant offcct on future regional climntc changc. ,

Confidcncc is highcr in hcmisphcric to continental scalc projnctiorts o f climate changc (Section F. 3.1) than at regional scalcs (Section F. 3.2). whcre confidcncc remains low.

E3.1 Continental scale patterns I n IPCC (1990). estimrttes of the patterns of future climntc changc were prescnted. the most robust of which relatcd .to continental and larger spatial scalcs. The results xere based on CCM experiments .which included the effect of greenhouse gases, but did not take into account the effects of aerosols. The following provides some details of the changcs on continental scales in experiments with 'greenhouse gases alone (generally represented by a , l'ju, sr ir, c: erh;: i n U?,) and increases in Freenhouse gas and aerosol conc. entrations (using aerosol concentration derived from the IS92a Scenario). It is important to realise that, in contrast to the many model results with COz alone, there are only two recent coupled atmosphere ocean model , land in most seasons and most regions;

a n enhanced global mean hydrological cycle:

i incremed precipitation in high latitudes in . winter. ,

Including the effects of aerosols in simulations of future climate leads to a somewhat reduced surface warming, mainly in middle latitudes of the Northern Hemisphere. The maximum winter warming in high northern latitudes is less extensive (compare _Figures 22 and 23). Hon. ever. adding the cooling effect of aerosols is not a simple offset' t o . t h e w a r m i n g effect of greenhouse gases, but. significantly affects some of the continental scale patterns of climate change.. This is most noticeable in summer where the cooling due to aerosols tends to weaken monsoon circulations. For esample. when the effects of both greenhouse gases and aerosols are included, Asian 'summer monsuon m i d d l L! C: XX~: ~.+~F R S i n earlier simulations with 'only' the effect of '.

greenhouse gases represented, Asian summer monsoon rainfall increased. Conversely. 'the addition of aerosol effects leads to a n i n c r e a s e i n 1 \ x.

4T Figure 22; The patterrl of surface temperature change projected at the time of CO, doubiiny frorll a transient corrpled model e. rperiment.

180" 90'W 0' 90'E , 180'

6 4 2 0 2 4 6 3 Figure 23: The pntterrl oJsurjiace temperature change projrcted by a transient corlpled rnodrl at the time of CO, donbliny rr'tm both CO, and nerosol  concentration increases are taken into account.

6 4 2 2 ' 4 6 Precipitation over southern' Europe, whereas decreases are found in simulations with greenhouse gases only. These changes will be sensitive to the aerosol scenario used, and the details of the parametrization of the radiative effects of aerosol. Other forcings. including that due to increases in tropospheric ozone, soot and the indirect effect of sulphate aerosols have been neglected and could influence these results. In general, regional projections are also sensitive to model resolution and are affected by large natural variability. Hence confidence in regional projections remains low. With increases in C02 only. two coupled atmosphere ocean models show a pattern of SST (sea surface temperature) chanp, precipitation change and anomalies in kind and ocean currents that resemble the warm phase of ENSO, a s y e l l as the obscrw! decadal t i n s c a l z SST anonmlics of the 1980s and early 1990s. This is characterised by a reduction of the. east west SST gradient in the tropica1, Pacific. though the magnitude of this effect varies among models. Soil moisture Although there is less confidence in simulated changes in soil moisture than in those of temperature, some of the 'results concerning soil moisture are dictated 'more by changes in precipitation and evaporation than by the detailed response .of the surface scheme of the climate model. All model .simulations, i h e t h e r they are forced with increased concentrations of greenhouse gases and' aerosols. or with increased greenhouse gas concentrations alone. produce predominantly increased soil moisture in high northern latitucles in winter. Over the northern continents in summer. the changes in soil moisture are sensitive to the inclusion of aerosol effects. . ,

Oceancirculation ' .

In resp0: ise io iiicrcnsiIlg grceniiuuse gases, I I I U >L models show a decrease in the strength of the northern North Atlantic oceanic circulation further reducing the strensh of the warming around the North Atlantic. The increase in precipitation in high latitudes decreases the surface salinity, inhibiting 0 Land use changes are also believed to have a the sinking of water at high latitude, which drives significant impact on t e m p e r a t u r e a n d t h i s c i r c u l a t i o n .  precipitation changes, especially in the tropics and subtropics. Climate model experiments have E3.2 Regional scale patterns  shown the likelihood of substantial local climate Estimation of the potential impacts of climate change associated with deforestation in the change on human infrastructure and natural Amazon, or desertification in the Sahel. Changes ecosystems requires projections of future climate in land use on small scales which cannot be foreseen are expected to continue to influence regional climate. or continental means. Since IPCC (1990). a greater appreciation has Because of these problems, no information on been developed of the uncertainties in making future regional climate change is presented here. projections at the regional scale. There are several However. this situation is espected to in1pror. e in the difficulties: future as a result ,of:

0 The global climate models used for future 0 more coupled atmosphere ocean model projections are run at fairly coarse resolution Aperirnents with aerosol effects included; and do not adequately depict many geographic features ' (such as coastlines. lakes and m o u n t a i n s ). s u r f a c e v e g e t a t i o n , a n d t h e interactions betxveen the atmosphere with the surface which become more important on 0 more refined scenarios for .aerosols and other regional scales. Considerable spread esists among model projections on the regional scale even when climate modcl espcrimcnts are E3.3 Changes in variability and extremes driven by the same future radiative forcing Small changes in the mean climate or 'climate tariability can produce relatively large changes in the frequency of cstrcrnc events (defined as events where a certain threshold is surpasscdl: it small change in the variability has a stronger effect than a similar change in the mean. improvements in models, both from increased resolution and improved representation of small scale processes:

There is much more natural variation in local c l i m a t e t h a n in cliniatc averaged over continental or larger scales. This variation arises interactions benveen the atmosphere and the. Temperature oceans (such as EXSO). and fro111 variations in A generill warming tcrlds to lead to a n inrrp; lsc in^ Soil moisture. Sea ice. and other COmpOnentS of high tempcr; ltllrc events allcl a decrease the climate system. Series or ensembles Of model in lolv temperatures (e. g., frost days). "

predictions started from different initial . from locally generated variability from storms. .

. superimposed variability to be determined.

8 Because of their uneven spatial distribution.

1 human induced tropospheric aerosols arc likely to^ greatly influence future regional climate change. At present. howver. there are very few projections of climate change with coupled atmosphere ocean models (the type of modcl that gives more reliable information on the effects of aeroso1s: Those that have been run . gas concentrations. In some, areas a numb& of simulations show there is also an increase in the probability of dry days and the length of dry spells (consecutive days without precipitation). Wllere mean precipitation decreases, the likelihood of drought increases. Xew results reinforce the view !::?.;! n$ r~.! c;: r! n translates into prospects for more severe droughts and/ or floods in some places and :c;::~.:>.! T::.!.) lq: f, irh inrl~ qrip the radiative aerosol effects: Year Figure 21: The gloilal mean surface temperature response to the CO, CoIrcentration pathuays leading to stabilisation at 450 (dashed curves) and 650 (solid crrrces) pprnu (see Figure 7a) for a climate sensiticily of 1.3. 2.5 and 4.5%. The changes shown are those nrising from CO, increases alone. The date of concentration stabilisation is indicated by the dot. Calculations assume the "obserued"% istory of forcing to 1930, including aerosol efleects and then C02 ronwntrotion increases only beyond 1990.

350 300 250 Y 6 a, F 200 m c 0 p 150 a, m

c% 100 50 0 2.01 Year Figrtre 25: The global'mean sea level response to the CO, concentration pothruays leading to stabilisntion at 450 (dashed C i l t ' C L ' s / c t r i i LS; i iauiiu' cwcrs) p p /m isre i; gi'r< i u j j i r ti c i i m (m sensitiritg of 1.5. 2.5 and 4.5%. The changes s h o m are those arisiny frour CO, irrcrenses alone. The date of concentration stnbi[ isation is indicated by the dot. Cnlculafiorls assunze the %bsersed" lrisrory of forcing to 1990, including aerosol eflects and then CO2 com eutrution increases only beyond 1990. Mid latitude storms In the felv analyses available, there is little agreement between models on the changes in storminess that might occur in a warmer tvorlcl. Conclusions regarding estreme storm events are '

obviously even more uncertain.

Hurricanes/ Tropical cyclones The formation of tropical cyclones depends not o! ll!. on sea surface temperature (SST). but also 011 a number of amospheric factors. Although some models non. represent tropical s t o m s with so111e realism for present day climate, thc state of rile science does not allow assessment of future changes.

El Nifio Southern Oscillation Several global coupled nlodels indicatc that the ENSO likc SST variability thcy simulatc continues with increased CO,. Associated with the mean increase of tropical SSTs as a result of increased greenhouse gzs conccntrations. there could be enhanced .precipitation variability gssociatccl Lcith ENS0 evcnts in the increased COz climate. especially over rhc tropical continents.

F. 4 Effects of stabilising greenhouse gas concentrations Possiblc global temperature and sca Icvrl rcsponse to the scenarios for s$ lbilising conccntrations discusscd i n Section R. 9.2 wcrc calcuiatod with rhe same upndling diffusion cnergy balance model used for the rcstrlts in Sections F. Z. 1 and F. 2.2. . ,

.For cach of the pathways lcacling to stabilisntion. the climate system shows consiclcrablc warming d u r i n g t h e 21st century. Figure 24 shoivs ternpcrature increases for thc cases which stabilise at concentrations of 650 and 450 ppmv for different clinlatc sensirivitics. Stabilisation of the concentration does not lead to an imrnediate stabilisation of rhe global mean temperature. The global nlean temperature is seen to continue rising for hundreds of\ ears after thc concentrations have stabiiiscd in Figure 24 due to long time scales in the ocean. As sh0u. n in Figure 25. the long term sea level i'iaa "C U l l l l l l l l l l .C ,,i I:, cLal1 111urs pL'u" ull11~~ u. 3 e c l level continues to rise. at only a slowly dcclininr rate, for many centuries after greenhouse gas ConcentrationS and temperatures ha5. e stnbiiised.

. .. F. 5 The possibility of surprises Unexpected external influences. such as volcanic , eruptions, can lead to unexpected and relatively sudden shifts in the climatic .state. Also, as the response of the climate system to various forcings can be non linear. its response to gradual forcing changes may b e quite irregular. Abrupt and significant changcs in the atmospheric circulation involving the Nor: h Pacific lvhich began about 1976 were described i n I K C (1990). .A re. lated esample is the apparent fluctuation in the recent behaviour of ENSO. with warm conditions prevailing since 1989. a pattern which 112s been unusual compared to previous ENSO bclhnviour. Another esample is the possibility that the LVest Antarctic ice shect might "surge". causing a rapid rise in sea level. The current lack of knowledge regarding the spwilic circumstances under which this might occur. cirhcr in total or in part. limits the ability to quantify this risk. Nonethelcss. the likelihood of a major sea level rise by the year 2 100 due to thc collapse of the West Antarctic icc sheet is considcrcd low. In tht: occans tho meridional overturning might Lvcakcn in a futurc climatc. This overturning (tho thcrmohalinc circnlntion) is driven in part by dcop convection in the northern North Atlantic Ocean and kccps the northern Sorth Atlilntic Occan several clcgrcc? s ~vrnrmcr than it would othorwisc be. Both tho study uf palacoclimate' from scclimcnt records and ice CDITS anclmoclclling studies with couplod climate rnivcEcls and ocean GCbls can be interpreted to suggest that the ocean circulation has been very different in the past. Both in these observations and in the ocean models. transitions between different types of circulation seem to occur on a time scale of a few decades, so relatively sudden changes in the regional (Sorth Atlantic, Lt'estern Europe) climate could occur, presumably mainly in response to precipitation and runoff changes gvhich alter the salinity. and thus the density. of the upper layers of the Sorth Atlantic. LVhether or not such it suclden change can actually be realisecl in response to global warming and how strong a perturbation is required to cause a transition bet! veen types of circulation are still the subject of much debate. In terrestrial ecological systems. there are Jhresholds in the sustained temperature and water availability at which one biological population is replaced by another. Sonic replacement. e. g., in tree species, is slow whilc some. e g . in micro organisms is rapid. Minimum temperatures exist for thc survival of organisms in winter. and the populations of such organisms may move polclvards as the climntc and especially night time temperatures warm. If the transitions are not orderly. sudclen shifts in ccosystcm functioning will occur. These may have impacts of direct human rclevancc (as cliscusscd in IPCC \V G I I (1995)) but may also havc surprising impacts on climate via effects on albcdd, acrosol forcing. the hydrological c dc. cvapotnu1spir; ttion. C02 rclcasc and methane cycling. for csamplo (see Sections 11.1 and D. 2). G. Advancing our Understanding An important long term goal is the accurate projection of regional climate change, so that  potential impacts can he adequately assessed. Progress towards this objective depends on determining the likely global magnitude and rate of human induced climate change. including sea level change. as well as the regional expressions of these quantities. The detection and attribution of human induced climate change is also most important. To achieve these objectives requires systemtic and sustained global observations of relevant variables. as well as requiring the effective co operaiion and participation of .many nations. The most urgent. scientific problems requiring attention concern:

(i) t h e rate and magnitudegf climate change and sea level rise:

0 the factors controlling the distribution o l clouds and their radiative characteristics; '

0 the hydrological cycle, including precipitation. evaporation and runoff;

0 the distribution and time evolution of ozone and aerosols and their radiative characteristics: .. 0 the response 'of terrestrial and marinc systems to . climate change and their positive and, negative fccdbacks;

t h e r e s p o n s e o f ' ice sheets anci glaciers to climate:

tho influence of human activities on cmissions:

the coupling between the atmosphere and ocean, and ocean circulation;

t h e f a c t o r s c o n t r o l l i n g t h e a t m o s p h e r i c concentrations of carbon dioxide a n d . o t h e r greenhouse gases: (i i ) detection and attribution of climate change:

0 systematic observations of key variables. and development of model diagnostics relating to climate change:

pataeoclirnatic time series to describe natura] variability of the climate system: 0 relevant p r "r y data to construct arid (i i i ) regional patterns of climate change:

land surface processes and their link to atmospheric processes;

@ coupling of scales between global climate modeis and regional and smaller scale models;

0 simulations with higher resolution climate models.

The research activities for each objective are strongly interconnected. Such research is and needs to be conducted by individual investigators in a variety of institutions. .as well as by co ordinated international efforts which pool natioEal resources and talents in order to more efficiently engage in large scale integrated field a n d m o d e l l i n g programmes to ndvancc our understanding.

.. . .. . .. . .. Glossary Aerosols

Climate change (FCCC usage)

Climate change (IPCC usage)

Climate sensitivity Diurnal temperature rangc Equilibrium climate experiment Equivalent C02 Evapotranspiration Greenhouse gas Airborne particles. The term has also come to be associated, erroneously, with the. propellant used in. "aerosol sprays'.

.A change of climate which is attributed directly or indirectly to human activity t h a t alters the composition of the global atmosphere ancl which is in addition to natural climate variability observed over comparable time periods.

Climate change as referred to i:: the observational record of climate occurs because of internal changes within the climate system or in the interaction betlveen its components, or because of changes in esternal forcing either for natural reasons or because of human activities. It is generally not possible clearly to make attribution between these causes. Projections of future climate change reported by IPCC g e d a l l y consider only the influence on climate of anthropogenic increases in greenhouse gases and other human related factors.

InlpCC reports, climate sensitivity usually refers to the long term (equilibrium) change in global m a n surface temperature following a doubling of atmospheric C02 (or equivalent C021 concentration. More generally, it refers to the equilibrium change in surface air .  temperature following a unit change i n radiative forcing (" C/ tVm~).

The difference between masimuln and minimuni. tcmpcraturo oycr a period of 21 hours.

411 cspcrimcnt where a step change is applied to the forcing of a climate model and the model is then allowed to redch a new equilibrium. Such esperimcnts proviclc information on thc difference between the initial and final states of the model. bu t not on the time dependent response.

The concentration of CO, that would cause 'the same amount of ,

radiative forcing as the given misture ofCO, and othcr grct: nhousc gases. ..

The combined process of evaporation from the Earth's surface and transpiration from vegetation.

A gas that absorbs radiation at specific wavelengths within the spectrum of radiation (infrared radiation) emitted by the Earth's surface and by clouds. The gas in turn emits infrared radiation from a level where the temperature is colder than the surface. Tilt: net effect is a local trapping of part o i the absorbed rnergy and a .

tendency to warm the planetary surface. LVater vapour (H, O), carbon dioxide (CO,). nitrous oside (S20), methane (CIf,) and ozone (0,) are'the primary greenhouse gases in the Earth's atmosphere.

5 3 .. . ,.. . . .. CLIMATE CHANC . 795  The Science of Climate Change ,I' ~

Ice cap A dome shaped glacier usually covering a highland near a !yarer divide.

 Ice sheet A g l a c i e r m o r e t h a n 50,000 Lrnz in a r e a f o n n i n g a continuous covey Over a land surface or resting on a continental shelf. .

Radiative forcing A simple measure of the importance of a potential climate change mechanism. Radiative forcing is the perturbation to the e n e y s  balance of the Earth atmosphere system (i n \\ 'm ') follo\ ving. for esample. a change in the concentration of carbon dioside or a change in the output of the Sun: the climare system responds to the racliative forcing so a s to re establish the energy balance. A p0sirii. e radiative forcing tends to warm the surface and a negative radiati\ e forcing tends to cool the surface. The radiative forcing is normally quoted as a global and annual mean value. .4 more precise definition of radiative forcing. as usecl in IPCC reports. is the perturbation of ,the energy balance of the surface troposphere system. after allowing for the stratosphere to re adjust to a stale of global mean radiative equilibrium (see Chapter 4 of IPCC (1994)). Sometimcs called "climate forcing".

Continental: 10  100 million square kilometrt! s (kmz) Regional: 100 thousand  10 million k m l ,

Local: less than 100 thousand km?  Spatial scales 3 Soil rnoislure  Wltcr storod in or a t the continclital surfwe and available for ctapor: ition. I n II'CC . (1990) a single storc (or "buc: kct") \vas comnionly uscd, in climate ~; odcls. Toclay's models which incorporafr! canopy and soil proccsscs vicw soil moisture a s t h e amount lwld i n cscoss of plant "wilting point".

Slralosp11cre Tho highly s t r a t i f i e d a n d s t a b l c region o f tho a t n w s p h c r e a b o v e t h e troposphcrc (qr. 1 extcnding from &ut 10 km to about 50 km.

Thermol~ alinc circulation h r g o scale density driven circulation i n the oceans. driven b> diffcrencc: s in temperature and salinity.

Transient clirnatc experiment An analysis of the time depenclent response of a climate model to a time varying change of forcing.

Troposphere T h e l o w s t p a r t o f t h e a t m o s p h e r e from the surface to about 10 km in altitudc in mid latitudes [ranging from 9 kn1 i n high latitudes to 16 km in the tropics on average) where. cloucls and "weather" phenomena occur. The troposphere is cieiineti as he region b b i w i tempcraturw gencrally decrease [virh i l p i y l r . .x /'

I .

References IPCC, (IntergoTernDen'tal Panel on'climate Change) 1990: Climate Change: The IPCC Scientific Assessment. J. T. Houghton, G. J. Jenkins and J. J.. Ephraums (eds.), Cambric& Urliversity Press, Cambridge, Tu'#, 365 pp.

IPCC, 1992: Climute Chctnge 1992: The Supplementary Report to the IPCC Scientific Assessment, J. T. Houghton, B. A. Callander and S. K. Varney (eds.), Cambridge University Press, Cambridge, UK. 198 pp.

IPCC, 1994: Climate Change 1994: Radiative Forcing of Climate Change and an EvalLlation of the IPCC 7.932 Elnission Scenarios, J. T. Houghton, L. G. Meira Filho, J.

Bruce, Iioesung Lee, B. A. Callander, E. Haites, N. Harris and K. Maskell (eds.), Cambridge University Press, C& bridge, UK, 339 PP. 3 IPCC WGII, 1995: Climate Change 1995  Impacts. Adaptations clnd Mitigations oj'

Climatt, Change: Scientific Technical Analyses: Contribrltion of Working Group 11 10 the Second Assessment Report of the Intergovernmental Panel o~< Climate Change, R. T. Watson, M. C. Zinyowcra and R. H. Moss (eds.), Cambridge University Press, New York, USA.. '

IPCC WGI, 1995: Climate Change 1995,  The. Science of Climate Change: Conlrihlion of Working Group I to the SecondAssessment Report of the Inter~ ooerrznlentctl Panel on Clilnnte Change. J. T. Houghton, L. G. Meira Filho, B. A. Callander, N. Ilarris, A. Kattcnberg and K. Maskell (eds.), Cambridge University Press, Cambridge, UK.

f Conference on Human Health and  Global Climate Change Summary of the Proceedings NationalScience Technology Council and the 'Institute of Medicinemational Academy of Sciences Written by Paul B. Phelps for the Institute of Medicine Valerie Setlow and Andrew Pope, Editors NATIONAL ACADEMY PRESS 'Washmgton, D. C. 1996 ..

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NATIONl ACADEMY PRESS 2101 Constitutioa. Avenue, N. W. * Washington, DC 20418 NOTICE: The conference that is the subject of this summary was approved by the Govemi Board of the National Research Council, whose members are drawn fiom the councils of the Natior Academy of Sciences, the National Academy of Engineering, and the Institute of Medicine. This report has been reviewed by a group other than the authors accordins to proredm approved by a Report Revie; r Committee consisting of members of t k Xziions: Ai; t& iiiY o i Science tile Wationai Academy of Engineering, and the Institute of Medicine. The Institute of Medicine was chartered in 1970 by the National Academy of Sciences to enli distinguished members of the appropriate professions in the examination of policy matters pertaining t the health of the public, In this, the Institute acts under both the Academy's 1863 congressional chartt responsibility to be an adviser tothefederal government and its own initiative, to identify issues c medical care, research, and education. Dr. Kenneth I. Shine is president of the Institute of Medicine, This summary was prepared by Paul Phelps for the Institute of Medicine. It summarizes th presentations and discussions that occurred during a 2 day conference (September 11 12, 1995) that wa organized and conducted in a collaborative effort between the Institute of Medicine, the Nationa Academy of Sciences, and the Naunal Science and Technology Council (NSTC), with support fron several member agencies of the NSTC. The views summarized in this report are those of the Conference participants and donot represent the views of the NSTC or the Institute of Medicine and Nationa. Academy of Sciences. This summary was reviewed for accuracy by the chairs of the individual sessioE of the, conference and by the chairs of the breakout group panels. Funding for the conference was provided by the National Academy of Sciences, the Office of Science and Technology Policy, the National Science Foundation, the National Aeronaiitics and Space Administration, the National Oceanic and Atmospheric Administration, the EnvironmentaI Protection Agency, the Department of Defense, the National Institutes of Health's Fogarty. Intemtiona1 Center and the National Institute of Environmental Health Sciences, the Department of Agriculture, the Centers for Disease Control and Prevention; the Agency for International Development, and the Department of Energy.

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