1

Report on Forcings for the C20C and EMULATE HadAM3 Experiments

Jeff Knight

10th March 2003

Introduction

This report briefly outlines the externally applied climate forcing used in the atmospheric generalcirculation model experiments for the Climate of the Twentieth Century (C20C) and European andNorth Atlantic Daily to Multideadal Climate Variability (EMULATE) projects. The different forcingmechanisms are split into two groups, natural and anthropogenic. In the natural group there isatmospheric forcing from sea-surface temperature (SST) and sea-ice anomalies, Milankovitch cycles,variations in the total solar irradiance and stratospheric volcanic aerosols. In the anthropogenic groupthere is changing atmospheric concentrations of greenhouse gases, changes in tropospheric andstratospheric ozone, the direct and indirect effects of atmospheric sulphate aerosols and changes in landsurface characteristics. The division between these two groups is slightly blurred due to the possibilitythat some of the changes in SSTs and sea-ice in recent decades may have arisen due to anthropogenicclimate change. Nevertheless, taking these two groups in turn, a description of how each of the forcingsis applied in the experiments is given.

Natural Forcings

(a) Milankovitch forcing.

HadAM3 calculates the solar flux and declination based on the parabolic orbital elements of the Earthat each time step. In addition, the model has been modified to include the principal long-term variationsin terrestrial rotation and orbit following Berger (1979). As the mean distance between the Earth andSun does not change, these do not affect the average global solar heating, but do alter the distributionof solar radiation as a function of latitude and season. Famously proposed by Milankovitch, these arechanges in the earth's obliquity, orbital eccentricity and the date of perihelion. The variation of theobliquity, or the inclination of the orbital axis to the plane of the earth's orbit, is shown in fig. 1 (toppanel). The obliquity shows near-periodic variations between about 22.3º and 24.3º with a time scale ofabout 40,000 years. This modulates the strength of the seasonal cycle, which is itself a result of theobliquity. Fig. 1 also shows the eccentricity, a measure of the deviation from circularity of the earth'sorbit, has declined significantly over the past 100,000 years, from 0.04 to its current value of about0.017 (middle panel). In the current epoch, the distance between the Earth and Sun at perihelion(closest approach) is about 3.4% less than that at aphelion (farthest), which corresponds to a differencein solar flux at the Earth of about 6.8%. With greater eccentricity, this difference was much larger inthe past. The last effect is the change of the date of perihelion through the year. Fig. 1 (bottom panel)shows that it shifts through the calendar year at an approximately regular rate, taking about 21,000years to cycle through the seasons. Mostly, this change arises through luni-solar precession, or the slowrotation of the Earth's axis about a fixed direction in space due to the combined gravitational influenceof the Moon and Sun. Thus the fixed seasons progressively occur in different parts of the Earth's orbit,giving the potential for slight reinforcement or damping of the seasonal cycle. In addition, there is asomewhat smaller contribution from a steady change in the alignment of the orbit itself due toplanetary perturbations. Currently, perihelion is in early January, and is slowly (25 minutes per year)becoming later in the year. With this being in the Northern winter/Southern summer, and aphelionoccurring in the Northern summer/Southern winter, the Southern seasons are amplified and theNorthern seasons damped.

The overall change in the seasonal distribution of radiation at the top of the atmosphere due toMilankovitch cycles can be very marked on paleaoclimate time scales, but is much less so on timescales associated with historical climate such as studied in C20C and EMULATE. Nevertheless, thesemay have had a discernible effect regionally, if not on the global mean climate. Fig. 2 shows thecalculated diurnal mean flux at the top of the atmosphere for each month of the year over the historical

2

climate time frame. The dominant effect over this period is the shift of the date of perihelion. Noticethat there is a relatively small effect in January and June, when the Earth is near peri- and ap-helion.This is not surprising as the Earth-Sun distance is nearly stationary at these times, and so does notchange very much as the perihelion advances. The principal changes are found in a fairly generalincrease in the flux in the early part of the year (Feb-Apr), with declines in the mid- and high-latitudeNorthern late summer (Jul-Aug) and early Southern summer (Nov-Dec) and in Oct-Nov generally.With maximum changes of up to about 1 W m-2 at certain latitudes, Milankovitch forcing has beenincluded as it is a may cause some discernible change in the simulation of regional climate and theseasonal cycle and has very small uncertainty.

(b) Solar Forcing.

The local solar irradiance at the top of the atmosphere is clearly a function of the intrinsic luminosity ofthe Sun as well as the astronomical parameters relating to the orbit and rotation of the Earth. It is wellestablished that the luminosity has varied by a very small fraction over the course of recent solarcycles, as do other solar properties, such as sunspots and the extent of the corona. Direct quantitativemeasurements of solar radiation are relatively recent, however, so a longer term history of solarbrightness has to be derived from proxies. For C20C and EMULATE we use a reconstruction based onobserved numbers of sunspots which is an updated version of the dataset published by Lean et al.(1995). Total solar irradiance specified in the model is shown in fig. 3 (upper panel). This is based on amean value for 1748-1991 of 1365 W m-2. The most important feature is an increase of about 2 W m-2

occurring between 1900 and 1950. Solar cycle variability is also present, but shows marked differencesfrom cycle to cycle, implying another source of decadal-interdecadal variability. Using the fact that theEarth's cross-section is one quarter of its surface area, and a planetary albedo of 0.3 (i.e. 0.7absorption), the radiative forcing from solar changes can be estimated (fig. 3, lower panel). The early20th century increase is therefore about 0.3 W m-2, which although more modest than the peak forcingfrom the Milankovitch cycles (0.7 times the numbers in fig. 2) has the advantage of being global. Onthe other hand, the solar changes are quite uncertain, as they rely on extrapolating the luminosity-sunspot number relationship to the secular time scale. In any case, the total extra-terrestrial (i.e. solarand Milankovitch) forcing of climate is rather small in these simulations. Over and above simpleradiative forcing, however, there is the possibility that the more substantial ultraviolet solar variabilitycould have a significant effect on climate is through its influence of short-wave stratospheric ozone. Inthe model simulations a small change is made to the shape of the solar spectrum to accommodate this,although it is doubtful whether the model has the stratospheric resolution or domain to capture thiseffect. Added to this, the model has a prescribed ozone field and so only allows changes in radiativeheating to be computed and not the chemical changes that would also result.

(c) Volcanic forcing.

Explosive volcanic eruptions can inject millions of tonnes of sulphur dioxide into the stratosphere,which is oxidised over a few weeks to form a layer of micron-scale aerosol droplets. These act toreflect sunlight and cool the surface. Major eruptions, such as that of Mt. Pinatubo in 1992, create largenegative radiative forcings (> 2 W m-2) for 2-3 years after the eruption. In isolation, these eruptions areimportant for climate in the following years, but not for decadal climate variability. The historicalrecord, however, shows these eruptions appear to be grouped in time, and together might haveproduced a significant decadal signature in the climate record. In the model we use a dataset of 0.55 µmvolcanic aerosol optical depth from Crowley (2000) which is based primarily on Greenland andAntarctica ice-core sulphate and conductivity measurements, cross-referenced by a catalogue of knowneruptions. The model has been modified to adapt these optical depths to aerosol mixing ratios for themodel's radiation scheme. As the aerosol loading is not known very accurately, the forcing is applied asaverages over four latitude bands (90ºS-30ºS, 30ºS-0º, 0º-30ºN, 30ºN-90ºN) for each month of theintegration. These optical depths are shown in fig. 4, which illustrates the occurrence of periods ofstrong and weak volcanicity. For example, the aerosol loading between 1920 and 1960 is much less, onaverage, than the period since 1960.

(d) Sea-surface temperature (SST) and sea ice.

3

The model uses the HadISST1.1 dataset to prescribe SST and sea-ice values, and is interpolated daily.HadISST is a monthly globally complete dataset based on quality-controlled SST and sea-iceobservations and spans the period January 1870 to present. Data-sparse regions are filled using areduced-space optimal interpolation technique. The basic dataset is supplied on a 1º grid and isinterpolated to the model grid (3.75º by 2.5º) before the application of the Karl Taylor variancecorrection procedure. This ensures that when the model interpolates daily SST and sea-ice values fromthe monthly data, the resulting monthly mean values are the same as in the original dataset. Unlike theother natural forcings, SST and sea-ice forcing is complicated by the fact that the ocean has a long-termclimate 'memory', acting to integrate climate variability and change. Thus it is not just an expression ofthe forcing of the atmosphere by internal oceanic processes, but is also a reflection of variabilityinternal to the atmosphere, as well as all the other forcings, including anthropogenic ones. Thus theterm 'natural' must be treated with some caution in relation to SST and sea-ice.

Anthropogenic Forcings

(a) Well-mixed greenhouse gases.

In the C20C/EMULATE runs the history of the following well-mixed greenhouse gases are specified:

Year CO2

(ppmv)

CH4

(ppm)

N2O

(ppb)

CFC11

(ppt)

CFC12

(ppt)

CFC113(ppt)

HCFC22

(ppt)

HFC125(ppt)

HFC134A(ppt)

1749 0.0 0.0 0.0 0.0 0.0 0.01859 284.5 0.486 425.71875 287.11890 292.6 0.5241900 294.21917 300.51920 0.5711935 307.81940 0.6081950 309.1 440.7 0.0 0.01957 0.6601960 315.1 83.1 126.61965 447.11970 323.9 284.8 501.3 0.01980 337.0 821.3 0.01984 461.51990 349.1 0.951 471.0 1249.0 1993.0 498.1 272.0 0.0 0.01995 356.11998 1242.0 2225.0 536.92000 364.7 0.966 483.2 1190.0 2183.0 524.0 490.0 9.1 123.62010 385.7 1.008 496.9 976.9 1978.0 465.7 561.0 23.7 344.0

Entries in the table are mass mixing ratios (kg/kg) except CO2 which is a volume mixing ratio. Only theturning points specified in the model are shown. Values at intermediate times are interpolated betweenthese times.

(b) Ozone

The model treats both tropospheric and stratospheric ozone changes. In the troposphere, values aretaken from simulations using the Met Office tropospheric chemical transport model STOCHEM.Stratospheric values are set to a seasonal climatology before 1970 and linear trends from Randel (1998)are applied after 1975. Half trends are applied between 1970 and 1975. These use a combination ofSAGE satellite data above 20 km and between 56ºS and 56ºN, ozonesonde trends at high latitudes andtrends below 20 km that make the total column ozone change equivalent to that derived from TOMS.Zonal mean ozone fields are specified for each month of the simulation. The annual mean ozonerelative to 1970 is shown in fig. 5. Tropospheric ozone is typically 25% less than in 1970 at the start ofthe simulations, but grows to 10-25% more than in 1970 by the end. In the stratosphere, extratropicalozone declines rapidly after 1970, with greatest loss at high latitudes, especially in the Southern

4

Hemisphere (> 50%). The seasonal variation includes the 'Antarctic ozone hole' in the SouthernHemisphere spring, and spring ozone loss in the Northern Hemisphere, but there are also importantcontributions to ozone loss in other seasons (not shown).

(c) The direct effect of sulphate aerosols.

The model includes a representation of the direct (radiative impact of tropospheric sulphate aerosols) aspart of a full sulphate aerosol scheme. The input to this is the geographical distributions of sulphurdioxide emissions and a climatology of the principal oxidants of the gas. The model then treats theformation, transport and fate of sulphate aerosols and calculates their radiative effects directly. Fig. 6shows the time evolving pattern of near-surface sulphur dioxide emissions. Initially, small andrestricted to Western Europe and Eastern North America, these emissions grow and spread until about1975 when emissions are limited in these places. Developing world emissions, particularly in Asia,increase after this time. The equivalent pictures for high level (~ 930 hPa) emissions, reflecting thedevelopment of high chimney stacks, is shown in fig. 7. Emissions at this level first appear in 1975 andstrongly reflect the surface emissions. A seasonally varying three-dimensional climatology the sulphurdioxide oxidants OH, HO2, H2O2 and O3 is used in the calculation of aerosol formation. Fig. 8 showsthe zonal mean of these oxidants for January (top 4 panels) and July (bottom 4). Concentrations tend tobe highest in the tropics, but there is a clear asymmetry between the Northern and Southern summers,with much larger oxidant concentrations in the north.

(d) The indirect effect of sulphate aerosols.

The model has been modified to accept specified changes in the cloud albedo representing the changesto cloud microphysical properties by high aerosols concentrations. Three-dimensional changes areinterpolated in the model from data supplied on 5 yearly intervals. Selected zonal mean cloud albedoperturbations are shown in fig. 9. The effects are mainly associated with low- to mid-level cloud. Earlyin the calculation the effect is restricted to the Northern Mid-latitudes, and grows until about 1975.During the latter part of the calculation period the indirect aerosol effect forcing spreads equatorwardand into the Southern Hemisphere. The spatial distribution of albedo anomaly at 800 hPa is shown infig. 10. This shows the forcing concentrated over the continents.

(e) Land use changes.

A land use dataset developed at the Hadley Centre is used to force the model surface characteristics.Changes in land use can affect the surface albedo, surface friction and the hydrological cycle. A rangeof 10 different surface characteristics are prescribed from 10 yearly input data. Amongst these is thefraction of the model gridbox covered by vegetation. Changes in this relative to 1970 values are shownin fig. 11. The most significant effects are the clearing of Eastern North America and Eastern Europe inthe last half of the 19th century and early 20th century. One effect of these changes is shown in fig. 12,which shows the change in surface albedo when deep snow is present as a result of vegetation changes.Forests appear dark in snowy conditions, but the snow covered fields and urban areas are much morereflective, enhancing the albedo. The type, as well distribution, of vegetation is also has impacts. Fig.13 shows changes in the mean height of the canopy, and confirms the tendency for removal of forest inEastern North America and Eastern Europe. In addition, by 2000, deforestation in the Amazon regionis evident. Fig. 14 shows the change in the surface roughness length associated with these changes,demonstrating the reduction in atmospheric drag over these regions.

Figures

5

Figure 1. Milankovitch cycle forcing in the model. Changes in terrestrial obliquity (upper panel),orbital eccentricity (middle panel) and in the date of perihelion (lower panel) are incorporated. Time is

shown on the abscissa in thousands of years, with negative values indicating dates BC and positivevalues AD. The year 2000 is denoted by the solid vertical line in each case.

6

Figure 2. Milankovitch forcing at the top of the atmosphere. Diurnal mean radiative flux anomaly withrespect to a reference time of 1975 for the 15th day of each month of the year.

7

Figure 3. Time series of total solar irradiance applied in the model (top panel). Note values after 1998are set equal to 1998. The estimated radiative forcing with respect to the period 1961-90 (lower panel)

is estimated using geometric arguments and an albedo of 0.3.

8

Figure 4. Volcanic aerosol optical depth as a function of latitude and time.

9

Figure 5. Annual zonal mean ozone change relative to ozone amounts in 1970 for 1855, 1895, 1935(tow row), 1975, 1995 and 2005 (bottom row). Contour intervals are 1, 2, 5, 10, 20, 50 and 80% in bothpositive and negative senses. Values on the ordinate are pressure normalised by its value at the surface.

10

Figure 6. Global near-surface sulphur dioxide emissions (10-10 kg m-2 s-1) for 1850-2000.

Figure 7. Global high-level sulphur dioxide emissions (10-10 kg m-2 s-1) for 1850-2000.

11

Figure 8. Zonal mean distributions of the climatological SO2 oxidants OH, HO2, H2O2 and O3. Valuesfor January are shown in the upper four panels, and for July in the lower four panels.

12

Figure 9. Zonal mean albedo changes representing the indirect aerosol effect for 1825-2025.

Figure 10. Specified distribution of aerosol-related albedo change at 800 hPa between 1825 and 2025.

13

Figure 11. Change in vegetation fraction relative to 1970 for 1840-1990.

14

Figure 12. As fig. 11, but for surface albedo with deep snow cover.

15

Figure 13. As fig. 11, but for the canopy height of vegetation (m).

16

Figure 14. As fig. 11, but for roughness length (m).