volcanic and solar forcing of climate change during the...

4094 VOLUME 16J O U R N A L O F C L I M A T E

q 2003 American Meteorological Society

Volcanic and Solar Forcing of Climate Change during the Preindustrial Era

DREW T. SHINDELL AND GAVIN A. SCHMIDT

NASA Goddard Institute for Space Studies, and Center for Climate Systems Research, Columbia University,New York, New York

RON L. MILLER

NASA Goddard Institute for Space Studies, and Department of Applied Physics and Applied Mathematics, Columbia University,New York, New York

MICHAEL E. MANN

Department of Environmental Sciences, University of Virginia, Charlottesville, Virginia

(Manuscript received 2 December 2002, in final form 15 May 2003)

ABSTRACT

The climate response to variability in volcanic aerosols and solar irradiance, the primary forcings during thepreindustrial era, is examined in a stratosphere-resolving general circulation model. The best agreement withhistorical and proxy data is obtained using both forcings, each of which has a significant effect on global meantemperatures. However, their regional climate impacts in the Northern Hemisphere are quite different. Whilethe short-term continental winter warming response to volcanism is well known, it is shown that due to opposingdynamical and radiative effects, the long-term (decadal mean) regional response is not significant compared tounforced variability for either the winter or the annual average. In contrast, the long-term regional response tosolar forcing greatly exceeds unforced variability for both time averages, as the dynamical and radiative effectsreinforce one another, and produces climate anomalies similar to those seen during the Little Ice Age. Thus,long-term regional changes during the preindustrial appear to have been dominated by solar forcing.

1. Introduction

While climate models have been able to simulate ob-served global annual average temperature changes dur-ing the past 150 yr reasonably well, it is much moredifficult to reproduce continental-scale patterns of cli-mate variations (Houghton et al. 2001). Understandinglocal changes is even more important than understand-ing global change, as impacts will be felt on continentalor smaller scales (hereafter, we apply ‘‘regional’’ tomean continental in scale). Given the many factors play-ing a role in climate change since the industrial revo-lution, it is extremely difficult to unravel the regionalclimate response to particular forcings during this pe-riod. For the few centuries prior to industrial era, how-ever, externally driven climate change is thought to havebeen forced primarily by only two factors: variation insolar output and volcanic eruptions (Crowley 2000; Freeand Robock 1999; Shindell et al. 2001a). These forcingslikely played a large role in the so-called Medieval

Corresponding author address: Dr. Drew T. Shindell, NASA GISS,2880 Broadway, New York, NY 10025.E-mail: [email protected]

Warm Period (MWP) and Little Ice Age (LIA) epochsof the last millennium, which saw significant climatechanges on at least regional scales. Understanding themagnitude and causes of the forced climate variations,and distinguishing them from unforced, internal vari-ability, is important for historical purposes, and is acrucial test for climate models attempting to predict fu-ture climate variations.

Many studies have attempted to attribute preindustrialclimate change to external forcings, primarily using rel-atively simple models. Recently, for example, energybalance models have been used to study the global av-erage response (Crowley 2000; Free and Robock 1999;Marcus et al. 1999; Reid 1997). The studies of Free andRobock (1999) and Crowley (2000), which includedboth solar and volcanic forcings, suggest that volcanicforcing played the dominant role. Studies using generalcirculation models (GCMs) of various complexity,which allow the investigation of both the global andregional climate response to external forcing, have gen-erally been restricted to examinations of historical solarforcing (Cubasch et al. 1997; Drijfhout et al. 1999; Shin-dell et al. 2001a). Here we present GCM simulations ofthe decadal-to-centennial-scale climate response to vol-

15 DECEMBER 2003 4095S H I N D E L L E T A L .

FIG. 1. Time series of zonal mean stratospheric aerosol optical thickness at 550 nm used in the transient volcanic experiment. The latterportion of these data are derived from SAGE II instrument observations. The last 9 yr, followed by 3 yr with no aerosol loading, were usedto create the 12-yr series used to repeatedly force the GCM for the Pinatubo and Tambora experiments.

canic forcing, and compare this to the response to solarforcing and to internal, unforced variations.

2. Experimental setup

Simulations were performed using a version of theGoddard Institute for Space Studies (GISS) GCM con-taining a mixed layer ocean with fixed heat transportsand a detailed representation of the stratosphere. Themodel contains parameterized stratospheric ozone pho-tochemistry (Shindell et al. 1999), which includesozone-related heterogeneous chemistry. Though themodel has relatively coarse horizontal resolution (88 lat-itude 3 108 longitude), it can be run for multiple, longsimulations and it reproduces the observed recent mul-tidecadal stratospheric trends, which seem to be closelylinked to regional climate change (Shindell et al. 2001b).The mixed layer ocean attains thermal equilibrium aftera few decades, while the time scales in the real oceancan be much longer, and of course does not allow forocean circulation changes. These features will influencethe results, especially in the North Atlantic and theSouthern Ocean where heat is transported into the deepocean. Future simulations will include a full dynamicalocean.

We examine the generalized climate response to vol-canic eruptions and its dependence upon the size, timeof year, and frequency of the eruptions. We comparethis with the long-term response to solar forcing as ev-idenced by the changes between the late Maunder Min-imum, a period of very low solar irradiance during thelatter part of the seventeenth century, and a century later(Shindell et al. 2001a). This relatively large solar irra-diance change makes an excellent test case as it occurredprior to sizeable anthropogenic impacts. Furthermore,estimates of its magnitude and timing are documentedby both historical observations and cosmogenic proxies.Our previous simulations demonstrated that the model’sregional response to solar forcing largely reproduced

the solar component of the surface temperature changederived by correlating reconstructions of irradiance andsurface temperatures (Shindell et al. 2001a). We nowinvestigate how well we can account for the total re-gional pattern of surface temperature change evident inclimate proxy reconstructions of the LIA, focusing onthe Maunder Minimum period.

Since precise historical data on the spatial and tem-poral distribution of aerosols from volcanic eruptionsare unavailable, we base our simulations upon morerecent observed stratospheric aerosol optical properties(spatial and temporal distribution, effective radius, andoptical thickness) using the GISS dataset (Hansen et al.1996; Sato et al. 1993; Fig. 1). The vertical profile ofthe aerosols covers 15–35 km in 5-km steps and opticalproperties are described at several key wavelengths (fur-ther information is available online at http://www.giss.nasa.gov/data/strataer/). We have performedan ensemble of five simulations using the observed timeseries of 1959–99 eruptions, and a separate run re-peatedly simulating the June 1991 Mt. Pinatubo erup-tion. To see the effect of a larger aerosol load and dif-ferent eruption date relative to the seasonal cycle, wealso forced the GCM with two estimates of the April1815 Mt. Tambora eruption. For Pinatubo, aerosols for1991–97 are based upon observations from the Strato-spheric Aerosol and Gas Experiment (SAGE II) instru-ment. This period includes the Pinatubo eruption andthe subsequent 6 yr of declining aerosol amounts, towhich we add 2 yr assuming an exponential decreasein optical thickness with a 1-yr decay constant, and then3 yr with no aerosol loading to make a 12-yr eruption/decay time series. To approximate the Tambora eruption,we use the same time series, but shifted two monthsearlier and with the aerosol amounts increased by afactor of 2 or 3 relative to Pinatubo (the aerosol amountis far more important than the timing of the eruption),which we refer to hereafter as Tambora 2P and 3P, re-spectively. These aerosol increases are based upon es-


timates from ice core information calibrated against op-tical depth information for modern eruptions, whichsuggests that the forcing from Tambora was roughlytwice that of Pinatubo (Crowley 2000), along with dustveil indices, which suggest a factor of 3 between theeruptions (Robock 2000). These estimates should cap-ture the climatically relevant stratospheric aerosol moreaccurately than a volcanic explosivity index, which in-dicates the strength of the eruption rather than the aero-sol injected into the stratosphere. The three volcanicaerosol series were used to drive model simulations inwhich the eruption/decay cycle was repeated 10 times(120-yr simulations) for each case. This repetition al-lows us to obtain good statistics on the model’s responseas well as to examine the long-term response to periodiceruptions. For brevity, we emphasize the Tambora 3Prun over the 2P run, as the 2P response is largely amuted version of the 3P response. Results are comparedwith a multicentury (250 yr) control run without forcing.All runs began from stable initial conditions taken froman earlier control simulation.

3. Impact of volcanic eruptions

The global annual average surface temperature re-sponse to volcanic eruptions is cooling, resulting fromincreased absorption and reflection of incoming short-wave radiation by stratospheric aerosols. Averaging allyears of the simulations together, the mean annual av-erage cooling was 20.358C for the periodic Pinatuboeruption, 20.778C for the periodic Tambora 2P eruption,21.098C for the periodic Tambora 3P eruption, and20.448C for the observed 1959–99 volcanoes. Thesefollow the average instantaneous radiative forcings,which were 20.47, 20.91, 21.39, and 20.44 W m22,respectively, almost linearly, though there is a sugges-tion of slightly larger climate sensitivity in response tothe more frequent 1959–99 eruptions. (The sensitivitymight be reduced with the introduction of a dynamicalocean model that is free to transport heat anomaliesbeneath the mixed layer.)

a. Winter warming

In contrast to the long-term global cooling, large re-gions of the NH extratropical continents warm in thePinatubo simulation during winters following eruptions(Fig. 2, top). This phenomenon is well known, havingbeen seen in both observations (Kelly et al. 1996; Ro-bock 2000; Robock and Mao 1992) and GCM studies(Graf et al. 1994; Hansen et al. 1996; Kirchner et al.1999; Rozanov et al. 2002; Stenchikov et al. 2002). Thisarises because volcanic forcing induces a shift towardthe high phase of a hemispheric-scale dynamic circu-lation pattern known as the Arctic Oscillation (AO) ornorthern annular mode (which is analogous to the NorthAtlantic Oscillation in the Atlantic sector, where it ismost strongly expressed). During the cold season, this

results in enhanced westerly advection of relativelywarm oceanic air over the continents and of cooler airfrom continental interiors to their eastern coasts (Fig.2, top). In some regions, meridional winds are modu-lated along with the increased westerlies, leading to en-hanced warming over Siberia and cooling over the Mid-dle East, for example. Surface temperatures respond op-positely to a low phase, with reduced westerly flow,which allows air from the northeast to flow into Europe,leading to cooling there. We define the model’s AOpattern as the leading empirical orthogonal function(EOF) of November–April monthly mean sea level pres-sure (SLP) in the control run. The model’s control runAO accounts for 26% of the variance in the SLP, similarto the 23% seen in observations (Thompson and Wallace1998). SLP differences between the volcanic and controlsimulations (e.g., Fig. 2, bottom) were then projectedonto the AO to obtain the mean AO change, which wedefine as the opposite of the average SLP change pole-ward of 608N. Thus, an increased AO corresponds to adecrease in Arctic SLP, which is accompanied by anincrease in midlatitude SLP and a strengthening of thewesterly zonal winds around 508–708N. Note that wesaw no significant changes in AO variability in our ex-periments, and throughout this paper a change in theAO refers to a change in the mean state rather than inthe variability.

The average response for the 10 winters immediatelyfollowing the Pinatubo eruptions is an AO increase of1.8 6 0.9 mb at the 95% confidence level. While theresults are statistically significant, the AO response inindividual simulations ranged from 22.0 to 14.4 mb.Assuming this range is realistic, comparison with a sin-gle real-world realization (an observed AO increase of0.9 mb following Pinatubo relative to the previous 7 yr)is of limited value (the probability of a positive responseof any size is 85%, and the two standard deviation rangeof a single eruption is 21.2 to 14.8). Similarly, therange of responses seen in the second year followingthe Pinatubo eruption was quite large (21.5 to 11.8).The difference between the observed 1992/93 increaseof 1.8 mb in the AO and the model’s mean response ofno AO change is thus within expected interannual var-iability. The observed winter response averaged overseveral eruptions provides better statistics, but becausethe aerosols injected by each eruption are different, thecomparison is still imperfect. An analysis of pressureobservations during the winter following the four largesttropical eruptions of the last 150 yr shows a similarpattern to that seen in the GCM, with statistically sig-nificant SLP changes of 24 to 26 mb over the Arcticand 12 to 14 mb over Europe during winter (Kelly etal. 1996). Note that the model’s SLP changes are largerin winter (December–February) than in the broader No-vember–April average, with decreases of 3–7 mb overthe Arctic, in accord with the observations. Addition-ally, both the GCM and the observations show a de-crease in SLP of 21 to 23 mb over North America


FIG. 2. (top) Average winter surface temperature response (Dec–Feb) and (bottom) SLP (Nov–Apr) in the 10 winters immediately followingeruptions for Pinatubo. Temperature responses are significant in the subtropics for values greater than about 0.3 K, and at midlatitudes forvalues greater than about 1.0 K, while the response poleward of ;758N is not significant.

(statistically significant in the observations, marginallyso in the GCM for December–February), breaking thezonal symmetry of the changes.

We can also compare the surface temperature re-sponse with observations. An analyses of measurementsfollowing the 12 largest eruptions during the nineteenthand twentieth centuries shows statistically significant

responses in several regions; namely, a warming of morethan 28C over north-central Siberia, a 0.58–18C warmingover the central United States, and weak (,18C) coolingin the Middle East (Robock and Mao 1992). The anal-ysis of Kelley et al. (1996) finds very similar results,also including a statistically significant cooling of 228to 248C at the eastern tip of Siberia (Robock and Mao


FIG. 3. As in Fig. 2, but for the Tambora simulations. Temperature changes in excess of about0.4 K in the subtropics and about 0.9 K at midlatitudes are statistically significant for both runs.The SLP response pattern in the Tambora 2P simulation is extremely similar to the 3P result.

find a similar cooling in this region, though theirs is notstatistically significant). The model’s surface tempera-ture response compares reasonably well with these val-ues (Fig. 2, top). The model underestimates the warmingover North America. However, the impact of the AO isless pronounced in this area than over Eurasia, makingthe signal harder to discern in both observations and themodel. Overall, our ability to reproduce the large dy-namical perturbation in the winter immediately follow-

ing eruptions increases our confidence in the GCM’slonger-term dynamical response to volcanic forcing.

For the year following the eruption, the 1.4 6 0.9mb AO response in the Tambora 3P simulation or the1.3 6 0.8 mb AO response in the Tambora 2P simulationare statistically indistinguishable from the Pinatubo re-sponse, although the surface temperature responses arequite different (Fig. 3). The Tambora eruption was muchlarger than Pinatubo (and occurred earlier in the year).


FIG. 4. Surface temperature response during the years immediately following eruptions. Values are given for the summer (Jun–Aug) averagefor (top left) Pinatubo and (bottom left) Tambora 3P and for the annual average for (top right) Pinatubo and (bottom right) Tambora 3P.

The direct radiative forcing dominates the net surfacetemperature changes in the 3P simulation, leading tosignificant cooling throughout most of the NH despitethe positive AO. Dynamical and radiative perturbationsare more nearly balanced in the 2P simulation, whichproduces a response that is fairly similar to the Pinatubosimulations, though superimposed upon a larger radia-tive cooling. The hemispheric mean temperature for theyear following the eruption dropped by 0.78C in the 3Pcase, and by 0.58C in the 2P case, comparable to (es-pecially 2P) observation-based estimates for 1816(Jones and Bradley 1992), known as ‘‘the year withouta summer’’ in Europe and eastern North America. Mod-eled continental cooling following Tambora’s eruptionis generally slightly larger during summer than duringwinter, especially over Siberia and the central UnitedStates where the AO impacts are largest, due to the rapidland temperature adjustment to the larger summertimesolar forcing and the absence of dynamical warming(Fig. 4). Similar seasonality is seen in the Tambora 2Pcase. Note that the dynamically induced winter conti-nental warming tendency partially offsets the enhancedradiative summer cooling over the continents, so thatin regions where the dynamical impact is largest suchas Siberia or the central United States, the annual av-erage response is weaker than over other continentalareas. However, there is still a very strong pattern ofcooling maximizing over continental areas in the Tam-

bora 3P case, which contrasts greatly with the annualaverage response to Pinatubo that clearly shows the win-ter warming pattern (Fig. 4). Annual average resultsfrom the Tambora 2P simulations fall approximatelymidway between the other runs, as for the wintertime.

b. Physical mechanisms

The radiative impact of volcanic aerosols is both tocool the surface and to warm the lower stratosphere.After Pinatubo, the model simulates a large winterwarming in the sunlit portion of the lower stratosphere(;28C), similar to the National Centers for Environ-mental Prediction (NCEP) observations (Stenchikov etal. 2002). This increases the meridional temperature gra-dient between the midlatitudes and the Arctic in thatregion. Since the eruption takes place in June, the oceanis given relatively little time to adjust to the radiativeforcing by the following winter, and the tropical surfacecooling then is only 20.18C. This leads to a very smallcooling of the tropical upper troposphere (due to itsclose connection with the surface in convective regions),which marginally reduces the meridional temperaturegradient in the tropopause region due to the reductionof tropopause height with increasing latitude. However,this has a weaker impact than the much larger lower-stratospheric heating in the sunlit atmosphere at slightlyhigher altitudes. The gradient change increases the west-


FIG. 5. The mean AO response as a function of year after thePinatubo and Tambora eruptions. Values are pressure-averaged north-ward of 608N relative to the mean for each simulation, which is 0.6(0.2) mb lower in the Tambora 3P (2P) simulation due to overallnegative forcing of the AO. Solid horizontal lines show the 95%confidence level for the mean AO response for 10 realizations.

erly winds in this region, which refracts upward-prop-agating planetary waves toward the equator (Shindellet al. 2001b). This in turn leads to an increase in pole-ward angular momentum flow, which drives the strongersurface westerlies associated with a positive AO anom-aly. This dynamical planetary wave feedback via thestratosphere has been seen in response to volcanic erup-tions in both observations and GCM simulations (Ko-dera 1994; Perlwitz and Graf 1995; Shindell et al.2001b). The warming associated with the enhanced AOis large enough to offset most of the surface cooling inthe extratropics for the Pinatubo simulation, so that thenet change in the year after the eruption (July 1991–June 1992) is 20.18C averaged over the entire NH ex-tratropics (Fig. 4). This is similar to observations andother GCM simulations (Hansen et al. 1996; Kirchneret al. 1999; Robock and Liu 1994; Robock and Mao1995).

The modeled ozone changes in response to Pinatuboaerosol are broadly similar to those observed (Randelet al. 1995), with around 10% column depletion pole-ward of 608N during the boreal late winter to springand about 4% loss at the NH midlatitudes. These lossesresult primarily from heterogeneous chemical activationof chlorine due to the increased surface area providedby volcanic aerosols, as in other models (Al-Saadi etal. 2001; Rozanov et al. 2002). To examine the influenceof the ozone changes on the dynamics, a second Pina-tubo run was performed without ozone changes (60-yrlength). This showed a similar AO response averagedover the entire winter, but the ozone depletion in theoriginal Pinatubo simulation reduced the AO responsein the early winter months (cooling of the sunlit lowerstratosphere offset a portion of the aerosol heating) andenhanced the response in the late winter (larger ozoneloss in the polar regions caused cooling there). An en-hancement of the AO during the late winter and earlyspring due to volcanic-induced Arctic ozone depletionhas also been reported in other simulations (Stenchikovet al. 2002).

In the Tambora experiments, the atmospheric com-position was changed to preindustrial conditions for thechemistry calculations. In such an atmosphere, with noanthropogenic chlorine, aerosol effects are minimal. Themain effect of heterogeneous processing is then to re-move nitrogen oxides from the lower stratosphere. Thisallows more hydrogen oxide radicals to exist, however.Because these influence ozone in opposite directions,the net impact is quite small (the Tambora runs showedozone changes less than one-fifth of the Pinatubo chang-es). We therefore did not include ozone changes in thetransient 1959–99 ensemble, as these were designed tobe used as representative simulations of the potentialpreindustrial climate response to volcanism.

As noted, the AO response to Tambora is similar tothe response to Pinatubo for the first winter, but thesurface cooling is much larger. The increased meridionaltemperature gradient due to the aerosol heating in the

lower stratosphere is partially offset by the cooling ofthe tropical upper troposphere, which reduces the me-ridional temperature gradient in the tropopause region.Thus, despite the much larger forcing, the AO responseis approximately the same magnitude as for Pinatubo.The net change in NH extratropical surface temperaturesfor the year after the Tambora eruption (using July–June,as for Pinatubo) is 20.88C in the 3P simulation (Fig.4), demonstrating that unlike after Pinatubo, the radi-ative cooling dominates over the dynamical AO-relatedwarming. The radiative forcing remains strong for sev-eral years due to the large size of the eruption. By thethird year in the 3P simulation, the aerosol heating ofthe lower stratosphere has declined from the 38–48Cseen in the first winter to a cooling of 20.048C, whilethe upper tropical troposphere has cooled substantially(228C) as the ocean adjusts more fully to the aerosolforcing. This leads to a reduced latitudinal temperaturegradient in the tropopause region, resulting in a reduc-tion of the westerly winds and a decrease in the AO(Fig. 5). While this appears over most of the years afterthe second, it only approaches statistical significance inthe third year, at the time when the ocean response islargest. The Tambora 2P simulation also shows a mar-ginally statistically significant response in its secondyear (Fig. 5). Given that there are 33 data points, randomchance would give 1.7 significant points at the 95%confidence limit, so we believe that for individual yearsonly the positive AO seen consistently in year one ofall three simulations is robust. Averaging over years 3–11, however, all three simulations show a statisticallysignificant negative AO response.

The long-term dynamical response is thus a combi-nation of the initial positive AO forced by stratosphericheating and the delayed negative AO response forcedby tropical surface cooling. Projecting the average SLPchange between the volcanic runs and the control onto


FIG. 6. Surface temperature response (K) to decreased solar irradiance during 1680 relative to 1780 and in the volcanic transient experiments.Values are given for the solar simulations for (top left) Nov–Apr, (bottom left) May–Oct, and (top right) the annual average. (bottom right)The annual average over the five ensemble members of the transient volcanic runs (1959–99) is shown for comparison. Note the differencein scale between the two columns. Cold-season temperature responses are significant in the subtropics for values greater than about 0.2 K,at midlatitudes for values greater than about 0.3 K, and poleward of ;758N for values greater than about 0.6 K. Warm-season and annualaverage values greater than about 0.2–0.3 K are significant. The solar forcing was 20.32 W m22, while the average volcanic forcing was20.44 W m22.

the control EOFs gives a 12-yr average AO response of20.2 6 0.4 mb for the Pinatubo simulation, and 20.86 0.5 and 20.5 6 0.4 mb for the Tambora 3P and 2Psimulations, respectively. For the more frequent vol-canic eruptions during the 1959–99 period, the ensem-ble average response is 0.3 6 0.1 mb. Thus the long-term response appears to be quite small compared tounforced decadal-scale variability, except perhaps in thecase of extremely large eruptions such as Tambora 3P.

4. Comparison of volcanic and solar forcings

In contrast to volcanic eruptions, long-term decreasesin solar irradiance lead to a strong negative AO responsein our model (Shindell et al. 2001a); a 1.1 6 0.4 mbdecrease for the Maunder Minimum versus a centurylater, which causes significant wintertime continentalcooling (Fig. 6). A similar signal is present in the sum-mer (Fig. 6) due to the long-term persistence of SSTanomalies and radiative cooling of the continents, butwith reduced amplitude compared with the winter be-cause there is no dynamical enhancement. AnotherGCM has also reported an enhanced temperature re-sponse to solar forcing over NH continents, consistent

with a reduced AO (Cubasch et al. 1997). The two pri-mary forcings for climate in the preindustrial period,solar variations and volcanic eruptions, thus have quitedifferent impacts in the model. For both volcanic erup-tions and decreased solar irradiance the global averageresponse is cooling. The regional responses are dissim-ilar, however (Fig. 6). The long-term effect of volcaniceruptions on regional surface temperature changes isextremely small, owing to the cancellation of dynamicaland radiative patterns of spatial response. Comparisonof two historical periods would likely yield even lessresponse to volcanic eruptions than in the simulationsdiscussed above, which are with respect to a control runwith zero volcanic activity. For regional changes, a sig-nificant negative AO anomaly on decadal time scales isforced by volcanoes only in the event of very largeeruptions, for which tropical surface and troposphericcooling overwhelm the heating by stratospheric aero-sols. Such eruptions occur extremely rarely, however,Tambora being the largest eruption in 500 years. Notethat while large high-latitude eruptions may also affectthe AO and surface climate for a year or two (Graf andTimmreck 2001; Robock and Mao 1995), they would


be expected to show a similar lack of spatial inhomo-geneity in their long-term responses.

Solar and volcanic forcings induce such different re-sponses because the stratospheric and surface influencesin the solar case reinforce one another but in the volcaniccase are opposed. Decreased solar irradiance cools thesurface. At the same time, the lower stratosphere coolsdue to decreased absorption of incoming solar radiationby ozone owing to both the irradiance reduction and toreduced photochemical production of ozone in the lowerstratosphere. In the example of the Maunder Minimumversus a century later, the sunlit portion of the lowerstratosphere cooled by 0.28–0.48C during winter. Theupper tropical troposphere cooled by 0.68–0.88C in re-sponse to the tropical surface cooling. These persistentcoolings led to a reduced temperature gradient betweenNH mid- and high latitudes in both the lower strato-sphere and the upper troposphere, forcing a negativeAO (Shindell et al. 2001a). In contrast, volcanic erup-tions cool the surface, but aerosol heating warms thesunlit lower stratosphere. This leads to an increased me-ridional gradient in the lower stratosphere, but a reducedgradient in the tropopause region. These may separatelycause enhancements of the AO on 1–2-yr time scalesvia the stratospheric warming, and reductions on de-cadal time scales via the surface cooling. However, inlight of these opposing physical drivers of AO changes,and the short-term nature of the large dynamical re-sponse to eruptions (Fig. 5), it is not surprising that thelong-term net dynamical effect of volcanic eruptions isminimal. Additionally, the typically short time scale ofvolcanic perturbations does not allow their radiative im-pact to be fully felt by the oceans or sea ice, in contrastto multidecadal solar forcing, further weakening theirlong-term effect. Since the impact of the volcanicallyinduced dynamical changes on the annual average tem-peratures is also reduced by the seasonally opposed in-fluences of the radiative response (summer continentalcooling) and stratospherically forced dynamical re-sponse (winter continental warming), it is reasonablethat the long-term annual average impact of volcanoesis relatively homogeneous spatially.

5. External forcing of historical climate change

a. Regional changes

Since the global average surface temperature responsebasically follows the radiative forcing for both volcanicand solar forcing (the model is slightly more sensitiveto solar than to volcanic forcing due to positive feed-backs from stratospheric ozone and the AO response),relatively simple models such as energy balance modelscan give reasonable estimates of global mean temper-ature variations in past centuries. Given that the regionalresponse is not always a simple function of the radiativeforcing, more complex models such as GCMs are re-quired to investigate regional climate change. All cli-

mate models are limited by the accuracy to which thehistorical forcing is known. While paleoindicators givean idea of the size of past eruptions, the spatial distri-bution of the aerosols and the exact timing of someeruptions are not well known. However, examining ourensemble of simulations driven by observed 1959–99eruptions, we find a good correlation between the globalsurface temperature response and the total forcing fordecadal time scales. Though our sample is small (4 de-cades), this suggests that the global response may notbe greatly sensitive to the precise details of the aerosoldistribution. More importantly, spatial inhomogeneitiesare relatively small. Comparison of the ensemble meanpattern of surface temperature anomalies relative to thezonal mean for each decade against the same patternfor the remaining decades shows no more areas of sta-tistically significant correlations (95% level) than ran-dom noise outside of a warming over Europe and acooling in the Bering Strait region. This argues againstthe presence of major regional anomalies associatedwith volcanically forced changes in the zonal flow inother areas.

b. The Maunder Minimum

The simulated external forcing of regional surfacetemperature anomalies can also be compared with pat-terns derived from a compilation of diverse proxy cli-mate indicators. We use model results from simulationsforced with reduced solar irradiance during the MaunderMinimum compared with a century later, and from ourvolcanic simulations (the twentieth-century transients,Pinatubo, and Tambora 3P). The latter serve as a gen-eralized response to volcanic forcing, which was ;0.16W m22 more negative during the late seventeenth cen-tury (1660–90) than the late eighteenth century (1770–90; Crowley 2000). Proxy data have been averaged overthese same decades to obtain decadal time scale patterns.

We compare the response in regions that exhibit large,coherent variations in the proxy-data record (Fig. 7).The unforced variability is the one-sided 95% confi-dence range for the occurrence of a regional anomalyof the same sign as the proxy data in any single 25-yrperiod in the 250-yr control run. The decadal-scale re-gional response in all the volcanic experiments is ex-tremely weak. The only significant responses are overEurope for the transient, over Europe and the NorthAtlantic for Pinatubo, and over Europe and North Amer-ica for Tambora 3P. Aside from the European warming,it is clear that there is no consistent pattern of regionalsurface response to volcanic forcing in these areas. Ad-ditionally, there is no evidence of a response that scaleswith the forcing, which ranges from 20.25 during1969–78 to 21.39 W m22 for Tambora 3P. Using themean response over the four decades of the transientensemble as broadly representative of the influence ofvolcanic activity, the volcanic response is well withinthe range of the unforced variability in all regions except


FIG. 7. Long-term regional annual average surface temperature changes. The upper bars show the mean regionalresponse to volcanic forcing in the transient ensemble for the 1960s (1959–68), 1970s (1969–78), 1980s (1979–88),and from the Pinatubo and Tambora 3P simulations. Decades were chosen so as to exclude large volcanic eruptions inthe last 2 yr, which would impact the radiative forcing but not yet be fully felt by the climate system. Values for alltemperature changes are given for the outlined regions as shown on the lower map compared to the mean change in theNH extratropics. The lower section shows 1660–90 vs 1770–90 regional temperatures in the multiproxy reconstruction.Model results for solar forcing changes are shown for the same time period. The volcanic response is averaged overthe 1960s to 1990s (forcing 5 20.44 W m22) as indicative of the long-term response to sizeable volcanic activity. The95% confidence levels for the decadal volcanic transients are 0.158C for North America, 0.188C for the North Atlantic,0.078C for Europe, and 0.128C for Siberia. They are ;0.058C for the longer Pinatubo, Tambora, and solar simulations.Note that the statistical significance of the GCM results comes from a signal-to-noise comparison over a large numberof experiments. Thus the response can be significant while still being less than the unforced variability range of anyone 2–3-decade period.

Europe. This occurs despite a relatively large averageforcing of 20.44 W m22. In contrast, the solar forcingcauses a large, statistically significant response in eachregion that has the same sign as the proxy data andexceeds unforced variability. Though these simulationsmay underestimate unforced variability on long timescales due to the lack of a dynamic ocean, our resultsare consistent with results from a comparison betweenclimate proxies and an unforced run of the Hadley Cen-ter coupled GCM (Collins et al. 2002), which also sug-gests that the majority of preindustrial climate changeover NH mid- to high-latitude continents resulted fromexternal forcing.

While paleoclimate reconstructions of surface tem-peratures in past centuries are uncertain, there is broadagreement among different methods and data sources,

at least at the level of the NH mean temperature vari-ations (Briffa et al. 1998, Crowley 2000; Jones et al.1998; Mann et al. 1999). On the regional scale, Euro-pean temperature estimates are most reliable, as histor-ical and a few long instrumental data series augmentthe more widespread proxy data such as tree rings inthe proxy network reconstruction. The reconstructed Eu-ropean regional temperature anomaly is thus a key testof a model’s regional response to forcing. Comparingthe Maunder Minimum period a century later, the morenegative volcanic forcing fails to produce even the rightsign of temperature change in this region, as all thevolcanic simulations generate a warming relative to theNH extratropical mean while the proxy data and theresponse to solar forcing show cooling outside the rangeof unforced variability (Fig. 7).


FIG. 8. (top) Reconstructed and (bottom) simulated annual average temperature difference between 1660–80 and 1770–90. The reconstructedsurface temperatures are based on a multiproxy estimate using tree rings, ice cores, corals, and historical data (Mann et al. 1998). Modelresults are based on the sum of the response in two simulations: one incorporating reconstructed solar irradiance changes during this periodand one using volcanic forcing scaled to changes over this time.

The reconstructed volcanic forcing during the lateMaunder Minimum as compared with the late eighteenthcentury is only about one-quarter the mean 1959–99volcanic forcing of 20.66 W m22 in the same recon-struction (Crowley 2000). Since the Sato et al. data donot extend back before 1850, we use the Crowley re-construction in our discussion of the centennial-scaleinfluence of volcanic forcing on climate. We believe,however, that it may represent a high-end estimate ofthe volcanic forcing during the past millennium (Satoet al. 1993; Crowley 2000).

Assuming that the responses can be linearly summed[given that the long-term volcanic influence on the non-linear AO–North Atlantic Oscillation (NAO) and henceon regional changes is very weak], the surface temper-

ature response to a forcing of this magnitude, estimatedsimply as one-quarter the response during the 1959–99experiments, can be added to that due to reconstructedsolar irradiance changes during this same period (Leanet al. 1995). This provides an estimate of the totalchange (Fig. 8). The result shows a regional pattern ofcontinental cooling due to a forced negative AO anom-aly. While the global mean change is 75% due to solarforcing and 25% to volcanic forcing, all of the AO re-duction is attributable to the solar forcing, as the vol-canic forcing actually gave a weak AO increase.

We compare the total response with that derived fromsurface temperature pattern reconstructions based on di-verse proxy data (Mann et al. 1998). A favorable agree-ment is found between the simulated and reconstructed


temperature anomalies over this period (Fig. 8), withboth showing large, spatially coherent anomalies in mid-latitudes. Agreement is especially good over NorthAmerica, where the impact of reduced northwesterliesmoving around the Rockies is clearly visible in bothcooling patterns. The model also reproduces the coolingthat arcs up through western Europe into Asia, thoughthe Middle East does not show the warming seen in thereconstruction (it is relatively warm compared to Eur-asia, however). While both the GCM and the recon-struction show warming over parts of the North Atlantic,the reconstructed anomaly is less than the unforced var-iability, which is quite large in this region. Additionally,we regard the model simulation results as being lessreliable over the oceans, as they do not include changesin ocean dynamics, which may be an important feed-back, at least in the Atlantic where surface anomaliesare buffered by communication with the deep ocean(Delworth and Dixon 2000). The model results also donot account for possible changes in the El Nino–South-ern Oscillation and potential extratropical responses,particularly with regard to the Pacific region. Never-theless, the basic pattern of anomalously cold continentsis clear in both the paleoclimate data and the modelresponse, and it is in these areas that both values aremost reliable. The amplitude of the cooling is larger inthe model, however. This version of the GISS GCM hasa climate sensitivity at the high end of the modeled range[;18C (W m22)21]. Additionally, the proxy reconstruc-tion is implicitly spatially smoothed through the use ofa limited eigenvector basis set to approximate large-scale temperature changes. Moreover, data availabilitylimits the size of this basis set in earlier centuries, lead-ing to a greater smoothing of regional temperature var-iations prior to the eighteenth century in particular. Notethat there is some independent evidence for a reducedNAO during the late seventeenth and eighteenth cen-turies from European historical reconstructions (Luter-bacher et al. 1999). Additional support for Atlantic sec-tor circulation change consistent with a reduced AO–NAO during the LIA comes from sediment cores (Kieg-win and Pickart 1999) and from historical evidence forgreater northeasterly flow of continental air into Europe(Wanner et al. 1995). Other proxies that have been re-lated to North Atlantic climate variability also appearto be correlated with solar variations on millennial timescales (Bond et al. 2001). Based on our simulations, wecaution that very large eruptions may add difficulty toreconstructing past AO or NAO indices based on tem-perature proxies because the rapid continental responseto very large radiative forcing yields a pattern of rela-tively cold continents during winter despite the positiveAO–NAO in the Tambora 3P run (Fig. 3).

c. Centennial-scale climate change

The relative importance of volcanic and solar forcingto long-term climate change can be examined more

broadly by comparing the reconstructed forcings (Crow-ley 2000) across other time periods. The seventeenth-century average is representative of the LIA. Comparedwith the period 1850–99, before the bulk of the an-thropogenic forcing, the seventeenth-century solar forc-ing anomaly is estimated to be 20.19 while the volcanicis 20.30 W m22. Compared with the period 1000–1299,which coincides with warmer hemispheric mean tem-peratures (Briffa et al. 1998; Jones et al. 1998; Mannet al. 1999) and the MWP, both the seventeenth-centurysolar and volcanic forcings were 20.16 W m22. In bothcases, the volcanic forcing was at least as large as thesolar. Since the global temperature change follows theradiative forcing fairly closely, this suggests that thevolcanic impact on the global mean LIA cooling wasimportant, as previously hypothesized (Crowley 2000;Free and Robock 1999). The global mean response tothe combined solar and volcanic forcing exceeds theunforced variability, which is ;0.198C (95% confidenceinterval), by a factor of 2–2.5 for the LIA and MWP.Although the climate sensitivity of this version of theGCM is relatively high, the forced component wouldstill exceed the unforced with a more typical sensitivityof ;3/48C (W m22)21, suggesting that much of the var-iability seen in climate reconstructions does not arisemerely from internal chaotic variations. We reiterate,however, that the absence of a fully coupled ocean inthese experiments limits the internal variability.

The presence of a sizeable decrease in solar irradianceduring the LIA relative to the MWP or the latter halfof the nineteenth century implies a significant regionalcooling over midlatitude NH continental areas. On theregional scale, the solar-induced anomalies scaled to theLIA forcing (roughly one-half the 20.32 W m22 dif-ference between the late Maunder Minimum and a cen-tury later) would be of the appropriate sign and in excessof the unforced variability in both Europe and NorthAmerica (Fig. 7). The appropriately scaled volcanic-induced anomalies (about one-third to two-thirds theforcing in the 1959–99 transients) would not meet thesecriteria in any region. The modeling results show con-tinental changes 1.5–3 times the NH mean change inresponse to a total forcing that is roughly two parts solarto one part volcanic (Fig. 8). While the solar-to-volcanicratio is between two-thirds and one for the forcingchange between the LIA and the MWP or between theLIA and the latter half of the nineteenth century, thisstill leads to continental cooling roughly 1.3–2.5 timeslarger than the NH mean.

Such differences between the NH mean and the con-tinental interiors are in accord with the contrast betweenreconstructions of full NH (land and ocean, Tropics andextratropics) annual mean temperature changes in pastcenturies based on diverse proxy (‘‘multiproxy’’) data(Crowley and Lowery 2000; Jones et al. 1998; Mannet al. 1998; 1999) and extratropical NH temperaturereconstructions based on tree-ring data alone (Briffa etal. 1998; Esper et al. 2002). These latter reconstructions


are representative of the more restricted continental re-gions of the NH, and target warm-season temperaturesthrough the tree-ring species and parameters employedin the reconstructions (Mann and Hughes 2002). Dueto the seasonally and spatially varying surface temper-ature responses to solar and volcanic forcing discussedearlier, continental proxies sensitive to warm-seasontemperatures emphasize the enhanced continental sum-mer cooling associated with the radiative response tovolcanic forcing, but miss the often-present dynamicalcontinental winter warming in response to those sameevents. This seasonal sensitivity thus effectively rectifiesthe temperature response to volcanic forcing, exagger-ating the recorded hemispheric (land and ocean) coolingassociated with volcanic forcing (Mann 2002). For thesereasons, tree-ring-based continental temperature recon-structions show evidence of far greater cooling in re-sponse to individual volcanic events (Briffa et al. 1998)and greater cooling on multidecadal-to-century timescales (Esper et al. 2002) than full hemispheric recon-structions. Our results suggest that these differences donot require any additional extratropical forcing or in-ternal multicentury oscillations (Bradley et al. 2001;Broecker 2001; Esper et al. 2002), as they can be ac-counted for by the dynamical response to both volcanicand solar variations.

6. Conclusions

Volcanic forcing is important for both regional cli-mate change on a year-to-year basis, and for long-termclimate change on a global scale. It appears unlikely toinduce sustained regional climate changes that exceedunforced variability, however. In contrast, solar forcingcreates long-term regional climate changes that aregreater than unforced variability, and which resemblethose seen in proxy-based reconstructions of historicaltemperature change during the Maunder Minimum. Theregional anomalies result primarily from a forced shiftin the AO.

During the last several decades, there has been anapparent upward trend in the AO (Thompson and Wal-lace 1998). Solar forcing, however, has increased littleif at all over this period, by less than 0.05 W m22 sincethe 1960s (the exact value depends upon the particularyears chosen for the trend calculations). Such a smallincrease in irradiance would cause an AO enhancementof about 0.1 mb, much smaller than the apparent trendof 3 mb, which is more likely to be attributable to in-creases in greenhouse gases, and to a lesser extent, Arc-tic ozone depletion (Shindell et al. 2001b).

Downward shifts in the AO are closely correlatedwith cold temperature extremes in Europe (Higgins etal. 2002; Thompson and Wallace 2001), consistent withthe historical impression that the LIA in Europe waspunctuated by short periods of extreme cold exceedingthat seen in the multidecadal averages of proxy datashown here (Pfister 1995). The global mean ‘‘Little Ice

Age’’ signal can therefore be attributed to both volcanicand solar forcing, but while volcanic forcing may haveplayed a major role in global-scale cooling, the muchlarger regional changes and probable changes in thefrequency of extreme cold events were likely drivenprimarily by solar variability.

Acknowledgments. Modeling of chemistry–climateinteractions and stratospheric dynamics at GISS is sup-ported by NASA’s Atmospheric Chemistry Modelingand Analysis Program, while climate modeling is fundedby the NASA Climate Program Office. GAS and DTSwere also supported by NSF Grant ATM-00-02267. Wethank M. Sato for providing the aerosol data, Scott Ruth-erford for assistance with analysis of the surface tem-perature reconstruction, David Rind for discussion, andthe three anonymous reviewers for their comments.

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