Atmosphere and ocean dynamics: contributors to the European Little Ice Age?
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The role of a reduction in the Atlantic meridional overturning and that of a persistently negative North Atlantic Oscillation in explaining the coldness of the European Little Ice Age (LIA) has been assessed in two sets of numerical experiments. These experiments are performed using an intermediate complexity climate model and a full complexity GCM. The reduction in the Meridional Overturning Circulation (MOC) of ca. 25% is triggered by a conventional fresh-water hosing set-up. A persistently negative NAO winter circulation, at NAO-index value −0.5, is imposed using recently developed data-assimilation techniques applicable on paleoclimatic timescales. The hosing experiments lead to a reduction in oceanic meridional heat transport and cooler sea-surface temperatures. Next to a direct cooling effect on European climate, the change in ocean surface temperatures feedback on the atmospheric circulation modifying European climate significantly. The data-assimilation experiments showed a reduction of winter temperatures over parts of Europe, but there is little persistence into the summer season. The output of all model experiments are compared to reconstructions of winter and summer temperature based on the available temperature data for the LIA period. This demonstrates that the hypothesis of a persistently negative NAO as an explanation for the European LIA does not hold. The hosing experiments do not clearly support the hypothesis that a reduction in the MOC is the primary driver of LIA climate change. However, a reduction in the Atlantic overturning might have been a cause of the European LIA climate, depending on whether there is a strong enough feedback on the atmospheric circulation.
KeywordsLittle Ice Age Meridional Overturning Circulation North Atlantic Oscillation Data assimilation
Recent reconstructions of European temperatures provide some context and allow an assessment of the amplitude of the natural climate changes that affected the continent over the last 500 years (Luterbacher et al. 2004). In particular, European climate between the 16th and 19th centuries has seen cold multi-decadal periods, within the period known as the “Little Ice Age” (LIA) (Bradley and Jones 1993). While changes in solar variability and volcanism are possibly main causes for the global cooling of the LIA (Jansen et al. 2007), feedbacks via atmospheric and ocean dynamics determine the regional climate response to a large extent. Over the North Atlantic sector, changes in the atmospheric circulation patterns (Luterbacher et al. 2001, 2002), sea surface temperatures (Keigwin 1996; Keigwin and Pickart 1999) and sea ice cover (Ogilvie and Jonsson 2001) were documented during the LIA, and links between the modes of variability of the atmospheric circulation and the LIA period were proposed (Luterbacher et al. 2002; Shindell et al. 2001). It has also been speculated whether changes in the ocean thermohaline circulation were responsible for the LIA coldness (Bianchi and McCave 1999; Broecker 2000). Still, the relative importance of these mechanisms in shaping the regional and global characteristics of the LIA climate needs to be clarified.
The North Atlantic Oscillation (NAO) is the main mode of winter atmospheric circulation variability in the North Atlantic (Hurrel 1995), and is coupled to the North Atlantic sea surface temperature (SST) through latent heat fluxes (Rodwell et al. 1999). The state of the NAO, measured as the difference in sea level pressure (SLP) between the Azores and Iceland, relates to the strength of the westerly circulation that carries warm and wet air from the Atlantic into western Europe. Reconstructions of the SLP fields back to 1,500 indicate that the negative phase of the NAO has been more active during extreme cold periods of the LIA, leading to reduced influence of the moist and warm zonal flow from the northeastern North Atlantic, which favored cool and dry European winters (Luterbacher et al. 2001, 2002). In addition, the recurrence of positive SLP anomalies centered over Scandinavia or northern Europe, which caused anomalous advection of cold air towards central and eastern Europe, enhanced the wintertime cooling (Luterbacher et al. 2001). Several reconstructions of the NAO index from independent proxy data (Rodrigo et al. 2001; Cook et al. 2002) support a link between the negative phase of the NAO and long lasting cold periods within the LIA. This hypothesis was substantiated by the modeling study of Shindell et al. (2001), who showed a relationship between external forcings (i.e. reduced irradiance) and the NAO. Basically, changes in the stratospheric temperature and wind anomalies modify the mid-latitude planetary waves refraction, which interact with the intensity of the westerly winds, and subsequent the NAO. The experiment confirmed that regional climate changes associated with the changing state of the NAO can be much stronger than hemisphere wide changes.
Recently, the opposite with a persistently positive phase of the winter NAO leading to warm climatic spells has been identified to have occurred in the Medieval Climate Anomaly (Trouet et al. 2009).
Variations in the meridional heat transported by the Meridional Overturning Circulation (MOC) in the North Atlantic affect the heat loss to the atmosphere at mid and high latitudes, and consequently influence the climate downstream. Although one study hypothesized that a slowdown of the MOC may have acted as an amplifying mechanism in the LIA (Bond et al. 1997), to date, the evidence for Holocene MOC variability from geochemical proxy data is inconclusive (Keigwin and Boyle 2000). Bianchi and McCave (1999) analyzed from sediment grain size the speed of the Iceland-Scotland overflow water (ISOW), one of the components of North Atlantic deep water, and found that since 8,000 years ago cold climate periods, including the LIA, coincided with a less intense overflow, and vice versa. On interdecadal time scales, the rate of deep water formation in the North Atlantic has been linked to the NAO (Dickson et al. 1996). The strong NAO minimum of the early 1960s lead to a reduction and ultimately a cessation of deep water formation in the Labrador Seas, while formation of deep water in the Greenland Sea reach a maximum in that period. Recently, observations by Boessenkool et al. (2007) showed an inverse correlation between the ISOW and the NAO, and suggest that interdecadal changes in Labrador Sea water play a key role in transmitting the NAO signal to the deep ocean overflows.
The present study aims to distinguish between a persistently negative NAO-type circulation and a reduction in the Atlantic MOC as the primary drivers of the cooling observed in the European LIA. Two types of experiments are conducted; one in which a climate is simulated with a constant and negative NAO-index in winter, and one with a reduced overturning. In order to distinguish between the two hypothesis for LIA climate change, we need to compare the model results with a reconstruction of LIA climate. We only compare simulated and reconstructed temperatures; other elements, like precipitation, are not included in this comparison. The temperature reconstruction used is from Luterbacher et al. (2004) and is based on a mixture of proxy-records and early instrumental temperatures. The experiments conducted in this study should be regarded as sensitivity experiments in an idealized framework. Nevertheless, the comparison between the experiments and reconstructed climate will shed some light on the question which process is more likely to have occurred and is more likely to contribute to the coldness of the European LIA. Moreover, the comparison with a reconstruction of climate may also result in the conclusion that neither hypothesis seems likely to be valid, indicating that the drivers for LIA climate change are more complex than suggested by the two simple hypotheses investigated in this study.
In the present study we use a General Circulation Model (GCM) of intermediate complexity and a state-of-the-art GCM to investigate the separate effects of a reduced MOC and of a persistently negative NAO phase on setting the anomalous coldness of the LIA. The difference in model complexity should give an estimate of the robustness of the results.
For the first experiment, a freshwater flux is added to the North Atlantic Ocean to force a reduction of the MOC. For the second experiment, we impose a negative NAO-type circulation representative of the LIA period (Luterbacher et al. 2002) on the model atmosphere. For this purpose, two data assimilation techniques are used that reproduce, in a time averaged sense, the prescribed perturbation in the large-scale atmospheric circulation. Both techniques leave the atmosphere free to respond in a dynamically consistent way to the change in climatic conditions; in particular, synoptic scale variability is not suppressed.
All experiments conducted in this study are idealized to the point that radiation changes, due to changes in dust loading related to explosive volcanic outbursts or changes in solar activity, are ignored. Changes in the radiation balance must have been important for explaining LIA climate (e.g. Grove 1988), and in that sense the experimental set-up is an oversimplification. However, the direct effects of changes in the radiation budget will lead to large-scale changes in climate. In this study we are interested in the patterns of LIA climate change over Europe. An additional motivation for this idealized setting is that combining more than one possible driver of LIA climate change in the experiments will complicate the interpretation of the results.
Recently, Sedlácek and Mysak (2009) attempted to model the response of the North Atlantic ocean in the LIA period. They prescribed constant windstress field associated with negative NAO type circulations in their intermediate complexity GCM. Their experiment has some similarities to the data-assimilation experiments performed in this study in that the impact of the anomalous character of the LIA atmospheric circulation on the ocean circulation is estimated. However, an advantage of the approach adopted in the current study is that the ocean-atmosphere-sea ice system remains coupled in a dynamically consistent way, while synoptic variability is maintained.
The paper is organized as follows. Section 2 briefly describes the models and experimental design. Section 3 includes a short summary of the data assimilation techniques used to force the NAO towards a long term negative value. Section 4 presents the results of the freshwater hosing and negative NAO-assimilated pattern experiment. Finally, Sect. 5 contains a discussion and conclusions.
2 Model description and experiment design
2.1 ECBilt-Clio Model description
The intermediate complexity model, ECBilt-Clio, is a coupled ocean-atmosphere-sea ice general circulation model (Opsteegh et al. 1998; Goosse and Fichefet 1999). The atmospheric component (ECBilt) resolves 21 wavelengths around the globe, and it has 3 levels in the vertical, at 800, 500 and 200 hPa. The dynamical part is an extended quasi-geostrophic model where the neglected ageostrophic terms are included in the vorticity and thermodynamic equations as a time dependent and spatially varying forcing. With this forcing the model simulates the Hadley circulation qualitatively correctly, and the strength and position of the jet stream and transient eddy activity become fairly realistic in comparison to other T21 models. The essentials of baroclinic instability are included, but the variability associated with it is underestimated compared to modern observations. The model contains simple physical parameterizations, including a full hydrological cycle.
The oceanic component (Clio) is a primitive equation, free-surface ocean general circulation model coupled to a thermodynamic-dynamic sea ice model and includes a relatively sophisticated parametrization of vertical mixing (Goosse et al. 1999). A three-layer sea-ice model, which takes into account sensible and latent heat storage in the snow-ice system, simulates the changes of snow and ice thickness in response to surface and bottom heat fluxes. The horizontal resolution of Clio is 3° × 3° and it has 20 unevenly spaced layers in the vertical.
2.2 HadCM3 Model description
HadCM3 is the third version of the Hadley Centre coupled Atmosphere-Ocean GCM and has been described in detail by Pope et al. (2000) and Gordon et al. (2000), where a validation of some climate parameters is also given. Here, only a brief description of the model is given. The atmospheric part of HadCM3 employs a longitude-latitude grid with a resolution of 3.75° by 2.5°. Vertically, the model resolves 19 layers in a hybrid coordinate system. The ocean component has a higher resolution of 1.25° by 1.25° on 19 unevenly spaced levels with increasing resolution near the ocean surface. HadCM3 produces a stable climate without the use of flux adjustments, with the exception of the deep ocean, where a small drift remains (Pardaens et al. 2003). The sea ice model uses a simple thermodynamic scheme and contains parameterizations of ice drifts.
In the first of the two sets of experiments performed with the ECBilt-Clio and HadCM3 models, a negative state of the NAO is imposed on the model’s atmosphere by forcing the NAO index towards a 30 year-mean value of approximately −0.5. For the ECBilt-Clio model, we use a data assimilation technique developed at the European Centre for Medium Range Weather Forecasts (ECMWF), which applies an adjoint model. This method has been successfully applied in recent paleoclimatology studies with the ECBilt-Clio model (van der Schrier and Barkmeijer 2005, 2007; van der Schrier et al. 2007). As it is not feasible to construct an adjoint model for a GCM as HadCM3, here we use an assimilation algorithm based on the Data Assimilation through Upscaling and Nudging (DATUN) or “pattern nudging” method (von Storch et al. 2000; Jones and Widmann 2003). A description of both data assimilation approaches and their differences is given in the Appendix.
Both data assimilation techniques assimilate large-scale changes in the atmospheric circulation by calculating an artificial forcing to the model tendencies. The method of calculating this additional forcing, as well as the actual forcing fields, differ between the two models. Both assimilation methods are designed specifically to modify the large-scale circulation without suppressing the synoptic scale variability, which can adjust in a dynamically consistent way.
In this experiment, we only assimilate a negative NAO-like circulation in the model and do not change irradiance, parameterizing solar activity or changes in atmospheric dust loading due to explosive volcanic eruptions. No action is taken to change the strength of the MOC either.
The second experiment consists of a freshwater hosing in order to reduce the mean strength of the MOC by about 25%. To achieve this, an artificial freshwater flux of 0.1 Sv in the HadCM3 model and of 0.075 Sv in the ECBilt-Clio model is added to the North Atlantic basin between 50°N and 70°N in a similar way to that described by Stouffer et al. (2006). The induced reduction in the overturning depends on both the magnitude and duration of the perturbation. The aim of the latter experiment is not to simulate a complete shut down of the MOC, as even for the extreme climate of the Last Glacial Maximum proxy data indicated a MOC reduction of no more than 30% compared to present climate (Weber et al. 2007). A decrease of that latter magnitude would be consistent with available Holocene paleoclimatic data (Bianchi and McCave 1999). A reduction of the intensity of the MOC by 25% for a simulation of LIA climate is then at the upper range of what could be considered as compatible with data. In the HadCM3 model, the freshwater pulse is implemented from year 100 onwards of the control run and is held constant during 100 years of integration, while for the ECBilt-Clio model we consider a simulation of 300 years, in which the perturbation is applied from year 50 onwards and is held constant over 50 years. For the ECBilt-Clio model, we analyze a 100-year period right after the cessation of the freswater inflow. The strength of the overturning remains (statistically) constant over this period. The inflow of freshwater is constant in time and a possible simulation of Heinrich events is not attempted. In this experiment, we only modify the strength of the MOC and do not assimilate a negative NAO-like circulation or change the irradiance in an effort to accommodate changes in solar activity or changes in atmospheric dust loading.
An effective way of modifying the mean strength of the MOC is by addition of freshwater to the model ocean. With this experiment, we do not claim that if a MOC change may have been instrumental for Europe’s cold climate during the LIA, that an increase in freshwater inflow must have occurred. The scope of this study does not include an analysis of the causes for a possible reduction in MOC during Europe’s LIA, nor for the causes of a persistently negative NAO-like circulation for that matter.
In addition to the above simulations, a 200 year control run at preindustrial conditions with HadCM3, which is itself a continuation of an earlier control run performed at the Hadley Centre, and a 100 years control run with the ECBilt-Clio model are performed.
2.4 Nudging method
Nudging was originally developed as a method for the assimilation of measurement data into weather forecasting models. The general idea in nudging is that the distance between the model state and the measurements of meteorological variables is minimized at every grid point of the weather forecasting model by adding correction terms to the model fields.
In contrast to conventional nudging approaches, the pattern nudging in the DATUN method is based on a decomposition of the model field into a climatology and a set of orthogonal modes as in Eq. (1) of the Appendix. The model dynamics are forced towards the large scale target pattern, without directly affecting components of the climate state that are not constrained by proxy data, and without suppressing synoptic-scale variability. A more expansive discussion on the approach can be found in Jones and Widmann (2003) and Widmann et al. (2009). The implementation of this technique in HadCM3 can be found in Kleinen et al. (2007). See Widmann et al. (2009) for a discussion of these two approaches for data assimilation on climatic timescales.
To adopt this approach for use in HadCM3, care has to be taken not to violate mass conservation in the model. A target pattern in terms of sea-level pressure could induce rather large mass fluxes. Therefore, it is undesirable to influence the sea level pressure field directly. While HadCM3 is a non-spectral model, where the winds in u and v direction are prognostic variables, these variables are transformed to vorticity and divergence. Changes in divergence would again introduce artificial mass fluxes into the model, but nudging vorticity avoids these artificial mass fluxes. A nudging pattern can therefore be determined by regressing the vorticity field onto the NAO index.
3.1 Slowdown of the meridional overturning circulation
In this section we analyze the climate response in the circum-North Atlantic region to a slowdown of the MOC by ca. 25%. In both the ECBilt-Clio and HadCM3 models, the intensity of the MOC is defined as the maximum of the meridional overturning streamfunction in the North Atlantic below a depth of about 500 m. For the ECBilt-Clio model, in addition to the 300 years simulation discussed here, several other simulations with a slightly different magnitude of the freshwater forcing and duration of the “hosing” were performed. All simulations show robust results in terms of MOC reduction, deep-water convection changes, and surface temperature response.
In the analysis of the results with the HadCM3 model, the means over the final 30 years of the hosing experiment are compared to the mean over the final 30 years of the control run, while for the ECBilt-Clio model we compare means over 100 years of the control and hosing simulation. Statistical significance in the difference of the means is tested using a standard Student t-test at a significance level of 90%. For simplicity, discussion focuses on the winter and summer seasons only.
The weakening of the MOC in the HadCM3 model induces a cooling in surface air and sea temperatures over most of the North Atlantic ocean (Fig. 2b). The region with maximum cooling (up to 3°C) lies between Iceland and Greenland, while temperatures of about 1–2°C colder are found over the Labrador Sea and in the center of the subtropical gyre around 40°N. Significant colder temperatures of 0.5–1°C extend over southern Greenland and in the north Africa-southern Asia region; warmer temperatures of about 0.4°C are seen over central Canada. As discussed in Stouffer et al. (2006), the HadCM3 model shows a warming of the high latitude North Atlantic caused by a northward shift of the deep-water convection and the associated increase in northward heat transport. On the other hand, much of the deep-water convection that takes place in the latitude belt between 50°N and 70°N is weakened by the freshwater hosing. In the summer season, the cooling extends further south in the North Atlantic and over the surrounding continental areas (Fig. 2d). In particular, temperatures are about 0.5–1°C colder over southwestern Europe and 0.3°C colder over central Russia.
The response in the atmospheric circulation is shown as differences in SLP between the hosing and control run. In winter, significant increases in the SLP appear in the southeastern North Atlantic, the Labrador Sea, and over a region extending from Scandinavia towards Eastern Europe, whereas a significant anomalous cyclonic circulation is seen over North America, with a maximum in the northwest (Fig. 3b). Negative SLP anomalies present over the Greenland Sea and Arctic Ocean respond to the warmer SST there. In summer, an anomalous anticyclonic circulation centered around 40°N covers the subtropical North Atlantic, and positive SLP anomalies are also found over Scandinavia and Russia (Fig. 3d). The increase in the anticyclonic circulation over the subtropical North Atlantic seems to induce a strengthening of the westerly winds south of 50°N and a weakening north of it, both in winter and summer, although these changes are rather small (0.1–0.2 m/s) and not significant.
3.2 Prescribed changes in the NAO index
The NAO index gives an indication of the predominant atmospheric circulation in the North Atlantic area. If the NAO index is positive, i. e. lower than average pressure over Iceland, and higher than average pressure over the Azores, weather in Europe is dominated by circulation from the Atlantic. This results in mild and moist winter weather in western Europe. Conversely, a negative NAO index leads to a much reduced influence of North Atlantic weather, thereby leading to comparatively cool and dry winters.
The anomalous winter temperatures (Fig. 11a) show the typical pattern associated with negative NAO-type circulations, with anomalous low temperatures over northern Europe and the southeastern USA, and anomalous high temperatures in the Mediterranean region and eastern North America.
In the data assimilation experiments, only winter circulation is modified. This could also affect summer conditions through an impact of the winter atmospheric circulation on components of the climate system with a large memory, like the upper ocean or the land surface. However, Fig. 12a shows that the impact on the summer temperatures in this experiment is modest.
Figure 11b shows the difference in surface temperature between the model experiment, where the winter (DJF) NAO index has been forced toward a 30 year mean value of −0.5, and the control run, where the NAO index is in a neutral mean state. Only the DJF NAO index was influenced, and in the HadCM3 simulation, temperature differences in MAM, JJA, and SON are very small, and not statistically significant. In DJF, on the other hand, some greater changes are apparent. West of the Atlantic, the negative NAO leads to cooling in the south-eastern part of the US, with lower SST in the North Atlantic up to −2°C, and a slight warming over Greenland. Over northeastern Europe and Russia temperatures are colder up to 1.5°C, with hardly any changes apparent over Scandinavia. In Central and Eastern Asia, a strong warming of up to 2°C is also seen (not shown).
Figure 12b shows that the impact on the summer temperatures on western Europe in this experiment cannot be distinguished from noise.
4 Comparison to a reconstruction of Little Ice Age climate
The hosing experiment in both models leads to negative SST anomalies over the North Atlantic in both winter and summer which modifies the exchange of heat from the ocean to the atmosphere. This will affect sea level pressure as well. In this sense, the reduction of the overturning circulation will feedback on the character of the atmospheric response. The location of the winter atmospheric circulation anomaly in the ECBilt-Clio model (Fig. 2a) suggests that advection of cold air from high latitudes over Scandinavia is a main factor contributing to the decrease in surface air temperatures over northwestern Europe. In the HadCM3 simulation, the anomalous high pressure center over Russia-eastern Europe is likely to produce the lower temperatures seen to the east of the Caspian Sea. It is interesting to note that the winter atmospheric circulation anomaly that develops in the hosing experiment with the ECBilt-Clio model exhibit some similarity with the large-scale SLP anomalies documented for several periods within the LIA (Luterbacher et al. 2002; van der Schrier and Barkmeijer 2005).
In both hosing experiments, changes in the North Atlantic wind stress curl associated with the winter circulation anomalies are not significant, implying that feedback mechanisms between the MOC reduction and the wind driven circulation via changes in the atmospheric circulation are not at work in the model experiments.
Assimilating a persistently negative winter season NAO circulation in the ECBilt-Clio model leads to a strong cooling over northwestern and northern Europe, while higher temperatures are observed over the Mediterranean region. This response is similar to what would be expected based on modern observations. The reconstruction of the winter temperature during the 1675–1715 and 1790–1820 periods shows severe cooling over the whole of Europe, whereas the simulation displays a north-south contrast.
The nudging experiment to impose a negative NAO index in the HadCM3 model in winter lead to decreased temperatures over eastern-Europe and northern Russia (Fig. 11b), a result that compares well with the temperature reconstructions for the 1790–1820 period (Fig. 13b). However, the simulation fails to show cooling over western Europe, a rather surprising result considering the link between the NAO and the LIA invoked in the literature (Shindell et al. 2001; Luterbacher et al. 2002). A cause for this could be related to the reorientation of the NAO pattern in this simulation, which results in advection of air from southeastern to western Europe, rather than advection of the colder continental air masses. On the other hand, the model shows negative SST anomalies over the western North Atlantic between 40°N and 50°N, which are consistent with a negative NAO index (Hurrel 1995).
The persistence of the change in temperatures into the summer season is, however, weak. In summer, when no data assimilation is performed, the simulation with the ECBilt-Clio model shows no significant cooling over western Europe, with only a weak warming over the Iberian Peninsula. The summer temperature change in the HadCM3 model is not statistically significant. This is in sharp contrast with the reconstruction of summer temperatures for the 1675–1715 and 1790–1820 periods, where a distinct cooling in western Europe is apparent, along with a warming in eastern Europe during the latter period.
5 Discussion and conclusions
In an effort to distinguish between a persistently negative NAO-type circulation and a reduction in the Atlantic MOC as the primary drivers of the cooling observed in the European LIA, two sets of experiments are performed. One in which a climate is simulated with a constant and negative NAO-index and one with a reduced overturning. Each set of experiment is done with an intermediate-complexity and a full complexity GCM. A comparison between the modelled temperatures and a reconstruction of LIA temperatures is at the basis of distinguishing between the two hypothesis.
This study shows that the hypothesis that LIA cooling is related to a circulation in the North Atlantic sector with a persistently negative NAO-index is unlikely to be valid. The model results do not convincingly reproduce the overall winter cooling see in Europe during the LIA period and persistence of this cooling into the summer season is weak.
The hosing experiments do not clearly support the hypothesis that a reduction in the MOC is the primary driver of LIA climate change. The hosing experiment with the ECBilt-Clio model shows a cooling which has similarities with the reconstructed temperatures for the LIA climate, but the largest cooling is off-shore. The reconstructions have their greatest cooling in eastern Europe which hints at a change in atmospheric circulation as an explanation for this pattern. Moreover, these experiments shows no similarity with the character of the reconstructed summer temperature signal either, which may have persisted from the preceeding winter, due to the conservative nature of surface ocean temperatures or soil moisture. However, the hosing explanation cannot be entirely discounted either. The experiments indicate that the feedback of changing SSTs due to the reduction in the overturning, lead to significant changes in atmospheric circulation with subsequent modification of the direct effect of changing ocean temperatures. The combination of the direct effect of cool oceanic surface temperatures and a modified atmospheric circulation remains a possible explanation.
A reconstruction of the winter atmospheric circulation for the 1790–1820 period clearly shows the anticyclonic anomaly in the northern North Atlantic (Lamb and Johnson 1959; van der Schrier and Barkmeijer 2005), which is in very good agreement with the anticyclonic circulation anomaly simulated here (Fig. 3a). However, in the present simulation the low pressure anomaly over the western North Atlantic and eastern US seaboard appears to be weaker and it does not correspond to the other aspects seen such as the anomalous cyclonic circulation over Central Europe. Nevertheless, the occurrence of high SLP anomalies with centers over northern Europe/Scandinavia is a recurrent feature of the atmospheric circulation in extreme cold periods of the LIA that has also been connected to the NAO variability (Luterbacher et al. 2002). A main conclusion from the present experiments with the ECBilt-Clio model is that the climatic response to a reduction in the strength of the MOC will directly impact temperatures in the North Atlantic sector, but it will also feedback on the atmospheric circulation. The latter will extort an additional modifying influence on European climate and its effect will be comparable, if not larger, than the direct effect of the MOC reduction.
In particular, Bjerknes (1965) argued that during the LIA, its climate was strongly influenced by the occurrence of anomalous high pressure centered in the North Atlantic, south of Iceland and anomalous low pressure off the eastern seaboard of the US.
In earlier experiments (van der Schrier and Barkmeijer 2005), the reconstructed winter atmospheric circulation for the 1790–1820 period was assimilated in the ECBilt-Clio model and that resulted in a strong resemblance of both winter and summer simulated temperatures with reconstructed temperatures. In the light of this earlier result, the present study suggests that the explanation of the coldness of this period hinges on the pattern of atmospheric circulation change, but not specifically a change in the NAO.
NCEP Reanalysis data were provided by the NOAA-CIRES Climate Diagnostics Center, Boulder, Colorado, USA, from their Web site at http://www.cdc.noaa.go. VP and GvdS were funded by the Netherlands Organisation for Scientific Research (ALW - NWO) and TK, TJO and KRB acknowledge support from the UK Natural Environmental Research Council (NERC). Funding for all authors is through the joint UK-NL RAPID Climate Change programme (To what extent was the Little Ice Age a result of a change in the THC? NE/C509607/1).
This article is distributed under the terms of the Creative Commons Attribution Noncommercial License which permits any noncommercial use, distribution, and reproduction in any medium, provided the original author(s) and source are credited.
- Bjerknes J (1965) Atmosphere-ocean interaction during the “Little Ice Age”. In: WMO-IUGG symposium on research and development aspects of longe-range forecasting WMO-No. 162. TP. 79, pp 77–88, Technical Note No. 66Google Scholar
- Jansen E, Overpeck J, Briffa KR, Duplessy J-C, Joos F, Masson-Delmotte V, Olago D, Otto-Bliesner B, Peltier WR, Rahmstorf S, Ramesh R, Raynaud D, Rind D, Solomina O, Villalba R, Zhang D (2007) Paleoclimate. In : Solomon S, Qin D, Manning M, Chen Z, Marquis M, Averyt KB, Tignor M, Miller HL (eds) Climate change 2007: the physical science basis. Contribution of working group I to the fourth assessment report of the intergovernmental panel on climate change. Cambridge University Press, Cambridge, pp 433–498Google Scholar
- Jones JM, Widmann M (2003) Reconstructing large-scale variability from palaeoclimatic evidence by means of Data Assimilation Through Upscaling and Nudging (DATUN). In: Fischer H, Kumke T, Lohmann G, Flöser G, Miller H, von Storch H, Negendank JFW (eds) The KIHZ project: towards a synthesis of Holocene proxy data and climate models. Springer, BerlinGoogle Scholar
- Kleinen T, Osborn T, Briffa K (2007) Assimilating NAO index states into GCMs: the case of the “Little Ice Age”. J Clim (submitted)Google Scholar
- Lamb HH, Johnson AI (1959) Climatic variation and observed changes in the general circulation. Part I and Part II. Geogr Ann 41:94–134Google Scholar
- Luterbacher J, Xoplaki E, Dietrich D, Rickli R, Jacobeit J, Beck C, Gyalistras D, Schmutz C, Wanner H (2002) Reconstruction of Sea Level Pressure fields over the Eastern North Atlantic and Europe back to 1500. Clim Dyn 18:545–561Google Scholar
- Shindell DT, Schmidt GA, Mann MA, Rind D, Wapple A (2001) Solar forcing of regional climate change during the Maunder Minimum. Clim Dyn 294:2149–2152Google Scholar
- Stouffer RJ, Dixon KW, Spelman MJ, Hurlin W, Yin J, Gregory JM, Weaver AJ, Eby M, Flato GM, Robitaille DY, Hasumi H, Oka A, Hu A, Jungclaus JH, Kamenkovich IV, Levermann A, Nawrath S, Montoya M, Murakami S, Peltier WR, Vettoretti G, Sokolov A, Weber SL, (2006) Investigating the causes of the response of the Thermohaline Circulation to past and future climate changes. J Clim 19:1365–1387CrossRefGoogle Scholar
- von Storch H, Cubach U, González-Rouco FJ, Jones MJ, Voss R, Widmann M, Zorita E (2000) Combining paleoclimatic evidence and gcms by means of data assimilation through upscaling and nudging (Datun). In: Proceedings of 11th symposium on global change studies, pp 1–4, AMSGoogle Scholar
- Widmann M, Goosse H, van der Schrier G, Schnur R, Barkmeijer J (2009) Using data assimilation to study extratropical Northern Hemisphere climate over the last millennium. Clim Past Discuss 5:2115–2156. http://www.clim-past-discuss.net/5/2115/2009/