Study of the ocean response to climate changes in XX

3. Results of Numerical Experiments

3.1. Response on Surface

[23] The sensitivity of a climate model to an increase in the concentrations of greenhouse gases is characterized primarily by an increase in globally averaged surface air temperature under CO₂ increased by 1% per year up to its doubling and by an equilibrium increase in surface air temperature under doubled CO₂ in the model with a homogeneous 50-m ocean layer. In the INM model, according to the data obtained from the experiment with CO₂ increased by 1% per year, warming is 1.57 K under doubled CO₂ (years 61-80) as compared to the same years of the control experiment. This value is close to the mean (1.61 K) over all models participating in similar experiments in the CMIP framework [Covey et al., 2000; Volodin et al., 2004]. The scatter in the data obtained with the models is sufficiently wide. For example, the minimum warming is 0.75 K, and the maximum warming is 3.77 K. However, the model experiments in the CMIP framework were carried out mainly in 2000-2001.

[24] According to the latest data, for 12 models that participated in such an experiment in 2004, the average warming is 1.81 K, the minimum warming is 1.46 K, and the maximum warming is 2.2 K (G. A. Meehl, 2005, private communication). This implies that the average warming in the models is slightly increased and the scatter in model data is significantly decreased. In this experiment, the INM model yields a warming value that is slightly smaller than the mean over all models. For these 12 models, the value of equilibrium warming in the model with the upper ocean layer under doubled CO₂ is known. The average warming is 2.95 K, and, for different models, its value varies from 2.10 to 3.95 K. For the INM model, the equilibrium warming is 2.10 K, which is the smallest value among the models under consideration. Comparison between the equilibrium and non-equilibrium responses of the model suggests that, in the INM model with the entire ocean, a thinner ocean layer (as compared to other models) is responsible for warming. This reduces the disagreement between the INM model and the mean over all models in the non-equilibrium experiment. The reasons why the sensitivity to doubled CO₂ in the INM model is significantly lower than its mean over all models are analyzed in [Volodin et al., 2004]. The main reason is an increase (under warming) in the amount of lower cloudiness under the conditions of a more frequent formation of inversion. In [Volodin et al., 2004], the probable sensitivity of the real climate system is estimated.

[25] As compared to the previous version of the INM coupled model [Volodin and Diansky, 2003], where the warming in the model with the entire ocean under CO ₂ increased by 1% per year was 1.0 K, the model's sensitivity was increased by a factor greater than 1.5. This is primarily due to the introduction of an interactive model of sea ice. As compared to analogous years of the control experiment, the warming in the model with the entire ocean under doubled CO₂ (experiment 2CO₂, years 201-220) at the end of calculation is 2.1 K, which coincides with the value of equilibrium warming in the model with a 50-m ocean layer. By this moment, warming at the surface can be considered to be almost equilibrium. The fact that the values of the equilibrium warming in the models with the entire ocean and a 50-m ocean layer almost coincide suggests that it is possible to study the equilibrium sensitivity with the aid of the models with a 50-m ocean layer. At the end of the experiment with quadruple CO₂ (4CO₂, years 271-290), the warming is 4.1 K as compared to analogous years of the control experiment. In this case, the warming at the surface can also be considered almost equilibrium. Comparison with the warming in the 2CO₂ experiment shows that the dependence of the equilibrium warming value on the logarithm of the CO₂ concentration increase is almost linear.

Figure 2

[26] Let us consider the temporal behavior of the globally averaged air temperature at the surface in the control experiment, as well as in the experiments modeling the climate of the 20th-22nd centuries (Figure 2). The mean air temperature for the control experiment is 12.6^o C, which is 1.4^o C below that obtained from the 1961-1990 observational data [Jones et al., 1999] and approximately 0.8-1.0^o C below that obtained for the second half of the 19th century, for the conditions under which the control experiment was performed. In the control experiment, the temperature trend is positive; its value does not exceed 0.2 K over 330 years of the experiment. In the experiment simulating the climate of the 20th century, already by the mid-20th century, warming is noticeable as compared to the control experiment. The temperature rise reaches 0.7-0.8 K, which is close to the warming observed (0.6-0.7 K). In the data obtained from the control experiment, there is no time interval within which the warming would be equally significant. It follows that the warming of the 20th century is most likely caused not by the atmosphere-ocean system's variability, but by external (with respect to this system) forcing. The same conclusion is drawn from the results of other models [Solomon et al., 2007].

[27] According to our model, during the 21st century, even under all forcing fixed at a level of the year 2000 (experiment 21), warming is expected to be approximately 0.6 K owing to a large thermal inertia of the ocean. The rise of temperature in experiments B1, A1B, and A2 is more noticeable (as compared to the year 2000) and reaches 2, 3, and 5 K, respectively, by the end of the 22nd century according to model data. In experiment A2, a slight additional warming during the years 2145-2165 is due to variations in the time step in the atmospheric block of the model.

[28] The temporal behavior of the globally averaged precipitation is shown in Figure 2b for the experiments carried out. In the control experiment, the amount of precipitation is about 2.80 mm day^-1. In experiment 20, the amount of precipitation is increased to 2.84 mm day ^-1 by the year 2000. By the end of experiment 21, the amount of precipitation is increased to 2.89 mm day^-1, and, by the end of experiments B1, A1B, and A2, the amount of precipitation is 2.95, 2.99, and 3.10 mm day^-1, respectively. In the experiments modeling warming, the amount of precipitation is increased mainly in proportion to the increase of the mean air temperature at the surface.

Figure 3

[29] Let us compare the temperature variation in experiment 20 with respect to the control experiment and the 1871-2000 temperature variation estimated from observational data [Jones et al., 1999] (Figure 3). The model adequately reproduces the magnitude of warming in the 20th century, which is about 0.6-0.7 K according to observational data. The features such as the warming in 1940-1950 and its slowing down in 1960-1970 are also reflected in the model. These features can be explained by the maximum solar constant and minimum volcanic aerosols in 1940-1950 and by the minimum solar activity and maximum volcanic aerosols in 1960-1970 (Figure 1). However, it follows from the temperature variation in the control experiment (Figure 2) that these features can be related to the natural variability of the atmosphere-ocean system. In the literature, it is also shown that these features may be due to both variations in the solar constant and volcanic-aerosol concentrations [Meehl et al., 2004] and the internal variability of the atmosphere-ocean system [Broccoli et al., 2003].

[30] Let us compare the temperature variation at the end of the 21st century according to the INM model with the data of other models from [Solomon et al., 2007]. Under scenario B1, the warming in 2091-2110 (with respect to 1991-2010) is 1.7 K in the INM model. According to the data of all models, under this scenario, the warming varies from 1.1 to 2.5 K (the mean value is 1.8 K). For scenario A1B, the warming is 2.3 K according to the INM model data and 1.8-3.8 K (the mean value is 2.8 K) according to the data of all models. For scenario A2, the warming is 3.2 K according to the INM model data and 2.5-4.7 (the mean is 3.6 K) according to the data of all models. For the three scenarios, the warming in the INM model is somewhat less than its mean value over all models, a finding that is in agreement with the data obtained from the experiment on simulation of a non-equilibrium response to doubled CO ₂. In the 22nd century, despite the conservation of external forcing at the level of the year 2100, the warming persists and is equal to 0.3 K for scenario B1, 0.5 K for scenario A1B, and 1.5 K for scenario A2. Slight temperature variations during the 22nd century in experiments A1B and especially B1 are due to the fact that the concentration of methane, whose contribution to the greenhouse effect is significant, is decreased in the second half of the 21st century according to these scenarios (Figure 1). Therefore, despite the increase in CO₂, the climate system by the year 2100 is not very far from equilibrium. At the same time, under scenario A2, increases in the concentrations of CO₂ and methane occur up to the very end of the 21st century; therefore, by the year 2100, the climate system is far from equilibrium, and the warming in the 22nd century under scenario A2 is greater than under scenarios A1B and B1.

Figure 4

[31] Temperature variations under a global warming are non-uniform over the surface. Figure 4 gives the difference of air temperatures at the surface over the years 2151-2200 of experiment A1B and over the years 1951-2000 of experiment 20. The warming is maximum in the Arctic and reaches 10 K. For the territory of Russia, the rise in air temperature is 5-7 K. Over the rest of the continent, the rise in temperature is 3-5 K. The lowest rise in air temperature (2-3 K) occurs over the tropical and southern oceans. Such a distribution of warming is characteristic of most of the models and is close to the results obtained by averaging the data of all models participating in the CMIP. In the INM model (Figure 4), variations in precipitation under warming are also characteristic of most of the models. The precipitation is increased (by 20-40%) in the middle and high latitudes of both hemispheres and also over the tropical Pacific. The precipitation is decreased over most of the subtropics and the tropical Atlantic.

Figure 5

[32] The peaks of warming observed in the Arctic and to the east of the Antarctic Peninsula are due to an intensive sea-ice melting at the end of summer. Figure 5 shows the area of sea ice for the Northern Hemisphere in March and September for the control experiment and experiments 20, B1, A1B, and A2. In March, the ice area remains almost unchanged for the control experiment; only high-frequency oscillations occur, and a small negative trend is noted. In experiments B1, A1B, and A2, the ice area is decreased in March. By the end of the 22nd century, the decrease reaches 20, 30, and 50%, respectively. In September, variations in the area of sea ice are even wider. Even by the end of the 20th century, the ice area in experiment 20 is decreased by 20-25% as compared to the control experiment. In the 22nd century, there is no ice in the Arctic in experiment A2; ice remains until September only in some years in experiment A1B; and, in experiment B1, ice remains, but its area is only 10-20% of that in the control experiment.

[33] According to observational data [Waple et al., 2004], at the end of the 20th century, in July-September, the area of the Arctic ice was 20-25% less than that in the mid-century and, in January-March; the ice area almost has not changed over the past 50 years. This is in agreement with the model data given in Figure 5.

Figure 6

[34] In Antarctica, the ice area is not decreased as rapidly as in the Arctic (Figure 6). In September, at the end of the 22nd century, the decrease of ice area is about 20, 30, and 40% for experiments B1, A1B, and A2, respectively. In March, at the end of the 22nd century, in experiments B1, A1B, and A2, the ice area is 50, 30, and 5%, respectively, of that in the control experiment. In addition, in Antarctica, a significant variability with amplitude of 20-30% of the mean area is observed for March in the control experiment.

[35] On the basis of the data obtained in experiments 20, 21, B1, A1B, and A2, the response of sea-level variation to variations in external forcing specified in these experiments is calculated. It was assumed that the sea-level variation involves the variations in water density (thermal expansion) and variations in the balance of accumulation and melting of ice in Greenland and Antarctica. To calculate the thermal component, it was assumed that, in 1871, the sea level was equal to 0 m, and the distribution of depth corresponded to that specified in the ocean model. Then, the variation in the mean water density resulting from variations in temperature and salinity and the related variation in sea level were calculated in accordance with the equation of state used in the ocean model.

[36] The contribution made to sea-level variations by the balance of ice in Antarctica and Greenland was calculated in the following way. The amount of precipitation in the form of snow S that had fallen out on the continental ice and the amount of melted ice M were calculated and averaged over 130 years of the control experiment (1871-2000). It was assumed that, in the control experiment, the continental ice is in equilibrium and the amount of ice R broken away from Greenland and Antarctica, on average over 130 years, exactly corresponds to its increase owing to the difference between its accumulation and melting: R=S-M. It was assumed that the rate of ice runoff to the ocean R calculated in such a way is the same for any year and for any numerical experiment. The variation in sea level because of variations in the continental-ice balance Ç was calculated as B=S-M-R. The total sea-level variation is the sum of the sea-level variation caused by thermal expansion and the variation of the balance of continental ice.

Figure 7

[37] Figure 7 shows the sea-level variation for experiments 20, 21, B1, A1B, and A2 as compared to the control experiment. According to model data, during the 20th century, the sea-level rise was 4 cm, which is approximately two times less than the estimate obtained from observational data [Solomon et al., 2007]. The underestimation may be due to the fact that the model does not take into account the melting of small mountain glaciers and that the ocean layer involved in global warming is thinner in the model than in the actual climate system. In experiment 21, the sea level continues to rise and reaches 11 cm by the end of the 21st century. By the end of the 22nd century, the sea-level rise for scenarios B1, A1B, and A2 is 27, 36, and 48 cm, respectively. Owing to a large thermal inertia of the oceans, the sea level continues to rise in the 22nd century, when external forcing are fixed, at a rate almost identical to that in the 21st century. According to model data, the major contribution to the sea-level rise is made by thermal expansion. The contribution of the variation in the balance of ice in Greenland and Antarctica does not exceed 2-3 cm. Moreover, for experiments 21, B1, and A1B, this contribution is mainly negative; i.e., at a comparatively slight global warming, the increase in snowfall exceeds that in the amount of melted snow.

[38] Comparison of the sea-level rise in the INM model with the data obtained with other models and given in [Solomon et al., 2007] shows that, among all models, the INM value is one of the lowest model values. Thus, for all models, the variation in the sea level in 2100 (as compared to 1990) is between 8 and 56 cm for scenario B1, between 13 and 69 cm for scenario A1B, and between 16 and 74 cm for scenario A2. In the INM model, these values are equal to 11, 16, and 18 cm for experiments B1, A1B, and A2, respectively. However, the last model comparison, whose results are to be presented in the fourth IPCC report being prepared, shows that the scatter in model data on this parameter is decreased mainly owing to a decrease in the upper bound of its estimate (J. Gregory 2005, private communication). According to the data obtained with 11 models, the sea-level rise in 2100 as compared to 1990 may vary from 9 to 27 cm for scenario B1, from 12 to 36 cm for scenario A1B, and from 15 to 36 cm for scenario A2. This comparison shows that, in the 20th century, the sea-level rise varies from 0 to 7 cm according to the data obtained with all models and is 4 cm according to the INM model. Thus, according to the data of this comparison, the sea-level rise in the INM model is within the same interval. We note that the scatter in the sea-level rise values among the models is wider than that for each individual model but for different scenarios.

3.2. Deep Ocean Response

[39] Here, we will consider the integral characteristics of ocean response. The ocean model as a component of the coupled model has a small trend in deep-sea layers, which is manifested in all characteristics. This trend is due to the features of the ocean model and to some imbalance of the heat and freshwater fluxes on the ocean surface. An incomplete balance between the river runoff and the precipitation is also responsible for this trend. Therefore, to reveal the response of the ocean to global warming, we shall treat its response as the averaged (over the past 30 years) differences in its characteristics obtained from the experiment under the most probable scenario A1B and the control experiment. Below, we shall associate these differences with the ocean responses to global warming.

Figure 8

[40] An important factor maintaining the Earth's climatic equilibrium is the meridional heat transfer (MHT) in the ocean. Figure 8a shows the MHT (averaged over the past 30 years of the control experiment) for the global ocean, the Atlantic, and the Pacific plus Indian oceans. Positive and negative MHT values correspond to northward and southward transfers, respectively. The latitude MHT distribution agrees well with the distribution averaged over all CMIP models [Covey et al., 2000; Solomon et al., 2007]. For comparison, the MHT values from [Macdonald and Wunsh, 1996] are given for the global ocean (circles), the Atlantic (triangles), and the Pacific plus Indian oceans (diamonds), which are in agreement with sparse direct observational data [Hall and Bryden, 1982]. In the new version of the coupled INM model, the MHT amplitudes were increased, which brought model values closer to observational data. For example, the global MHT was increased approximately by 0.5 PW for both Northern and Southern hemispheres as compared to the previous version of the coupled model [Diansky and Volodin, 2002]. The maximum of the calculated global mean annual MHT is about 1.8 PW for the Northern Hemisphere and -1 PW for the Southern Hemisphere at latitudes of about 10^o.

[41] The mean annual heat transfer in the Atlantic (Figure 8b, dashed curve) reaches its maximum 1.2 PW at 30^o N and then decreases with its turning eastward and with a decreased intensity of the Gulf Stream and North Atlantic current. In the Southern Hemisphere, the heat transfer is determined mainly by the processes occurring in the Pacific and Indian oceans. In these oceans, the joint heat transfer to the south (Figure 8b, dotted curve) reaches its maximum 1.4 PW at 15^o S. It should be noted that the heat transfer is increased in the Northern Hemisphere mainly due to the increased MHT in the Atlantic and in the Southern Hemisphere owing to heat transfer in the Pacific and Indian oceans.

[42] MHT intensification in the ocean is a positive effect of the modernized coupled model. Earlier, its value was among the smallest for the CMIP models. This intensification is caused by an increased intensification of meridional circulation in the upper ocean layers, which, in turn, is increased owing to both the strengthening of trade winds in the coupled model and the change (in numerically implementing the ocean model) to the C grid, which has a lower dissipation than the B grid used in the previous version of the model.

[43] Figure 8b shows the MHT variations caused by climate changes under scenario A1B for the global ocean, the Atlantic, and the Pacific plus Indian oceans. The character of the global response in the MHT formed so as to decrease the intensity of the global MHT in the control experiment. Figure 8b shows that the response in the global MHT is formed through the responses in the Atlantic and in the Pacific and Indian oceans, which are of quite a different character. In this case, the decrease in the MHT is caused mainly by decreased MHT in the tropics of the Pacific for the Northern Hemisphere and in the tropical Atlantic for the Southern Hemisphere.

Figure 9

[44] The spatial structure of the MHT response can be formed by the responses in the meridional circulation and/or the meridional distribution of ocean temperature. With the exception of a wavelike response in the MHT in the equatorial region, major variations in the MHT are caused by the corresponding variations in meridional circulation in the Atlantic and Pacific oceans (Figure 9). This figure shows the responses for the stream function of the zonally averaged circulation in the Atlantic and in the Pacific and Indian oceans against the background of the zonally averaged circulation itself.

[45] It follows from Figure 9 that the positive response in the MHT in the tropics of the Southern Hemisphere (Figure 8b) is caused by a positive cell of the response of circulation in the upper layer of the South Atlantic with its center at approximately 15^o S and a depth of 400 m, and the negative response in the MHT in the tropics of the Northern Hemisphere is caused by positive cell of the response of meridional circulation in the upper layer of the Pacific with its center approximately 15^oN and a depth of 300 m. Here, should be noted that the formation of both the MHT itself and its response is mostly affected by meridional circulation in the upper ocean layer because the near-surface sectors of circulation cells transport the near-surface water, which is warmer than the water transported by their deep-water sectors, and, therefore, determine heat transfer to a greater extent. Meridional circulation and its variability in the deep-water ocean layers have a weaker influence on the MHT because this circulation transports the water of deep ocean layers with a weak vertical stratification. Therefore, the extensive deep-sea negative cell of the response of meridional circulation in the Atlantic Ocean does not significantly affect the MHT. This is also true for meridional circulation in the Pacific and Indian oceans.

[46] A sufficiently significant peak in the response of meridional circulation in the Southern Ocean is caused (as shown in Figure 9b) mainly by the response in the sector of the Pacific and Indian oceans. This response corresponds to an increase of heat transfer to Antarctica, a phenomenon that must lead to an additional factor of sea-ice melting in this region.

[47] A wavelike response of the MHT in the equatorial region is caused by a significant rise in the OST in this region of the Pacific Ocean (see Figure 4a), when the warm water of this region is transported by the near-equatorial cells of meridional circulation.

[48] Here, for lack of room, we do not show the response of ocean salinity. Only note that the response of the sea surface salinity is determined primarily by a variation in the surface freshwater balance, a circumstance that is mainly due to variations in precipitation (see Figure 4b). In this case, the sea surface salinity decreases in the Pacific and mainly increases in the Atlantic (especially in the tropics). Such a redistribution of surface freshwater fluxes must lead to the rise of the sea level in the Pacific and to its decline in the Atlantic. The salinity response is reflected in the meridional freshwater transport, which is formed by both the response of meridional circulation and salinity variations mainly in the upper ocean layers.