Impact of atmospheric CO 2 doubling on the North Pacific Subtropical Mode Water

Transcription

1 GEOPHYSICAL RESEARCH LETTERS, VOL. 36, L06602, doi: /2008gl037075, 2009 Impact of atmospheric CO 2 doubling on the North Pacific Subtropical Mode Water Hyun-Chul Lee 1,2 Received 19 December 2008; revised 12 February 2009; accepted 18 February 2009; published 20 March [1] In order to investigate responses of the North Pacific Subtropical Mode Water (NPSTMW) to climate change, the impact of atmospheric CO 2 doubling with 1% annual increase is examined using a coupled climate model (GFDL CM2.1). Under the CO 2 forcing, the surface waters in the formation region of NPSTMW and the core layer of NPSTMW become warmer and freshener. The total volume of NPSTMW increases by about 40% due to warming. The inter-annual and decadal variability of NPSTMW is significantly correlated with the variability of the Kuroshio heat transport, and less correlated with variability of sea surface temperature and winter monsoon index in comparison with the control run. Citation: Lee, H.-C. (2009), Impact of atmospheric CO 2 doubling on the North Pacific Subtropical Mode Water, Geophys. Res. Lett., 36, L06602, doi: /2008gl Introduction [2] North Pacific Subtropical Mode Water (NPSTMW) with a characteristic temperature range of 15 C 19 C, is located between the seasonal and permanent thermoclines in the North Pacific [Bingham, 1992; Yasuda and Hanawa, 1997; Taneda et al., 2000]. NPSTMW is formed by surface cooling and associated vertical mixing processes in the winter season to the south of the Kuroshio and its extension region, and advected to the southeastward by the subtropical gyre [Suga and Hanawa, 1990; Bingham, 1992]. Yasuda and Hanawa [1997] and Taneda et al. [2000] showed that NPSTMW formation by surface cooling in winter is closely related to variability of the wintertime wind stress field and Ekman divergence associated with the intensification of the Aleutian low. [3] Because the formation process occurs through air-sea interaction, the inter-annual and decadal variability of NPSTMW is greatly affected by climate change. It has been reported that the inter-annual temperature variability of the NPSTMW could be controlled by the spinning-up of the Pacific subtropical gyre [Yasuda and Kitamura, 2003]. Also it has been suggested that the variability of ventilation of subtropical mode water modulates the decadal climate fluctuations in the tropical Pacific [Gu and Philander, 1997]. [4] The oceanic responses to global climate change with an atmospheric CO 2 increase at rate of 1% per year have been studied by Stouffer et al. [2006] and Manabe et al. [1991]. Global warming decreases the rate of ventilation in the North Atlantic and the Southern Ocean due to the 1 Geophysical Fluid Dynamics Laboratory, NOAA, Princeton, New Jersey, USA. 2 High Performance Technology, Inc., Reston, Virginia, USA. Copyright 2009 by the American Geophysical Union /09/2008GL increase of the surface stratification [Stouffer et al., 2006; Gnanadesikan et al., 2007]. Through the wintertime ventilation process, the NPSTMW exchanges buoyancy flux with atmosphere. This interaction is important to inter-annual and decadal variability in the North Pacific. The oceanic response of the North Pacific subsurface water mass to global warming, however, has not been clarified. This study investigates how climate change from increasing atmospheric CO 2 can impact the spatial distributions, and interannual variability of NPSTMW in a coupled climate model. 2. Numerical Model [5] The numerical climate model applied here is the Geophysical Fluid Dynamics Laboratory (GFDL) coupled climate model (CM2.1). The GFDL CM2.1 is a fully coupled climate model using ocean, atmosphere, land and sea ice components without flux adjustment [Delworth et al., 2006]. [6] The Ocean model component is the GFDL Modular Ocean Model (MOM4) [Griffies et al., 2003]. The horizontal grid resolution is one degree longitude and latitude south of 65 N (1/3 the latitudinal resolution of in Tropics). The vertical resolution is 50 levels. The sea ice component is the GFDL Sea Ice Simulator which uses three vertical layers with a modified Semtner three-layer scheme for the thermodynamics [Winton, 2000]. The atmospheric model component (GFDL AM2.1) uses the finite volume dynamical core [Lin, 2004] with a terrain-following Lagrangian control-volume discretization. The model resolution is 2 by 2.5 (latitude and longitude) using 24 vertical levels. The land component considers the land cover type by the potential natural vegetation classification [Milly and Shmakin, 2002]. The components are coupled using the GFDL Flexible Modeling System. A detailed description of the model formulation and characteristics are provided by Delworth et al. [2006]. [7] In order to investigate the impact of an idealized atmospheric CO 2 doubling on oceanic conditions, 200 year model results from a control experiment and a CO 2 doubling experiment were compared. The control run uses 1990 atmospheric conditions, with CO 2 and other greenhouse gases, solar constant and aerosols held at the 1990 level, that is the same configuration as one of Stouffer et al. [2006]. The CO 2 doubling experiment (hereafter 2CO 2 ) is identical with the control run, except the CO 2 concentration is increased at the rate of one percent per year and thus doubled after 70 years. After the initial doubling, the CO 2 concentration remains constant. 3. Results [8] As shown by Stouffer et al. [2006], the surface air temperature increase in response to the CO 2 doubling is L of5

2 Figure 1. Fields of winter time average (DJF, year ) of (a) difference of SST ( C) between 2 CO 2 and control experiment (shade) and the water age of 2CO 2 at 300 m depth (contour), (b) difference of SSS with unit of psu (shade), the Kuroshio axis of the local current speed maximum of control run (black line) and 2CO 2 (red line), (c) Montgomery potential (10 3 JKg 1 ) on the isopycnal surface (s q = 25.5) of 2CO 2 (shade) and its difference from the control run (contour), and (d) planetary scale potential vorticity (10 10 m 1 s 1 ) in the layer between s q = and 25.5 (shade) and its difference from the control run (contour). larger in high latitude than in the equatorial region of the North Pacific. In the northern Bering Sea, the warming is about 8 C in the winter season (DJF). The sea surface temperature (SST) in the northwestern Pacific also increases due to the atmospheric CO 2 doubling. Figure 1a shows the difference of annually averaged winter SST due to the CO 2 doubling. The SST of the northwestern Pacific increases by C in the Kuroshio extension region (Figure 1a). The water age distribution at 300 m depth shows a newly ventilated water mass locates south of the Kuroshio extension. Here, the water age is the elapsed time since the water mass was out of contact with the sea surface. This distribution well agrees with observational results of Yasuda and Kitamura [2003]. [9] Within 2CO 2, the annually (year ) and zonally (140 E 130 W) averaged westerly wind stress over the North Pacific (30 N 55 N), increases its strength by 3.1% over the control run. Due to the increase of westerly wind stress, the southward Ekman transport increases by 0.3 Sv (1Sv = 10 6 m 3 s 1 ). However, the strength of anticyclonic wind stress curl over the subtropical gyre (17 N 43 N) actually decreases by 5.7%, and the southward Sverdrup transport decreases by 2.8 Sv. Based on the linear dynamics, the total change of southward Ekman and Sverdrup transports would decrease by 2.5 Sv due to the wind field change of the CO 2 forcing. [10] The sea surface salinity (SSS) in the northwestern Pacific is decreased by the atmospheric CO 2 doubling (Figure 1b). This freshening is mainly induced by increased precipitation in the western equatorial Pacific [Stouffer et al., 2006]. In the western equatorial Pacific, the boreal winter precipitation increases in the western equatorial Pacific by 2 Kg m 2 day 1 and in the high latitude region north of 40 N by 0.5 Kg m 2 day 1. Over the subtropical gyre, precipitation is decreased by 0.6 Kg m 2 day 1. The freshening anomalies of the western equatorial Pacific are advected to high latitudes by the Kuroshio, and SSS in the formation region of the NPSTMW decreases by 0.15 psu. [11] Geostrophic streamlines can be illustrated by the Montgomery potential defined as p/r + gz, where p is pressure, r is water density, and g is gravitational acceleration. As shown in Figure 1c, the pattern of Montgomery potential on the s q = 25.5 Kg m 3 isopycnal surface is consistent with the analytical solution of the ventilated thermocline theory [Pedlosky, 1996]. It is shown that there is a pool region in the west and a shadow zone in the east. In 2CO 2, the winter outcropping region of the isopycnal surface moves to the north and the Montgomery potential 2of5

3 Figure 2. Seasonally averaged fields (from May to December, year ) along the core layer of NPSTMW for (a) the control run temperature (shading, C) and depth of the core layer (contour, meter), (b) 2CO 2 temperature (shading, C) and depth of the core layer (contour, meter), (c) control run salinity (shading, psu) and current vector (ms 1 ), and (d) 2CO 2 salinity (shading, psu) and current vector (ms 1 ). increases. The increase of the Montgomery potential in the ventilated region is much larger than in the shadow zone, and therefore the gradient of geostrophic streamline between ventilated and shadow region increases. [12] There is nearly uniform potential vorticity (PV) in the most central region with low PV (Figure 1d). The low PV tongue occurs in the boundary between the ventilated and shadow region. According to Pedlosky [1996], this homogenized pattern is produced by lateral mixing across the boundary. Under the CO 2 doubling scenario, the PV in this layer decreases, and the vertical density stratification weakens. In comparison to the control run, PV is more homogeneous, due to the increasing gradient of geostrophic stream function (Figure 1c) and lateral mixing across the boundary. Because of the homogenization and the southward extension of the pool region, the largest decrease of PV occurs near the boundary with the shadow zone. [13] Observational studies [Hanawa and Kamada, 2001; Taneda et al., 2000; Yasuda and Hanawa, 1997] have defined the core layer of NPSTMW by the layer with minimum vertical temperature gradient with in the temperature range of C. The same definition of the NPSTMW in the depth range of m is used here to compare with observational results. Figure 2 shows that the NPSTMW core layer expands to the north and east, and the total volume of NPSTMW calculated from the water mass C and m depth, increases by 38.3% in comparison to the control run. [14] In2CO 2, the 200 year averaged annual volume and salt transport of the Kuroshio are slightly decreased by 1.3% and 1.5% respectively, but the Kuroshio heat transport (KHT) increases by 6.1% in the control run. These changes of average are statistically significant at 95% significance level. Here, the heat transport is calculated by r o C p RR D(U T) dydz, where C p is specific heat. U and T is zonal velocity and water temperature and integrated between 29.5 N and 32 N at 132 E. This increase of the KHT causes warming in both the surface water and the core layer in the northwestern Pacific and results in an expansion of the core layer of the NPSTMW. In comparison with control run, the core layer of 2CO 2 becomes freshened by 0.2 psu in the region southeast of the Japan. As a result, the NPSTMW becomes warmer and fresher in response to the atmospheric CO 2 doubling. [15] The inter-annual and decadal variability of NPSTMW is controlled by atmospheric decadal variability and Kuroshio transport. NPSTMW is formed in the winter season and the winter SST over the ventilation area of the mode water is an important control factor of the air-sea heat flux for the NPSTMW formation process [Taneda et al., 2000; Hanawa and Kamada, 2001]. 3of5

4 Figure 3. (a) Correlation between CLT and SST of control run (solid) and 2CO 2 (dashed) and (b) correlation between CLT and Kuroshio heat transport of control run (solid) and 2CO 2. The CLT is averaged from May to December, and SST is winter time average (DJF). Kuroshio heat transport is an annual average. CLT and SST are spatially averaged over E and N, and Low-pass filtered (10-year). Variabilities (X s ) of CLT and SST were standardized by the average (X ) and standard deviation (s), X S =(X X )/s. [16] In control run, the CLT of the NPSTMW is well correlated with the decadal variability of SST averaged over the area. The correlation between CLT and SST is 0.90 at zero time lag (Figure 3a). In response to CO 2 forcing, the correlation between CLT and SST weakens to This suggests that warming and freshening reduces the influence of decadal atmospheric variability on the NPSTMW formation. [17] The KHT plays an important role in the formation of NPSTMW and the variability of the CLT [Hanawa and Kamada, 2001; Taneda et al., 2000]. In both the control and 2CO 2, the maximum correlation between the variability of the NPSTMW CLT and the KHT are 0.63 and 0.54, respectively (Figure 3b). The decadal variability of NPSTMW CLT is significantly correlated with the KHT in both results. In 2CO 2, the influence of atmospheric interaction on the decadal variability of CLT is reduced compared to the control run. [18] The East Asian winter monsoon index (EAWMI) is defined as the difference of zonal wind speed at the 300 hpa level between two regions U 300Pa (110 E 170 E and 27.5 N 37.5 N), and U 300Pa (80 E 140 E and 50 N 60 N). When EAWMI increases, the Aleutian low deepens and the westerly wind in the Kuroshio region increases [Jhun and Lee, 2004]. [19] It is noticeable that in both runs negative linear correlation mainly occurs in the Kuroshio extension region (Figures 4a and 4b). Negative coefficient of linear correlation means cooling (warming) of the core layer of NPSTMW and strengthening (weakening) of the East Asian winter monsoon. The inter-annual variability of NPSTMW is modulated by the variability of the East Asian winter Figure 4. Linear correlation coefficients with zero time lag between CLT and (a) EAWMI of the control run, (b) EAWMI of 2CO 2, (c) KHT of the control run, and (d) KHT of 2CO 2. Linear correlation coefficients displayed when significance exceeds 95% level. 4of5

5 monsoon and the winter atmospheric conditions over the formation region. This suggests that the NPSTMW is formed by winter convection in this region, and this result is consistent with the observations [Suga and Hanawa, 1990; Bingham, 1992]. In comparison with control run, the CO 2 forcing weakens the correlation between CLT and EAWMI, and the influence of winter atmospheric condition on the inter-annual variability of NPSTMW also decreases. [20] Figure 4c shows that in the control run the correlation between KHT and the variability of CLT is high in the southeastern corner of the core layer and small in the Kuroshio extension region. This may suggest that KHT in the upper layer does not directly affect NPSTMW just below the Kuroshio extension region. Instead of this, the influence of the KHT is anti-cyclonically advected by the gyre after subduction and appears in the southeastern region of the core layer. [21] By the CO 2 forcing the correlation between CLT and the KHT is strengthened in a wider region along the eastern rim of the core layer (Figure 4d). In the formation region, the inter-annual variability of the KHT is significantly correlated with the variability of NPSTMW. The maximum correlation occurs in the southeastern region by the same process as the control run, and the variability of CLT in this region would be a dynamical response of the subtropical gyre to the wind field change. The model results show that by the atmospheric CO 2 doubling and global warming, the decadal variability of NPSTMW is significantly affected by the KHT along the eastern boundary of the core layer. 4. Summary and Discussion [22] The impact of the global warming on the NPSTMW is investigated by using the GFDL coupled climate model, CM2.1, forced by an idealized doubling of atmospheric CO 2. By CO 2 doubling, the sea surface layer becomes warmer and freshener in the Kuroshio extension region at which the NPSTMW is formed in the winter. In response to increase CO 2, the NPSTMW extends to the north and east, and the total volume of NPSTMW increases about 40%. This is mainly due to increased KHT. The correlation of decadal variability between CLT and SST in the Kuroshio extension region weakens under the increased CO 2 forcing. The variability of the NPSTMW is mainly determined by the variability of the KHT, and the influence of the atmospheric variability is reduced by the CO 2 doubling. [23] The NPSTMW is formed by the convectional mixing process in winter. By the 2CO 2 forcing, the surface layer in the northwestern Pacific is warming and freshening (Figure 1), and the static stability of the upper layer increases in this region. This can reduce the convectional mixing in winter and the influence of atmospheric variability for a formation of NPSTMW. The linear correlation between CLT and EAWMI consistently shows that CO 2 forcing weakens the impact of the decadal atmospheric variability on the NPSTMW. The decadal variability of NPSTMW is significantly modulated by the KHT and the variability of the subtropical circulation. [24] Acknowledgments. I thank R. Stouffer for invaluable comments and also A. Gnanadesikan, A. Rosati and T. Delworth for many helpful discussions. Many constructive comments and suggestions from anonymous reviewers greatly improved this manuscript, and I am deeply grateful for many helps. References Bingham, F. M. (1992), Formation and spreading of subtropical mode water in the North Pacific, J. Geophys. Res., 97, 11,177 11,189. Delworth, T. L., et al. (2006), GFDL s CM2 global coupled climate models. Part I: Formulation and simulation characteristics, J. Clim., 19, Gnanadesikan, A., J. L. Russell, and F. Zeng (2007), How does ocean ventilation change under global warming?, Ocean Sci., 3, Griffies, S. M., M. 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Kitamura (2003), Long-term variability of the North Pacific Subtropical Mode Water in response to spin-up of the subtropical gyre, J. Oceanogr., 59, H.-C. Lee, Geophysical Fluid Dynamics Laboratory, NOAA, Princeton, NJ , USA. (hyun-chul.lee@noaa.gov) 5of5