Honours Thesis: Impacts of Latitude Shifts in the Southern. Ocean Westerly Winds on Past and Present. Climates

Transcription

1 Honours Thesis: Impacts of Latitude Shifts in the Southern Ocean Westerly Winds on Past and Present Climates StéphanieDupré Department of Physical Oceanography/Meteorology School of Mathematics University of NSW Australia Supervisor: Dr. Matthew H. England October 28, 24

2 1 Abstract The climate response to latitude shifts of the Subpolar Westerly Winds (SWWs) in the Southern Hemisphere is analysed in a coupled climate model of intermediate complexity. The control experiment is a present-day model run forced by the National Centers for Environmental Prediction and National Center for Atmospheric Research (NCEP-NCAR) Reanalysis climatological wind fields. We then run two experiments out of equilibrium, one with a poleward shift, and the other with an equatorward shift, of 5.4 in latitude. Of particular interest is the response of ocean circulation, water mass changes, climate, and ocean carbon uptake to the shifted winds. This has relevance to the interpretation of past and future climates, as wind shifts are projected under enhanced greenhouse forcing, and the wind axis appears to have oscillated in paleoclimate records. We find an increase in the formation of Antarctic Intermediate Water (AAIW) and a decrease in Subantarctic Mode Water (SAMW) production for a poleward wind shift, in the South Pacific and South Indian Oceans. The opposite scenario occurs for the equatorward wind shift; namely increased SAMW formation and reduced production of AAIW. The solubility pump of globally dissolved inorganic carbon (DIC) is enhanced (reduced) by 12 ppm (23 ppm) due to a southward (northward) latitude shift in the SWWs. This is due to the winds control of overturning in the Southern Ocean; as the wind maximum moves poleward, cooler waters are advected in the surface layer, altering convective processes in the ocean s interior. The opposite occurs in response to an equatorward wind shift. Other important adjustments are noted in regional ocean and climate patterns, confirming that subtle latitude shifts in the Subpolar Southern Hemisphere westerlies can significantly alter our climate system.

3 Contents List of Figures 4 List of Tables 14 1 Introduction 15 2 Model Description The Climate Model Wind Forcing Experimental Design and Diagnosis 24 4 Control Simulations Meridional Overturning Horizontal Circulation Temperature-Salinity Properties Oceanic Response to Wind Forcing Horizontal Transport Streamfunction Global Meridional Overturning

4 3 5.3 Atlantic Meridional Overturning Surface Temperature and Salinity Resposne Global Interior Temperature and Salinity Response Climate Response Precipitation Northward Heat Transport Sea-Ice Variability Water Mass Response Control experiment Poleward shift anomaly Equatorward shift anomaly Carbon Uptake Change in Carbon Dioxide CO 2 in a Paleoclimate Context Summary and Conclusions 83 Appendix 87 Appendix A: Full set of Figures Appendix B: Table of Abbreviations Acknowledgements 146 Bibliography 147

5 List of Figures 2.1 The interacting climate processes and subcomponents of the University of Victoria Earth System Climate Model (adapted from Weaver (22)) Zonally averaged zonal wind stress in Pascals (Pa) for various wind climatologies, plotted against latitude (courtesy of Peter Oke, 23) Zonally averaged zonal wind stress for the control case (black); poleward shift (blue), and equatorward shift (red) The four defined regions for the passive tracers for water masses: south of 6 S traces Antarctic Bottom Water; 5 S to 6 S traces Antarctic Intermediate Water, and 4 S to 5 S traces Subantarcitic Mode Water. Deep water from the North Atlantic Ocean can be traced in the region north of 4 S, along with other water masses originating outside the Southern Ocean Global meridional overturning for the control case, in 2 Sverdrup (Sv) intervals Atlantic Ocean meridional overturning for the control case, in 1 Sv intervals Horizontal transport streamfunction for the control case, in 1 Sv intervals a) Sea Surafce Temperature (SST) for the control case and b) the difference between the control and Levitus experiments; in degrees Celsius ( C)

6 5 4.5 a) Sea Surface Salinty (SSS) for the control experiment and b) the difference between the control and Levitus experiments; units in practical salintiy units (psu) Horizontal transport streamfunction (Sv) difference between (a) the poleward shift and control case, and (b) the equatorward shift and control case; intervals are 5 Sv Temperature cross-section at 54.9 S in C, with the Drake Passage marked by the bold dashed line. Note that the depth axis is non-uniform and shows the depth levels in the Modular Ocean Model Global meridional overturning for a) the poleward-control and b) equatorwardcontrol differences; intervals are.5 Sv and the zero line is in bold Atlantic Ocean meridional overturning for a) the poleward-control, in.25 Sv intervals, and b) equatorward-control, in.1 Sv intervals Control simulation for (a) Sea Surface Temperature ( C) and (b) Sea Surface Salinity (psu). Only every second vector is shown, with vectors less than 5% of the maximum velocity omitted. The scale is indicated in the bottom left corner Poleward wind shift minus the Control case for (a) Sea Surface Temperature ( C) and (b) Sea Surface Salinity (psu). Only every second vector is shown, with vectors less than 5% of the maximum velocity omitted. The scale is indicated in the bottom left corner

7 6 5.7 Equatorward wind shift minus the Control case for (a) Sea Surface Temperature ( C) and (b) Sea Surface Salinity (psu). Only every second vector is shown, with vectors less than 5% of the maximum velocity omitted. The scale is indicated in the bottom left corner Control plots for a) zonally averaged temperature; intervals in 1 C with the zero line in bold, and b) zonally averaged salinity; intervals in.1psu Zonally averaged temperature for a) poleward-control, and b) equatorwardcontrol. Intervals in.1 C, with the zero line in bold Zonally averaged salinity for a) poleward-control, and b) equatorwardcontrol. Intervals in.1psu, with the zero line in bold Precipitation for a) the control case and b) the National Centers (NCEP) Reanalysis observations, in kgm 2 s 1. The equivalent colourbar range in mm/year is to 442 mm/year Precipitation for a) the poleward shift-control case and b) the equatorward shift-control case, in kgm 2 s Northward heat transport in Peta Watts (P W ) for each of the experiments: control (black); poleward shift (blue), and equatorward shift (red). Also shown are the differences: poleward - control (blue dashed line) and equatorward - control (red dashed line) Sea-ice thickness for a) the control experiment, b) poleward-control and c) equatorward-control; units in metres

8 7 7.1 Global average control plot of (a) Antarctic Bottom Water (AABW), (b) Antarctic Intermediate Water (AAIW), and (c) Subantarctic Mode Water (SAMW) tracer concentrations, respectively. (d) Atlantic Ocean average for North Atlantic Deep Water (NADW) for the control experiment Horizontal cross-section of Antarctic Bottom Water (AABW) tracer below the depth of 3m Horizontal cross-sections for the control experiment for the Antarctic Intermediate Water (AAIW) tracer Horizontal cross-sections for the control experiment for the Subantarctic Mode Water (SAMW) tracer Horizontal cross-sections for the control experiment for the North Atlantic Deep Water (NADW) tracer Changes due to the poleward shift in the Subpolar Westerlies in Antarctic Intermediate Water (AAIW) and Subantarctic Mode Water (SAMW) Changes in mode and intermediate waters averaged over the Pacific Ocean Conceptual diagram of water masses during a) the control experiment and for each of b) the poleward and c) the equatorward wind shifts in latitude Changes due to the equatorward wind shift in Antarctic Intermediate Water (AAIW) and Subantarctic Mode Water (SAMW) Surface dissolved inorganic carbon (DIC; in µmol/kg) for a) the Control experiment and b) Poleward shift - Equatorward shift; units in µmol/kg Surface dissolved inorganic carbon (DIC; in µmol/kg) for a) Poleward - Control and b) Equatorward - Control; units in µmol/kg

9 8 8.3 Globally averaged dissolved inorganic carbon (DIC) in µmol/kg for a) the control case; and differences between b) the poleward shift and the control case, c) the equatorward shift and the control case, and d) the poleward and equatorward shifts Atlantic Ocean average dissolved inorganic carbon (DIC) in µmol/kg for a) the control case; and differences between b) the poleward shift and the control case, c) the equatorward shift and the control case, and d) the poleward and equatorward shifts Pacific Ocean average dissolved inorganic carbon (DIC) in µmol/kg for a) the control case; and differences between b) the poleward shift and the control case, c) the equatorward shift and the control case, and d) the poleward and equatorward shifts Indian Ocean average dissolved inorganic carbon (DIC) in µmol/kg for a) the control case; and differences between b) the poleward shift and the control case, c) the equatorward shift and the control case, and d) the poleward and equatorward shifts Carbon Dioxide concentration in the atmosphere over the past 4, years, in parts per million volume (ppmv; adapted from PETIT et al. (1999)) a) Zonal averaged temperature; intervals in.2 C with zero line in bold, and b) global overturning; intervals in 2Sv, for the poleward-equatorward shift difference A1 Sea surface temperature (SST) for a) the Control experiment and b) Poleward shift - Equatorward shift; units in C

10 9 A2 Sea surface temperature (SST) for a) Poleward shift - Control and b) Equatorward shift - Control ; units in C A3 Sea surface salinity (SSS) for a) the Control experiment and b) Poleward shift - Equatorward shift; units in psu A4 Sea surface salinity (SSS) for a) Poleward shift - Control and b) Equatorward shift - Control ; units in psu A5 Zonally averaged temperature for a) the Control experiment and b) Poleward shift - Equatorward shift; units in C A6 Zonally averaged temperature for a) Poleward shift - Control and b) Equatorward shift - Control ; units in C A7 Zonally averaged salinity for a) the Control experiment and b) Poleward shift - Equatorward shift; units in psu A8 Zonally averaged salinity for a) Poleward shift - Control and b) Equatorward shift - Control ; units in psu A9 Global meridional overturning (MOT) for a) the Control experiment and b) Poleward shift - Equatorward shift; units in Sv A1 Global meridional overturning (MOT) for a) Poleward shift - Control and b) Equatorward shift - Control ; units in Sv A11 Atlantic Ocean overturning a) the Control experiment and b) Poleward shift - Equatorward shift; units in Sv A12 Atlantic Ocean overturning for a) Poleward shift - Control and b) Equatorward shift - Control ; units in Sv

11 1 A13 Horizontal transport streamfunction for a) the Control experiment and b) Poleward shift - Equatorward shift; units in Sv A14 Horizontal transport streamfunction for a) Poleward shift - Control and b) Equatorward shift - Control ; units in Sv A15 Precipitation for a) the Control experiment and b) Poleward shift - Equatorward shift; units in kgm 2 s A16 Precipitation for a) Poleward shift - Control and b) Equatorward shift - Control ; units in kgm 2 s A17 Defined tracer regions for a) south of 6 S and b) 5 S to 6 S A18 Defined tracer regions for a) 4 S to 5 S and b) north of 4 S A19 Control case for Antarctic Bottom Water (AAIW): a) below 3m, and b) for the global average A2 Control case for Antarctic Intermediate Water (AAIW) at a) 6m and b) 793m A21 Control case for Antarctic Intermediate Water (AAIW) at a) 112m and b) 1257m A22 Control case for Antarctic Intermediate Water (AAIW) for the a) Pacific Ocean and b) Atlantic Ocean averages A23 Control case for Subantarctic Water (SAMW) at a) 433m and b) 6m A24 Control case for Subantarctic Water (SAMW) at a) 793m and b) 112m A25 Control case for Subantarctic Water (SAMW) for the a) Pacific Ocean and b) Indian Ocean averages

12 11 A26 Control case for North Atlantic Deep Water (NADW) at a) 1528m and b) 1825m A27 Control case for North Atlantic Deep Water (NADW) for the a) global and b) Atlantic Ocean averages A28 Poleward shift - Control for Antarctic Bottom Water (AABW) for the a) Pacific Ocean and b) Atlantic Ocean averages A29 Poleward shift - Control for Antarctic Intermediate Water (AAIW) at a) 292m and b) 433m A3 Poleward shift - Control for Antarctic Intermediate Water (AAIW) for the a) global and b) Pacific Ocean averages A31 Poleward shift - Control for Subantarctic Mode Water (SAMW) for the a) 292m and b) 433m A32 Poleward shift - Control for Subantarctic Mode Water (SAMW) for the a) Indian Ocean and b) Pacific Ocean averages A33 Poleward shift - Control for North Atlantic Deep Water (NADW) for the a) Indian Ocean and b) Atlantic Ocean averages A34 Equatorward shift - Control for Antarctic Bottomw Water (AABW) for the a) Pacific Ocean and b) Atlantic Ocean averages A35 Equatorward shift - Control for Antarctic Intermediate Water (AAIW) at a) 292m and b) 433m A36 Equatorward shift - Control for Antarctic Intermediate Water (AAIW) for the a) Pacific Ocean and b) Atlantic Ocean averages

13 12 A37 Equatorward shift - Control for Subantarctic Mode Water (SAMW) at a) 292m and b) 433m A38 Equatorward shift - Control for Subantarctic Mode Water (SAMW) for the a) Pacific Ocean and b) global averages A39 Equatorward shift - Control for North ATlantic Deep Water (NADW) for the a) Indian Ocean and b) Atlantic Ocean averages A4 Poleward shift - Equatorward shift for Antarctic Bottom Water (AABW) for the a) global and b) Pacific Ocean averages A41 Poleward shift - Equatorward shift for Antarctic Intermediate Water (AAIW) at a) 177m and b) 292m A42 Poleward shift - Equatorward shift for Antarctic Intermediate Water (AAIW) for the a) global and b) Pacific Ocean averages A43 Poleward shift - Equatorward shift for Subantarctic Mode Water (SAMW) at a) 177m and b) 292m A44 Poleward shift - Equatorward shift for Subantarctic Mode Water (SAMW) for the a) Indian Ocean and b) Atlantic Ocean averages A45 Poleward shift - Equatorward shift for North Atlantic Deep Water (NADW) for the a) Indian Ocean and b) Atlantic Ocean averages A46 Surface dissolved inorganic carbon (DIC) for a) the Control experiment and b) Poleward shift - Equatorward shift; units in µmol/kg A47 Surface dissolved inorganic carbon (DIC) for a) Poleward - Control and b) Equatorward - Control; units in µmol/kg

14 13 A48 Globally averaged dissolved inorganic carbon (DIC) for a) the Control experiment and b) Poleward shift - Equatorward shift; units in µmol/kg A49 Globally averaged dissolved inorganic carbon (DIC) for a) Poleward - Control and b) Equatorward - Control; units in µmol/kg A5 Atlantic Ocean average dissolved inorganic carbon (DIC) for a) the Control experiment and b) Poleward shift - Equatorward shift; units in µmol/kg A51 Atlantic Ocean average dissolved inorganic carbon (DIC) for a) Poleward - Control and b) Equatorward - Control; units in µmol/kg A52 Pacific ocean average dissolved inorganic carbon (DIC) for a) the Control experiment and b) Poleward shift - Equatorward shift; units in µmol/kg A53 Pacific Ocean average inorganic carbon (DIC) for a) Poleward - Control and b) Equatorward - Control; units in µmol/kg A54 Indian Ocean average dissolved inorganic carbon (DIC) for a) the Control experiment and b) Poleward shift - Equatorward shift; units in µmol/kg A55 Indian Ocean average dissolved inorganic carbon (DIC) for a) Poleward - Control and b) Equatorward - Control; units in µmol/kg A56 Sea-ice thickness for a) the Control experiment and b) Poleward shift - Equatorward shift; units in metres A57 Sea-ice thickness for a) Poleward - Control and b) Equatorward - Control; units in metres

15 List of Tables 3.1 Experimental design with control, poleward and equatorward shifts in the Subpolar Westerly Wind (SWW) maximum displayed in degrees south Transport values for overturning cells Transport values (in Sv) for selected currents in the Southern Hemisphere. 31

16 Chapter 1 Introduction The present climate system has been witnessing a shift in the magnitude and location of the Subpolar Westerly Winds (SWWs), on both seasonal (Ruijter and Boudra, 1985) and longer time scales [Thresher (22); Shulmeister et al. (24), and Fyfe (23)]. Yet little is known about the climate s response to these shifts. The SWWs are mainly driven by meridional temperature and pressure gradients across the Southern Hemisphere (Shulmeister et al., 24). They act as a primary forcing mechanism over the Southern Ocean, affecting horizontal currents and water masses, as well as controlling the latitude of extratropical lows and rainfall bands in the atmosphere (Karoly, 23). They are also believed to influence North Atlantic Deep Water (NADW) outflow [Toggweiler and Samuels (1995); Rahmstorf and England (1997); and Furue and Suginohara (22)], but the extent of the significance of this role is unclear. In general, previous assessments of the variations in the SWWs have focused on changes in the strength of the wind maximum, with only one study to date (Oke and England, 24) analysing the impacts of changes in the latitude of the westerlies. Oke and England (24) performed a transient run of a 5.4 poleward shift in the latitude of the SWWs in an ocean-only general circulation model (GCM).

17 16 They found a decrease in Antarctic Intermediate Water (AAIW) formation, changes in the depth, size and position of the Deacon Cell, and changes in temperature and salinity throughout the ocean. However, their study included no atmospheric feedback, thermodynamic/dynamic sea-ice model, or thermochemical land-ice model. Here we will examine the climate s response to a poleward and equatorward shift in the latitude of the Subpolar Westerlies in a coupled climate model of intermediate complexity. From the mid-197s, the ocean s circulation around Antarctica underwent a noticeable change, coinciding with a shift in Pacific Ocean circulation and an increase in temperatures in the troposphere above the South Pole (Neff, 24). Further, Gibson (1992) noted a southward shift of.18 latitude per year in the Southern Hemisphere s subtropical wind maximum, from the late 197s. This is thought to be directly related to the enhanced Greenhouse Effect and global warming, since the subtropical winds indicate changes in meridional temperature gradients (Gibson, 1992). Increasing levels of atmospheric greenhouse gases coincide with stronger SWWs, and higher (lower) pressure levels at mid (high)- latitudes (Karoly, 23). The earth has been warming over the past 5 years with temperatures at mid-latitudes near the Antarctic Circumpolar Current (ACC) rising at a rate of.17 C/year, which is faster than the mean global increase rate (Gille, 22). The mean surface air temperatures have also been rising at a larger rate at high latitudes (Vaughan et al., 21), increasing vegetation cover, reducing glacier size, and lowering snow cover. These climate changes affect salinity properties due to changes in rainfall and evaporation (Bindoff and McDougall, 1994). The 15 to 2% decrease in precipitation over southwest Western Australia (SWWA) has been connected to a poleward shift in rain-bearing weather systems (Karoly, 23), which may be linked to the observed

18 17 shifting of the SWWs maximum towards higher latitudes. Changes in the SWW stress are believed to influence transport across the ACC due to changes in the density field (Borowski et al., 22). Gnanadesikan and Hallberg (2) noted that increases in the strength of the SWWs would drive a greater Ekman flux at mid-latitudes, altering overturning at intermediate depths in the Northern Hemisphere, and consequently increasing ACC transport. The SWWs can change the ocean s gyre structure (Bindoff and McDougall (1994); Oke and England (24)) and large-scale circulation near Antarctica, which may advect warmer air into the region (Vaughan et al., 21). Further, changes in the SWWs can result in a change in the properties of mixedlayer waters being transported into the thermocline via subductive processes (Bindoff and McDougall, 1994), which can alter intermediate and mode water-mass properties (Saenko and England, 23). The SWWs have been linked to future global warming and to changes in paleoclimates. Aeolian sediment records during the Last Glacial Maximum (LGM) showed enhanced upwelling near the east coast of New Zealand due to increased strength in westerly wind circulation (Shulmeister et al., 24). Also, the position of the SWWs is noted to have been located further northward (41 S), compared to its current position of 55 S, during the LGM (Toggweiler et al., 24). Bostok et al. (24) suggested that a shift in the SWWs caused a similar shift the Tasman Front, resulting in temperature and circulation changes. The Tasman outflow is affected by variations in wind stress, with a southward shift producing a stronger outflow, and vice versa for a shift towards the equator (Rintoul and Sokolov, 21). For a broader discussion of literature related to the SWWs, please refer to the Literature Review included at the end of this thesis.

19 18 The future and paleo climate changes above have motivated us to assess the long-term response of the ocean and climate system to shifts in the latitude of the westerlies. Until now, this type of experiment has been solely attempted by Oke and England (24) in an ocean GCM, with only a transient response analysed, and no atmospheric feedback or comprehensive sea-ice model. We propose to perform two experiments, one shifting the SWW stress equatorward, and the other shifting it poleward. In both experiments, the SWWs will be shifted 5.4 latitude, to coincide with the shift chosen by Oke and England (24), and the maximum wind speed will not be altered. The equatorward shift will allow us to assess the role of past changes; where as present and future climate impacts will be determined from results of the poleward wind shift. The difference between the two shifts will allow us to discuss the affects in a paleoclimate context, because the SWW maximum has changed by approximately 1 since the Last Glacial Maximum The University of Victoria (UVic) Earth System Climate Model (ESCM) has been chosen to run this climate simulation. The model couples a full ocean general circulation model with a single-layer atmosphere model and a sea-ice model. Diagnostic analyses of water mass and carbon dioxide changes will be conducted so that variations in the ocean s response can be further examined. The rest of the thesis is divided as follows. A brief description of the model, and the experimental design used, are outlined in Sections 2 and 3, respectively. The control simulation is compared to recent studies in Section 4. Sections 5 to 8 discuss the results of the experiments, with respect to overturning; rainfall and heat fluxes; water masses, and carbon uptake in the ocean. A summary and conclusions of the thesis are presented in Section 9. A list of abbreviations used in this thesis and a complete set of modelled

20 diagnostic figures are included in the Appendix. 19

21 Chapter 2 Model Description 2.1 The Climate Model The model used is the the University of Victoria Earth System Climate Model (hereafter known as the UVic ESCM) version 2.6, with a spherical grid resolution of 1.8 (meridional) by 3.6 (zonal). It is an intermediate complexity model, with a dynamic atmospheric energy-moisture-balance model (EMBM), a thermodynamic/dynamic sea-ice model, and a 3-dimensional ocean general circulation model coupled. A detailed description of the UVic ESCM is given in Weaver et al. (21), so the model s components are only briefly summarised here. The single layer atmospheric EMBM uses some concepts from Fanning and Weaver (1996). It transports heat through diffusion and moisture through advection. The model detects significant changes in incoming solar, and outgoing longwave, radiation, due to changes in the concentration of carbon dioxide. To assess the carbon uptake we included a calculation of the dissolved inorganic carbon in the ocean model. In this analysis, the atmospheric carbon dioxide concentration was set to the pre-industrial level of 28 ppm.

22 21 The sea-ice model combines multi-level thermodynamics and a rotated coordinate system, which allows a comprehensive evaluation of processes occurring at high latitudes. The use of elastic-viscous-plastic rheology (Hunke and Dukowicz, 1997) improves model computational efficiency and relates the internal ice stresses to ice deformation and thickness. The 19-level Geophysical Fluid Dynamics Laboratory (GFDL) Modular Ocean Model version 2 (MOM2; Pacanowski (1995)) is based on Navier-Stokes equations and the Boussinesq and Hydrostatic approximations, with depth levels ranging between 25 and 5m. The isopycnal tracer diffusivity was set at cm 2 /s. Brine rejection is parametrised during sea-ice formation to improve the representation of deep and bottom water properties, however, the NADW forms further south than observed(on the equatorward side of the Greenland/Scotland Ridge), which results in greater sea-ice cover forming in the high latitudes of this region. Coupling occurs every two ocean and four atmospheric time steps. The submodels (Figure 2.1) are coupled through heat fluxes and the exchange of water at the air/sea and sea/ice interfaces. Wind stress is applied over both the ocean and ice, with sensible heat and evaporative fluxes being controlled by wind speed. The Gent and McWilliams (199; hereafter known as GM) mixing scheme has been included in the UVic ESCM to allow for a parametrisation of the effects of eddy-induced mixing and reduce errors caused by coarse resolution (Santoso and England, 24). Furthermore, climate drift is moderated, because GM mixing reduces deep convection in the ACC and elsewhere in the high latitudes of the Southern Ocean (Hirst et al., 2). It can also improve the vertical structure of model temperature and salinity by decreasing exaggerated convection at high latitudes, and produce more realistic properties for water masses in the Southern Hemisphere (Sørensen et al., 21).

23 22 Figure 2.1: The interacting climate processes and subcomponents of the University of Victoria Earth System Climate Model (adapted from Weaver (22)). 2.2 Wind Forcing The wind climatology used to force the model is specified by the National Centers for Environmental Prediction and National Center for Atmospheric Research (NCEP-NCAR) Reanalysis, which is a global reanalysis of observed atmospheric fields. The data used by NCEP-NCAR date from 1958 to 1997 (including selected monthly fields for 1948 to 1957) with records post 1979 taken from modern satellite data (Kistler et al., 21). The recent Oke and England (24) study used the Hellerman and Rosenstein (1983) wind data, which is an in situ based analysis containing only ship-based data up to The Southern Hemisphere is not documented as thoroughly as the Northern Hemisphere, so

24 23 there are large discrepancies in the maximum wind stress region at 4-65 S between the Hellerman and Rosenstein (1983) and NCEP-NCAR reanalysis data, as shown in Figure 2.2. For comparison, also shown are the zonal mean zonal wind data supplied by the European Centre for Medium-Range Weather Forecasts (ECMWF) and the analysis of the 1mb winds of the ECMWF, during , by Trenberth et al. (199) [denoted as TLO]. The wind stress maximum of the NCEP-NCAR data is both stronger and postioned further south (by approximately 5 in latitude), compared to the HR data. Figure 2.2: Zonally averaged zonal wind stress in Pascals (Pa) for various wind climatologies, plotted against latitude: Hellerman and Rosenstein (1983)[HR]; Trenberth et al. (199) [TLO]; National Centers for Environmental Prediction [NCEP], and the European Centre for Medium-Range Weather Forecasts [ECMWF] (courtesy of Peter Oke, 23). See text for further details.

25 Chapter 3 Experimental Design and Diagnosis Recent observations have implied that the westerlies in the Southern mid-latitudes have shifted, both north and south between the Last Glacial Maximum and the present (Shulmeister et al., 24), with implications of winds shifting poleward in the future. Wyrwoll et al. (2) assessed westerly wind shifts in latitude at the LGM through storm track records and Bostok et al. (24) looked at shifts in the westerlies as being the cause of a similar shift in the Tasman Front. Recent westerly shifts have been observed by Seager et al. (21) on decadal timescales, and Gibson (1992) concluded that the subtropical wind maximum oscillated in a the study of wind stress patterns in a 16-year ( ) wind climatology record. This evidence is motivation to perform experiments where in the SWW maximum is shifted poleward and equatorward to allow us to analyse past, present and future climate responses. Table 3.1 shows the three experiments that were conducted for the present study. A control case was run, with two other experiments, that only differed in the latitudinal position of the SWW maximum. For the control case, the SWW maximum is at 52.2 S and the poleward (57.6 S) and equatorward (46.8 S) anomalies have the SWW maximum

26 25 Experiment CONTROL CASE POLEWARD SHIFT EQUATORWARD SHIFT Latitude of the SWW maximum 52.2 S 57.6 S 46.8 S Table 3.1: Experimental design with control, poleward and equatorward shifts in the Subpolar Westerly Wind (SWW) maximum displayed in degrees south. shifted 5.4 south and north, respectively. Figure 3.1 shows the zonal mean windstress for the Southern Hemisphere in each of the experiments, between 3 and 7 S, which encompasses the SWW range. Each experiment was run for 2 computational years to produce equilibrated solutions. Two diagnostic runs were also performed to identify some of the mechanisms of climate change throughout the main experiments. The first was designed to trace water masses, and the second to detect changes in the dissolved inorganic carbon (DIC) in the ocean and determine carbon uptake. Figure 3.2 illustrates the four regions defined for the passive water-mass tracer diagnosis. Antarctic Bottom Water (AABW) was traced in the region south of 6 S; Antarctic Intermediate Water (AAIW) between 6 and 5 S; Subantarctic Mode Water (SAMW) in the range 5 to 4 S, and North Atlantic Deep Water (NADW) was traced north of 4 S. The four tracers were initialised at zero from the equilibrated control restart, and run for 2 years, with the maximum SWW positions applied accordingly. Each fractional tracer concentration was reset to 1. at the tracer source regions of Figure 3.2, and set to zero outside of its defined region. Carbon uptake in the ocean is brought about by two pumps - the biological pump and the solubility pump. The biological pump includes the process of organisms near the ocean s surface obtaining CO 2 through photosynthesis, then sinking to the cooler deep ocean, where CO 2 is stored. The solubility pump, in contrast, is driven by physical and chemical processes, where CO 2 is absorbed by cold polar waters,

27 26 Figure 3.1: Zonally averaged zonal wind stress for the control case (black); poleward shift (blue), and equatorward shift (red) in Pascals (Pa). sinks, and returns to the surface elsewhere through convection and upwelling. In terms of our wind shift experiments, only the solubility pump was analysed to allow us to examine temperature changes and alterations in the ocean s circulation, in isolation of the biological response.

28 Figure 3.2: The four defined regions for the passive tracers for water masses: south of 6 S traces Antarctic Bottom Water; 5 S to 6 S traces Antarctic Intermediate Water, and 4 S to 5 S traces Subantarcitic Mode Water. Deep water from the North Atlantic Ocean can be traced in the region north of 4 S, along with other water masses originating outside the Southern Ocean. 27

29 Chapter 4 Control Simulations 4.1 Meridional Overturning Overturning Maximum Transport (Sv) Cell Control Poleward Equatorward case wind shift wind shift Deacon Cell AABW production Abyssal Cell NADW production NADW outflow Table 4.1: Transport values (in Sv) for overturning cells: the Deacon Cell; Antarctic Bottom Water (AABW) production cell; the Abyssal Cell; North Atlantic Deep Wtaer (NADW) production cell; NADW outflow. The transport values for the overturning of bottom and deep water cells is shown in Table 4.1, with the results of each experiment included. Figure 4.1 shows global meridional overturning (MOT), with the AABW production Cell, Deacon Cell and Abyssal Cell indicated. Observations suggest a transport value of 8.1Sv (1Sv = m 3 /s) for AABW production (Orsi et al., 1999), which is in agreement with our results. The maximum transport in the Deacon Cell for the control experiment is 31.2 Sv, which is

30 29 higher than Oke and England (24), since we do not include GM eddy-indcued advection in the calculation of meridional overturning. This result, however, is typical in similar models (Döös and Webb (1994); Rahmstorf and England (1997)). Observations of the Abyssal Cell are sparse, due to its deep location. Studies have modelled its maximum transport to be in the range of 13Sv to 15Sv (Döös and Webb (1994); )), which is slightly lower than the UVic model. Changes of less than 1Sv occur in the the bottom water Global Overturning Figure 4.1: Global meridional overturning for the control case, in 2 Sverdrup (Sv) intervals. and abyssal cells due to shifts in the latitude of the SWWs. However, the direct effect of wind stress on the Ekman layer causes fluctuatins of approximately 2.5Sv in the Deacon Cell due to a poleward and equatorward wind shift. Overturning in the Atlantic is addressed in Figure 4.2, with the NADW production (5 to 13m) and outflow (1 to 15m) cells indicated. The formation of deep water

31 3 Atlantic Overturning Figure 4.2: Atlantic Ocean meridional overturning for the control case, in 1 Sv intervals. in the North Atlantic in our model is 15.6 Sv, which agrees with the observed range of 14 to 2 Sv (Gordon, 1986). Outflow of deep water into the South Atlantic is 15 to 2 Sv (Rintoul (1991); Gordon (1986)), which is higher than the 1.1 Sv shown in Figure 4.2. Overall our model produces comparable results for overturning. 4.2 Horizontal Circulation Figure 4.3 shows the horizontal transport streamfunction and the transport values for some of the Southern Hemisphere currents are displyed in Table 4.2, along with the results of transport for the control simulation and the wind shifts. The range of transport values observed are 44 to 85 Sv for the Agulhas Current (Toole and Raymer (1985);Grundlingh (198); Toole and Warren (1993)); 3 to 56 Sv for the Brazil Current (Maamaatua-

32 31 Observed Maximum Transport (Sv) Current Transport Control Poleward Equatorward Range (Sv) case wind shift wind shift Brazil 3 to Agulhas 44 to EAC 17 to ACC 97 to Table 4.2: Transport values for selected currents in the Southern Hemisphere: The Brazil Current, the Agulhas Current; the East Australian Current (EAC), and the Antarctic Circumpolar Current (ACC) through the Drake Passage. Values for the observations, control case, and each of the wind shifts are shown. Transport Streamfunction (Sv) Latitude (Degrees) Figure 4.3: Horizontal transport streamfunction for the control case, in 1 Sv intervals.

33 32 iahutapu et al., 1998); 17 to 27 Sv for the East Australian Current (Mata et al., 2); 97 to 154 Sv for the Antarctic Australian Current (ACC) through the Drake Passage (Whitworth (1983); Whitworth and Peterson (1985); Orsi et al. (1995)). For each cae, the modelled values lie within the observed ranges. 4.3 Temperature-Salinity Properties The sea surface temperature (SST; Figure 4.4) and salinty (SSS; Figure 4.5) are compared to the World Ocean Atlas (1994) data sets produced by the National Oceanographic Data Center (NODC) Ocean Climate Laboratory. The SST differs significantly in the southeast Pacific, in the Drake Passage region, and in the west Pacific along coastal margins. However, the UVic model compares favourably with the obsevred data in the central Pacific, along the ACC and in the South Atlantic. Surface salinity is higher in the model, compared to the observaed in the Arctic Ocean, which partly increases sea-ice formation in this region. The Southern Hemisphere salinity levels are in good agreement with Levitus, which leads to improved representation of water masses (Weaver et al., 21).

34 33 Latitude (Degrees) Latitude (Degrees) a) Control b) Control Levitus Figure 4.4: a) Sea Surafce Temperature (SST) for the control case and b) the difference between the control and Levitus experiments; in degrees Celsius ( C).

35 34 Latitude (Degrees) Latitude (Degrees) a) Control b) Control Levitus Figure 4.5: a) Sea Surface Salinty (SSS) for the control experiment and b) the difference between the control and Levitus experiments; units in practical salintiy units (psu).

36 Chapter 5 Oceanic Response to Wind Forcing 5.1 Horizontal Transport Streamfunction Figure 5.1a shows that a poleward shift in the winds causes horizontal transport in the ACC to shift southwards, which is consistent with observations made by Oke and England (24). The transport increases by up to 2 Sv south of 6 S. The ACC can only shift up to about 3 of latitude, because it is constrained by the Drake Passage and the Campbell Plateau, south of New Zealand (Toggweiler et al., 24). Figure 5.2 shows the crosssectional view of temperature due to the poleward wind shift at latitude 54.9 S within the ACC. A poleward wind shift leads to 2 C colder water being transported downwards through Ekman pumping in the ACC, which is transported along the Drake Passage. However, transport in the South Pacific gyre weakens causing surface waters, north of the control SWW maximum (latitude 52.2 S), to warm, as dipicted in the western side of the Drake Passage in Figure 5.2a. The change in the horizontal streamfunction, as a result of an equatorward wind shift, is shown in Figure 5.1b, with transport decreasing by a maximum of 1 Sv at the surface

37 36 south of the ACC. When the SWWs are positioned at a more northerly latitudes, the ACC also shifts northward, decreasing transport and downwelling warmer waters into the ocean s interior, as shown in Figure 5.2b in the Drake passage region. The South Pacific gyre increases by a maximum of 2 Sv, which transports cooler waters in the gyre, including the region west of the Drake Passage. The Subpolar Westerly Winds change the Ekman flux, which drives temperature changes at the surface and in the ocean s interior.

38 37 a) Poleward shift Control: Transport Streamfunction (Sv) Latitude (Degrees) b) Equatorward shift Control: Transport Streamfunction (Sv) Latitude (Degrees) Figure 5.1: Horizontal transport streamfunction (Sv) difference between (a) the poleward shift and control case, and (b) the equatorward shift and control case; intervals are 5 Sv.

39 38 a) Poleward shift Control: Temperature b) Eqautorward shift Control: Temperature Figure 5.2: Temperature cross-section at 54.9 S in C, with the Drake Passage marked by the bold dashed line. Note that the depth axis is non-uniform and shows the depth levels in the Modular Ocean Model. 5.2 Global Meridional Overturning Variations in the global meridional overturning cells, due to poleward and equatorward shifts in the SWWs, are addressed in Figure 5.3. Bottom water production (Table 4.1) reduces (increases) by only.5sv (.3Sv) when the westerlies shift poleward (equatorward). The Abyssal cell overturning increases by up to.9 Sv due to a poleward shift, but decreases by over 2 Sv due to the equatorward shift. The mechanisms controlling the circulation intensity in the Ayssal cell remain unclear. The Deacon Cell, located within the interior of the Antarctic Circumpolar Current, is the cell in the Southern Hemisphere that is most strongly controlled by the mid-latitude westerlies. North of 65 S, the SWWs drive an Ekman flux, pushing surface waters northward, and this water sinks to depths of 1 to 3 km between 3 and 45 S, due to a strong flow at the circumpolar frontal regions (Speer et al., 2). In the ocean s interior, the water travels southward and finally upwells just south of the ACC, closing the circulation of the Deacon Cell (Döös and Webb, 1994). Since the SWWs are the primary mechanism transporting surface waters northward, a

40 Figure 5.3: Global meridional overturning for a) the poleward-control and b) equatorwardcontrol differences; intervals are.5 Sv and the zero line is in bold a) Poleward Shift Control (Sv) b) Equatorward Shift Control (Sv)

41 4 change in the wind stress latitude will strongly affect the size and position of the Deacon Cell. Figure 4.1a shows the control global meridional overturning, with values of up to 31.2 Sv across the surface in the ACC. Figure 5.4a shows that a poleward wind shift in the SWWs causes the Deacon Cell to extend a further 1m in depth and shift in unison with the westerlies, and increase in transport by 2.4Sv. Conversely, the equatorward wind shift (Figure 5.3b) causes the Deacon Cell to shift northwards, reduces its depth from 37m to 28m, and decreases maximum transport by 2.5Sv. 5.3 Atlantic Meridional Overturning There exists controversy about the extent to which the SWWs affect the outflow and overturning of deep water in the North Atlantic. Toggweiler and Samuels (1995) suggested that the SWWs play a dominant role in upwelling, outflowing NADW out of the circumpolar zone and into the surface Ekman layer. Stronger SWWs then theoretically bring about more deep water formation in the North Atlantic and more outflow of NADW into the South Atlantic. Against this, the model results of Rahmstorf and England (1997) suggest that thermohaline forcing is the primary mechanism responsible for deep water outflow into the Southern Ocean. According to Furue and Suginohara (22), changes in the strength of NADW are not directly related to northward Ekman transport due to the SWWs and NADW upwelling in the Southern Ocean does not flow into the surface Ekman layer (Furue and Suginohara, 22). The experiments conducted by Toggweiler and Samuels (1995) and Oke and England (24) were more sensitive to wind forcing, because temperature and salinity are set from restoring boundary conditions (Rahmstorf and England, 1997). With a model set up using more realistic heat and freshwater fluxes,

42 41 a) Poleward Shift Control: Atlantic Overturning (Sv) b) Equatorward Shift Control: Atlantic Overturning (Sv) Figure 5.4: Atlantic Ocean meridional overturning for a) the poleward-control, in.25 Sv intervals, and b) equatorward-control, in.1 Sv intervals.

43 42 the thermohaline circulation appears to be the dominant driver of NADW formation and outflow. However, in most previous studies, the wind climatologies used, such as Hellerman and Rosenstein (1983), are no longer regarded as the most up-to-date data records. In addition, our study experiments include an ice model with sea-ice feedbacks, a current wind climatology (NCEP/NCAR Reanalysis) and realistic atmosphere-ocean feedbacks. Hence, it produces climatic responses of a higher degree of accuracy than previous studies that have seeked to determine the impacts of changes in the SWWS in the Southern Ocean. Inspite of the large changes in the Subpolar Westerly wind stress around 35 (Figure 3.1), North Atlantic Deep Water outflow changes by less than 1 Sv for the poleward and equatorward wind shifts (Table 4.1). NADW production decreases by a maximum of.8 Sv for the polewardwind shift and control differences, with almost no change whenthe equatorward wind shift is applied (Figure 5.4). By incorporating similar restoring boundary conditions to Rahmstorf and England (1997), for temperature, we realistically include a thermal feedback that keeps North Atlantic Deep Water less sensitive to changes due to the SWWs seen by Oke and England (24) and Toggweiler and Samuels (1995).

44 Surface Temperature and Salinity Resposne The Subpolar Westerlies alter the ACC zonally and the Ekman flux Meridionally, as shown in the control simulations for sea surface temperature and salinity, with the surface velocity vectors over laid (Figure 5.5). The maximum velocity along the ACC is.2ms 1. Figure 5.6a addresses the changes in temperature and salinity at the surface due to poleward wind shift. The superimposed velocity vectors show a shift southwards in the ACC and a stronger velocity field of up to.4ms 1. In the Drake Passage region, the poleward wind shift strengthens the Falkland Current around the southern tip of South America, which cools the surface waters. The East Australian Current (EAC), which flows south along the East coast of the Austrlian continent extends further southwards, accounting for the temperature rise of approximately 1 C, and salinity decrease of greater than.2psu, between Australia and New Zealand. The decrease of 8Sv in overturning (Figure 5.3a) at the surface, between 4 S and 5 S, causes a clear decline in salinity in the South Inidan basin and the increase in temperature. Weaker evaporation (figure not shown) along the west coast of the South American continent decreases salinity levels. The thermohaline circulation slows, in the ocean s interior in the North Atlantic, which decreases SST due to a weakened North Atlantic drift (Manabe and Stouffer, 1988). Warmer water is transported north in the Gulf Stream, then mixes with the cooler water, thus reducing salinity. The affects of an equtorward wind shift on sea surface temperature and salinity is shown in Figure 5.7, with the ACC and EAC further northwards. The increase in the surface velocity, towards the equator, in the EAC tranports colder waters into the region west of New Zealand. The path of the ACC now extends above New Zealand (Figure

45 44 5.1b), which cools the east coast. Further, water transported through the Drake Passage originates from a more northward position, which accounts for the temperature rise south of South America. The strong increase in overturning by 11Sv at the surface, in the latitude range 3 S and 5 S reduces SST in the southern parts of each ocean basin. The enhanced SSS along South America s west coast is caused by a similar increase in evaporation (figure not shown). There are minimal changes in the North Atlantic overturning, hence SST and SSS properties vary little here. Shifts in the latitude of the Subpolar Westerly Winds drive changes in the Ekman velocity and the ACC, which, in turn, affect temperature and salinity at the ocean s surface. 5.5 Global Interior Temperature and Salinity Response Figure 5.8 shows temperature and salinity in the ocean s interior for the control simulation. The MOT results can be attributed to shifts in the latitude of the Subpolar Westerly winds. The MOT then forces changes in the interior temperature and salinity distributions (Figures 5.9 and 5.1). The decrease in temperature in the upper 15 m of the subtropics is due to 1.5Sv increase in overturning in the same location. The winds drive a stronger Ekman flux over sea-ice along the coast of Antarctica (Figure?- P-C: ice figure!!!- not yet positioned correctly!!), which increases brine rejection and hence, salinity, along the continental shelf. An equatorward shift in the SWWs leads to strengthening in overturning at midlatitudes, which in turn cools surface waters by a maximum of.8 C. Changes in overturning also raise salinity levels as shown clearly in Figure 5.3b. Changes in temperature can be primarily caused by variations in MOT, where as salinity in the ocean s interior is

46 45 a) Control: Sea Surface Temperature ( o C) m/s b) Control: Sea Surface Salinity (psu) m/s Figure 5.5: Control simulation for (a) Sea Surface Temperature ( C) and (b) Sea Surface Salinity (psu). Only every second vector is shown, with vectors less than 5% of the maximum velocity omitted. The scale is indicated in the bottom left corner.

47 46 a) Poleward shift Control: Sea Surface Temperature ( o C) m/s b) Poleward shift Control: Sea Surface Salinity (psu) m/s Figure 5.6: Poleward wind shift minus the Control case for (a) Sea Surface Temperature ( C) and (b) Sea Surface Salinity (psu). Only every second vector is shown, with vectors less than 5% of the maximum velocity omitted. The scale is indicated in the bottom left corner.

48 47 a) Equatorward shift Control: Sea Surface Temperature ( o C) m/s b) Equatorward shift Control: Sea Surface Salinity (psu) m/s Figure 5.7: Equatorward wind shift minus the Control case for (a) Sea Surface Temperature ( C) and (b) Sea Surface Salinity (psu). Only every second vector is shown, with vectors less than 5% of the maximum velocity omitted. The scale is indicated in the bottom left corner.

49 controlled by sea-ice. 48

50 Figure 5.8: Control plots for a) zonally averaged temperature; intervals in 1 C with the zero line in bold, and b) zonally averaged salinity; intervals in.1psu Control: Zonally Averaged Temperature (C) Control: Zonally Averaged Salinity (psu)

51 5 a) Poleward shift Control: Zonal Average Temperature (C) Latitude (Degrees) b) Equatorward shift Control: Zonal Average Temperature (C) Latitude (Degrees) Figure 5.9: Zonally averaged temperature for a) poleward-control, and b) equatorward-

52 51 a) Poleward shift Control: Zonal Average Salinity (psu) Latitude (Degrees) Latitude (Degrees) b) Equatorward shift Control: Zonal Average Salinity (psu) Figure 5.1: Zonally averaged salinity for a) poleward-control, and b) equatorwardcontrol. Intervals in.1psu, with the zero line in bold.

53 Chapter 6 Climate Response 6.1 Precipitation Before analysing the model precipitation response to shifted the Subpolar Westerly Winds, we will briefly assess the control simulation of global rainfall rates, as shown in Figure 6.1. Our model is agreeable with observations in most regios, except in the equatorial latitude band. The small differences are due to the sparse NCEP data available in the Southern Ocean. Figure 6.2 indicates changes in precipitation due to the poleward and equatorward shifts as compared with the control experiment. There are two ways in which rainfall can change due to latitudinal wind shifts. Either, the wind shifts vary sea surface temperature, which changes evaporation and hence precipitation, or the wind shifts alter moisture advection, which then changes rainfall. There is freshening of surface waters in the ACC in the South Pacific due to a poleward wind shift, which is likely due to the increase in SST in the same region (Figure 5.6a). A decrease in moisture advection appears to reduce precipitation between latitudes 35 and 45 S, with temperatures changes at a minimum.

54 53 8 Control: Precipitation ( kg/m 2 s ) x Latitude (Degrees) Interpolated NCEP Precipitation ( kg/m 2 s ) x Latitude (Degrees) Figure 6.1: Precipitation for a) the control case and b) the National Centers (NCEP) Reanalysis observations, in kgm 2 s 1. The equivalent colourbar range in mm/year is

55 54 8 a) Poleward shift Control x Latitude (Degrees) b) Equatorward shift Control 8 x Latitude (Degrees) Figure 6.2: Precipitation for a) the poleward shift-control case and b) the equatorward shift-control case, in kgm 2 s 1.

56 55 An equatorward wind shift causes a strong increase in precipitation by kgm 2 s 1 (25 mm/yr), between 3 S and 38 S in the South Indian Ocean, however SST in this region has minimal changes, indicating that moisture advection plays the dominant role. However, the decrease in rainfall in the latitude band 4 to 6 S corresponds with a reduction in temperature in the surface layer. Our experiments clearly affect precipitation levels in the Southern Hemisphere, due to synonymous changes in sea surface temperature and moisture advection.

57 Northward Heat Transport The northward heat transport (NHT) for each experiment, and the differences between each shift and the control case, are addressed in Figure 6.4. Although the SWW maximum is positioned at 5 S, the greatest amount of heat transport occurs near 4 S, where a poleward (equatorward) shift in the SWWs leads to a.9p W decrease (.25P W increase), which is synonymous with a 7% increase (19% decrease) in the northward transport of cold water. Oke and England (24) achieved.27p W decrease for their poleward trend, which is a much stronger response than our model. We have used an atmospheric thermal feedback, allowing temperature to adjust to the wind shift, rather than in Oke and England (24), where restoring boundary conditions are used, making temperature (and hence heat transport) more sensitive to the SWW shift. The equatorward shift produces a higher change than the poleward shift, which we can attribute to the greater change of 5Sv in the South Pacific gyre circulation (Figure 5.2) and the wind stress, at latitudes between 3 and 5 S. NADW outflow changes minimally, so we disregard this as a mechanism in changing heat transport. Figure 5.4 shows that the eqautorward shift increases MOT by a maximum of 11Sv near 4 S, where as the poleward shift decreases overturning by only 8Sv, which would also account for the stronger response in NHT for the northward shift in the SWWs.

58 57 NORTHWARD HEAT TRANSPORT Total heat transport (PW) Latitude (degrees north) Figure 6.3: Northward heat transport in Peta Watts (P W ) for each of the experiments: control (black); poleward shift (blue), and equatorward shift (red). Also shown are the differences: poleward - control (blue dashed line) and equatorward - control (red dashed line).

59 Sea-Ice Variability The control and difference plots for sea-ice are addressed in Figure.... To explain the low atmospheric CO 2 levels, due to reduced deepwater ventilation, Sigman and Boyle (2) hypothesised that a cooler climate would cause the westerlies in the Southern Hemisphere to shift northward, decreasing deepwater upwelling into the surface layer near Antarctic and replacing it by upwelling of water from intermediate depths into the subantarctic surface. This would cause a fresh, stable ice-covered surface to form along the coast of Antarctica, which further reduces deepwater ventilation and, hence, CO 2 outgassing (Stephens and Keeling, 2). On the other hand, interglacial periods would lead to higher levels of atmospheric CO 2. The global overturning change due to an equatorward shift (Figure 5.4b) supports the theory made by Sigman and Boyle (2), with overturning decreasing by 8 Sv at the surface in the Antarctic region, but increasing in the subantarctic waters. Sea-ice thickness is addressed in Figure 8.6. An increase in ice-thickness during an equatorward shift at the LGM corresponds with lower amounts of DIC at the 6 S latitude, and the opposite holds for a poleward shift. The change in sea-ice thickness is a response to changes in CO 2 (Toggweiler et al., 24), with a positive feedback resulting.

60 59 a) Control 4 Latitude (Degrees) b) Poleward shift Control.4 Latitude (Degrees) c) Equatorward shift Control.4.4 Latitude (Degrees) Figure 6.4: Sea-ice thickness for a) the control experiment, b) poleward-control and c) equatorward-control; units in metres.

61 Chapter 7 Water Mass Response 7.1 Control experiment The results for water masses for the control case in horizontal and vertical plots are shown in Figures 7.1 to 7.5. We note that the regions defined as mode water (4 to 5 S) and intermediate water (5 to 6 S) are predominantly for the Pacific and Indian Oceans, since formation of AAIW in the Atlantic lies between 4 and 5 S. The primary focus here is on changes in AAIW and SAMW, since the differences in NADW and AABW, due to latitudinal shifts in the SWWs, are only minimal (3-5% change), and are therefore not discussed in this section. The associated figures showing changes for these two water masses are located in AppendixA. Bottom Water forms in the Weddell and Ross Seas along the coast of Antarctica (Figure 7.2) through processes of brine rejection. Sea-ice forms along the coast of the Antarctic continent, leaving the water beneath high in density. The cold saline surface waters sink to the bottom of the ocean and spread northward (Figure 7.1a). The formation of deep water is addressed in Figure 7.5, showing the Atlantic average for the tracer entering the surface north of 4 S. NADW reaches depths

62 61 of 3m and originates in the Norwegian and Greenland seas in the North Atlantic Ocean. Three mechanisms are responsible AAIW formation, namely convection, subduction, and subsurface mixing (Sørensen et al., 21). AAIW forms in the Southern Hemisphere between 5 and 6 S, through the sinking of dense cold water and subsurface mixing. At 8-1m depth, around 4 S, the water mass is transported northward to 3 N in the North Atlantic (Santoso and England, 24). Intermediate water in the South Pacific separates into a north and south component. In the northern branch, AAIW moves northward by an eastern boundary current as part of the South Pacific subtropical gyre. The southern branch sends AAIW eastward via the ACC, through the Drake Passage, then turning northward as it flows into the South Atlantic. Then part of the intermediate water joins the South Atlantic subtropical gyre (Santoso and England, 24). Subduction at mid-latitudes acts as the primary mechanism for the formation of AAIW (Sørensen et al., 21), and England et al. (1993) conclude that intermediate water is renewed by warm large-scale advection near Chile, which leads to convective overturning and the formation of deep mixed layers in AAIW. Figure 7.3 shows the formation of AAIW in the model to be in the south-east Pacific, then moving through the Drake Passage, into the south Atlantic, which agrees with observations. SAMW is formed north of the Antarctic Circumpolar Front, in the Subantarctic Front (see Figure 2 in Literature Review attached) between 45 and 5 S; in the same region where thick mixed layers and a band of heat loss to the atmosphere exist. The formation of SAMW in the control experiment (Figure 7.4) is evident in the south-east Indian Ocean, then extending northwards, which coincides with observations [Banks et al. (22);

63 62 Rintoul and England (22)]. Dense, warm waters originating in the Tasman Sea mix with cooler waters from the Southern Ocean to form the mixed layer of mode water, with the resulting convective overturning leaving temperatures constant throughout the upper layers of the water column. The model s main locations of SAMW formation are in the Southeast Pacific and Southeat Indian Oceans, in good agreement with observations. Global Average: AABW Global Average: SAMW Global Average: AAIW Atlantic Ocean Average: NADW Figure 7.1: Global average control plot of (a) Antarctic Bottom Water (AABW), (b) Antarctic Intermediate Water (AAIW), and (c) Subantarctic Mode Water (SAMW) tracer concentrations, respectively. (d) Atlantic Ocean average for North Atlantic Deep Water (NADW) for the control experiment.

64 Figure 7.2: Horizontal cross-section of Antarctic Bottom Water (AABW) tracer below the depth of 3m. Figure 7.3: Horizontal cross-sections for the control experiment for the Antarctic Intermediate Water (AAIW) tracer at depths (a) 433, (b) 6, (c) 793, and (d) 112 metres.

65 64 Figure 7.4: Horizontal cross-sections for the control experiment for the Subantarctic Mode Water (SAMW) tracer at depths (a) 292, (b) 433, (c) 6, and (d) 793 metres. Figure 7.5: Horizontal cross-sections for the control experiment for the North Atlantic Deep Water (NADW) tracer at depths (a) 1528, (b) 1825, (c) 2148, and (d) 2497 metres.

66 Poleward shift anomaly Changes due to the poleward shift for AAIW are shown in Figures 7.6a to 7.6f. The presence of AAIW increases in the upper 12m, due to colder waters being driven to greater depths by the poleward shift. The increase in formation can be seen further in Figure 7.7a, in the upper 1m of the south Pacific. Formation of SAMW, on the other hand, decreases (Figure 7.6d to 7.6f) in the south Pacific and south-east Indian Oceans. Figure 7.7b shows that the decrease in mode water occurs in the upper 1m, implying that the SWWs play a significantnt role in changing the formation and path of these water-masses. The conceptual diagram of the processes occurring for AAIW and SAMW during the poleward shift are shown in Figure 7.8b. As the SWW maximum shifts southward, so too does the Ekman transport. This now advects waters that are much colder. This increases convection at 5 to 6 S, thus enhancing the formation of AAIW. However, with the SWW maximum further southward, the thick deep mixed layer, signifying formation of SAMW, decreases significantly at the surface. Further, the wind stress, and hence Ekman transport, is weaker north of 54 S (Figure 3.1). The upwelling region for Circumpolar Deep Water (CDW) also shifts polewards due to the SWW shift, affecting the position of the ACC. Figure 5.6b shows the zonal average salinity due to the poleward shift. A poleward shift increases salinity south of the ACC, which enhances bottom water formation.

67 Figure 7.6: Changes due to the poleward shift in the Subpolar Westerlies in Antarctic Intermediate Water (AAIW; a-c) and Subantarctic Mode Water (SAMW; d-f) at depth levels of 177, 292 and 433 metres for intermediate water, and 292, 433 and 6 metres for mode water. 66

68 67 Poleward shift Control: AAIW Equatorward shift Control: AAIW Poleward shift Control: SAMW Equatorward shift Control: SAMW Figure 7.7: Changes in water-mass tracers averaged over the Pacific Ocean, for the poleward-control (plots a and b) and equatorward-control (plots c and d) for Antarctic Intermediate Water (AAIW) and Subantarctic Mode Water (SAMW)

69 Figure 7.8: Conceptual diagram of water masses during a) the control experiment and for each of b) the poleward and c) the equatorward wind shifts in latitude. Antarctic Intermediate Water (AAIW), Subantarctic Mode Water (SAMW), Antarctic Bottom Water (AABW), and Circumpolar Deep Water (CDW) are shown, with the Ekman transport in the surface layer illustrated by the blue arrow, and the Subpolar Westerly Wind (SWW) maximum (SWWs) shown by the circle, with a black dot in the centre, representing wind flowing out of the page (see text for further details). 68

70 Equatorward shift anomaly The changes in AAIW due to an equatorward shift are shown in Figures 7.9a to 7.9c, with the Pacific Ocean averages shown in Figures 7.7c. There is a substantial decrease of AAIW in the south-east Pacific in the upper 13m, since an equatorward shift transports warmer waters in the Ekman layer, reducing deep convection. SAMW formation increases (Figures 7.9d to 7.9f, and Figure 7.7d), as do temperatures in the upper 15m (Figure 5.5c). A northward shift in the SWW maximum drives warmer waters in the Ekman layer, which sink via subsurface-mixing and convection. Figure 7.8c is a conceptual representation of the water mass processes that change due to an equatorward shift in the latitude of the SWWs. Warmer waters are advected in the surface layer, which reduces convection and the formation of AAIW. Mode water is enhanced and CDW shifts northward, as shown by the cold tongue of water between 4 and 5 S in Figure 5.5c. Bottom water formation decreases with an equatorward shift due to a reduction in surface salinity (Figure 5.6c) in the Southern Ocean. Changes in intermediate and mode water have clearly occured due to shifts in latitude in the wind stress, which drives the Ekman flux in the surface layer, changing convective processes and hence the formation of these water masses. A poleward shift reduces SAMW, due to the southward migration of Ekman transport away from mode water formation zones. This increases AAIW formation. Conversely, an equatorward shift in the SWWs sees AAIW formation reduced, and an increased Ekman pumping and convection of SAMW.

71 7 Figure 7.9: Changes due to the equatorward shift in Antarctic Intermediate Water (AAIW; a-c) and Subantarctic Mode Water (SAMW; d-f) at depth levels of 292, 433 and 6 metres for intermediate water, and 433, 6 and 793 metres for mode water.

72 Chapter 8 Carbon Uptake 8.1 Change in Carbon Dioxide Only the physical processes of the solubility pump are examined and discussed here. Figures 8.1 and 8.2 shows the surface dissolved inorganic carbon (DIC) for the control case and the changes due to a poleward and equatorward wind shift. A poleward shift in the SWWs decreases DIC by up to 15 µmol/kg in the latitude band 3 S to 55 S. On the other hand, when the SWWs shift equatorward, the surface DIC increases in this region. Comparing these changes to those in rainfall (Figure 6.2) shows that the winds vary carbon uptake and atmospheric CO 2, which then changes precipitation levels in the Southern Hemisphere. An increase in atmospheric carbon levels increase mean sea level pressure (Cai et al., 23), which decreases rainfall, and a converse effect occurs for reduced carbon in the atmosphere. Our model simulations indicate that wind shifts in latitude can alter surface carbon concentration, which is partly responsible for differences in precipitation.

73 a) Control: Surface DIC (umol/kg) b) Poleward shift Equatorward shift: Surface DIC (umol/kg) Figure 8.1: Surface dissolved inorganic carbon (DIC; in µmol/kg) for a) the Control experiment and b) Poleward shift - Equatorward shift; units in µmol/kg.

74 a) Poleward shift Control: Surface DIC (umol/kg) b) Equatorward shift Control: Surface DIC (umol/kg) Figure 8.2: Surface dissolved inorganic carbon (DIC; in µmol/kg) for a) Poleward - Control and b) Equatorward - Control; units in µmol/kg.

75 74 Plots (a) of Figures 8.3 to 8.6 show the control dissolved inorganic carbon (DIC); averaged globally and for each ocean basin. Higher DIC levels exist at the poles due to colder waters having a higher solubility of carbon dioxide, from where the DIC is overturned into the deep ocean via convection. Eventually, CO 2 is outgassed in the warmer waters of the equatorial zone via abyssal upwelling of deep, cold carbon-rich waters. Plots (b) & (c) of Figures 8.3 to 8.6 show the poleward and equatorward wind shift changes for dissolved inorganic carbon in the ocean s interior, respectively. A poleward shift increases DIC in the ocean s interior, but CO 2 decreases at the surface in the subantarctic zone. This change in the surface layer occurs primarily in the Pacific, and to a lesser degree in the Indian Ocean. An equatorward shift results in low DIC in the deep ocean, with a strong increase in the Indian and Pacific Oceans in the region between 4 S and 55 S. These surface changes are directly related to the overturning in the mid-latitudes (Figure 5.4), with stronger MOT increasing carbon uptake in the ocean. Bottom water formation changes, brought about by differences in salinity (Figure 5.7) at the poles, corresponds with DIC levels in the deeper ocean. A poleward shift causes colder and denser surface waters, which increase AABW formation (see full set of figures in Appendix A) and carbon uptake, whereas an equatorward shift would decrease the DIC levels of the ocean s interior. We note that the bottom water changes are small (< 1%), as are the changes in CO 2, however, small changes in DIC in the ocean are significant. The Atlantic Ocean s DIC composition changes minimally due to the SWW shifts. We note here that changes in SST (Figures 5.3a & 5.3b) are large in the Atlantic Ocean, indicating that temperature and meridional overturning effects are compensating.

76 75 a) Control b) Poleward shift Control c) Equatorward shift Control d) Poleward shift Equatorward shift Figure 8.3: Globally averaged dissolved inorganic carbon (DIC) in µmol/kg for a) the control case; and differences between b) the poleward shift and the control case, c) the equatorward shift and the control case, and d) the poleward and equatorward shifts.

77 76 a) Control b) Poleward shift Control c) Equatorward shift Control d) Poleward shift Equatorward shift Figure 8.4: Atlantic Ocean average dissolved inorganic carbon (DIC) in µmol/kg for a) the control case; and differences between b) the poleward shift and the control case, c) the equatorward shift and the control case, and d) the poleward and equatorward shifts.

78 77 a) Control b) Poleward shift Control c) Equatorward shift Control d) Poleward shift Equatorward shift Figure 8.5: Pacific Ocean average dissolved inorganic carbon (DIC) in µmol/kg for a) the control case; and differences between b) the poleward shift and the control case, c) the equatorward shift and the control case, and d) the poleward and equatorward shifts. The primary mechanisms resposible for changes in carbon concentration are the terrestrial pump in the atmosphere, and the biological and solubility pumps in the ocean. As the earth s climate shifts to and from glacial and interglacial periods, the atmospheric carbon dioxide concentration can fluctuate between 1 ppmv (parts per million volume) and 125 ppmv, as shown in the Vostok Ice Core record in Figure 8.7. We calculated the global ocean DIC in our model to be 25,625 P gc (Peta grams of carbon), with a 25P gc, (12 ppmv in the atmosphere), increase for the poleward shift. The equatorward shift reduced global DIC by 38P gc (23 ppmv). A shift of 11 in latitude of the SWW maximum, which is the observed change between glacial and interglacial periods (Toggweiler et al., 24), would result in a change of 35 ppmv in atmospheric CO 2, which accounts

79 78 a) Control b) Poleward shift Control c) Equatorward shift Control d) Poleward shift Equatorward shift Figure 8.6: Indian Ocean average dissolved inorganic carbon (DIC) in µmol/kg for a) the control case; and differences between b) the poleward shift and the control case, c) the equatorward shift and the control case, and d) the poleward and equatorward shifts. Figure 8.7: Carbon Dioxide concentration in the atmosphere over the past 4, years, in parts per million volume (ppmv; adapted from PETIT et al. (1999)).