Dependence of tropical precipitation on changes in cross-equatorial atmosphere and ocean heat transport in global warming simulations.

Transcription

1 Dependence of tropical precipitation on changes in cross-equatorial atmosphere and ocean heat transport in global warming simulations Ashly Spevacek A thesis submitted in partial fulfillment of the requirements for the degree of Master of Science University of Washington 2017 Committee: Dargan Frierson, Chair Christopher Bretherton Daehyun Kim Program Authorized to Offer Degree: Department of Atmospheric Sciences

2 Copyright 2017 Ashly Spevacek

3 University of Washington Abstract Dependence of tropical precipitation on changes in cross-equatorial atmosphere and ocean heat transport in global warming simulations Ashly Ann Spevacek Chair of the Supervisory Committee: Dr. Dargan Frierson Department of Atmospheric Sciences The Intertropical Convergence Zone (ITCZ) is an area of convection encircling the earth near the equator where the northern and southern hemisphere (SH) Hadley cells converge. If a forcing (e.g. aerosol emissions decline) results in a greater net influx of heat into one hemisphere the Hadley cells will shift across the equator to transport heat from the warmer to the cooler hemisphere resulting in a shift of the annual mean location of the ITCZ. Shifts in the annual mean location of the ITCZ can significantly impact the nearly 3 billion who live in the tropics and depend on the ITCZ as a source of freshwater. Recent studies have shown the extent to which the Hadley cells will shift is dependent not only on the magnitude of the forcing but also the cross-equatorial ocean heat transport. In this study, we use Coupled Model Intercomparison (CMIP) output to determine the thermal forcings that result in hemispheric heating asymmetries

4 and induce an ITCZ shift in the 21 st century. Using a fully coupled model we also determine the atmosphere will transport most of the total cross-equatorial heat transport if the thermal forcing is applied in the northern hemisphere.

5 TABLE OF CONTENTS Chapter 1. Introduction Background ITCZ shifts in observations Last glacial maximum modeling studies ITCZ shifts in CMIP simulations Ocean-atmosphere c-eq heat transport split Causes for hemispheric energy imbalance Guiding questions Chapter 2. CMIP5 RCP8.5 moist static energy budget The global atmosphere energy budget Errors in the global climate models Chapter 3. Attribution of ITCZ shifts Quantifying ITCZ shifts in the 21 st century Attribution of the change in c-eq heat transport Effects of the AMOC slowdown Effects of the aerosol clean-up and ice-albedo feedback Effects of the longwave radiative feedback Result of competing forcings Extratropical forcing and feedbacks Ocean-atmosphere c-eq heat transport split Chapter 4. GFDL-MOM experiments Models and simulations i

6 4.2 C-eq atmosphere and ocean heat flux response Surface and TOA flux response Chapter 5. Summary and conclusion Bibliography ii

7 LIST OF FIGURES Figure 1.1. Atmosphere, ocean and net poleward heat transport. [Fassullo and Trenberth, 2008] Figure 2.1. CMIP5 RCP8.5 sublimation in W m -2. Sublimation values in kg m -2 s -1 were converted by multiplying by the latent heat of sublimation (2.83e6 J kg-1). Means shown are for 2006 through Figure 2.2. Difference between the latent heat flux and evaporative heat flux for CMIP5 RCP8.5. Mean upward latent heat flux at the surface minus the evaporation rate (kg m-2 s-1) x 2.3e6 J kg- 1 for 2006 through Figure 2.3. Upward latent heat flux at the surface minus the evaporation rate (kg m-2 s-1) x 2.3e6 J kg-1 and sublimation (kg m-2 s-1) x 2.83e6 J kg-1 for CMIP5 RCP8.5 models. Mean shown for 2006 through Figure 3.1. Change in c-eq atmospheric heat transport versus change in tropical precipitation asymmetry defined as the precipitation from 0 to 20N minus precipitation from 0 to 20S Figure 3.2. Change in precipitation asymmetry over each ocean basin versus the total change in c- eq AHT. Values over land excluded when calculating change in precipitation asymmetry over ocean basin. Change calculated as mean for 2081 through 2100 minus the mean for 2006 through Figure 3.3. Change in tropical precipitation (mm day -1 ). Change calculated by subtracting the mean for 2006 through 2025 from the mean for 2081 through Figure 3.4. Attribution of the change in c-eq AHT. For all terms a northward c-eq AHT (positive value) is induced by an increase in heat entering the southern hemisphere atmosphere. For the OHT term, a positive value indicates the northward OHT decreases, inducing a northward AHT. For the ocean storage term a positive value indicates the heat uptake in the northern hemisphere is greater than the heat uptake in the southern hemisphere, inducing a northward c-eq AHT. Change is calculated as the mean c-eq AHT for 2081 through 2100 minus the mean c-eq AHT for 2006 through Figure 3.5. Change in upward surface flux (LW + SW + sensible + latent heat fluxes). Change calculated as the mean for 2081 through 2100 minus the mean for 2006 through Figure 3.6. Change in c-eq AHT induced by Pacific/Indian and Atlantic OHT. For all terms a northward c-eq AHT (positive value) is induced by an increase in heat entering the southern iii

8 hemisphere atmosphere. Change is calculated as the mean c-eq AHT for 2081 through 2100 minus the mean c-eq AHT for 2006 through Figure 3.7. Change in incoming SW radiation due to clear-sky scattering changes calculated using the APRP method. Change calculated as the mean for 2081 through 2100 minus the mean for 2006 through Models with an asterisk (*) include the indirect effects of aerosols on cloud albedo and lifetime as well as direct aerosol radiative. Models with a plus sign (+) only include indirect effects of aerosols on cloud albedo in addition to direct radiative effects. Models without a symbol only include direct aerosol radiative effects Figure 3.8. Change in incoming SW radiation due to cloud changes calculated using the APRP method. Change calculated as the mean for 2081 through 2100 minus the mean for 2006 through Models with an asterisk (*) include the indirect effects of aerosols on cloud albedo and lifetime as well as direct aerosol radiative. Models with a plus sign (+) only include indirect effects of aerosols on cloud albedo in addition to direct radiative effects. Models without a symbol only include direct aerosol radiative effects Figure 3.9. Change in c-eq AHT induced by SW radiation changes. For all terms a northward c-eq AHT (positive value) is induced by an increase in heat entering the SH atmosphere. Change is calculated as the mean c-eq AHT for 2081 through 2100 minus the mean c-eq AHT for 2006 through Figure Change in c-eq AHT induced by LW radiation changes. For all terms a northward c- eq AHT (positive value) is induced by an increase in heat entering the SH atmosphere. Change is calculated as the mean c-eq AHT for 2081 through 2100 minus the mean c-eq AHT for 2006 through Figure C-eq AHT induced by aerosol Change in precipitation asymmetry versus change in c-eq AHT induced by clear-sky scattering, albedo, northward AMOC heat flux, ocean heat content changes and longwave changes Figure Change in precipitation asymmetry versus change in c-eq AHT induced by clear-sky scattering, albedo, upward ocean heat flux and northward AMOC heat flux changes Figure Change in c-eq AHT induced by radiative changes calculated using hemisphere totals versus change in c-eq AHT calculated with excluding 15 S to 15 N. A positive value indicates the radiative change results in a greater increase in heat entering the SH than NH inducing a northward AHT Figure Change in zonal mean column integrated heat flux (TOA down + surface flux up) into the atmosphere. Change calculated as the mean for 2081 through 2100 minus the mean for 2006 through iv

9 Figure Change in zonally integrated northward OHT. Change calculated as the mean for 2081 through 2100 minus the mean for 2006 through Figure Change in zonally integrated northward OHT in the Atlantic basin. Change calculated as the mean for 2081 through 2100 minus the mean for 2006 through Figure Northward AHT induced by changes in the upward ocean (surface) flux Figure Change in c-eq AHT versus change in c-eq OHT by basin. Values over land are omitted. Change calculated as the mean for 2081 through 2100 minus the mean for 2006 through Figure 4.1. Heat fluxes applied at the ocean surface in the GFDL-MOM5 simulations. For each simulation, except for SOcn10, the total PW added was constant at PW. For the SOcn10 run a total forcing of 0.6 PW was added. Three ensembles were run with each setup for 100 years..63 Figure 4.2. Change in northward atmosphere heat transport and ocean heat transport (PW, experiment control) Figure 4.3. Experiment minus control precipitation (mm/day) Figure 4.4. Experiment minus control precipitation asymmetry versus northward c-eq AHT and northward c-eq OHT Figure 4.5. Ensemble mean experiment minus control surface flux response for GFDL-MOM5 experiments Figure 4.6. Ensemble mean experiment minus control sea surface temperature response for GFDL- MOM5 experiments Figure 4.7. Experiment minus control zonal mean column integrated energy flux into the atmosphere, upward surface flux and net downward TOA (LW + SW) radiation (W m -2 ) Figure 4.8. Change in net downward TOA (LW + SW) radiation and low cloud fraction (experiment control). Values shown are mean values for the 100 period the experiments were run Figure 4.9. Experiment minus control northward c-eq OHT (PW) split up by ocean basin Figure Change in zonal wind stress (experiment control). Values shown are mean values for the 100 period the experiments were run v

10 LIST OF TABLES Table 2.1. Global atmosphere (ATM) energy imbalance and c-eq AHT values for beginning and end of 21st century RCP8.5 CMIP5 GCMs Table 2.2. Change in global latent heat flux due to sublimation and c-eq AHT due to sublimation for RCP8.5 models in which sublimation data is available Table 3.1. Change in c-eq AHT and c-eq OHT. A positive Δc-eq AHT indicates the increase in heat entering the southern hemisphere is greater than the increase in heat entering the northern hemisphere, inducing a northward c-eq AHT. A negative Δc-eq OHT induces a northward AHT Table 3.2. Change in c-eq AHT and the change in AHT induced by changes in the upward surface fluxes. A positive Δc-eq AHT due to upward surface fluxes value indicates the change in upward heat flux in the SH is greater than The upward surface heat flux includes ocean storage and northward OHT effects Table 3.3. C-eq OHT due to changes in northward heat transport and due to hemispheric storage asymmetries Table 3.4. Change in c-eq OHT by basin and change in c-eq AHT Table 4.1. GFDL-MOM5 simulations with the total heat added and region in which the heat was added Table 4.2. Change in c-eq AHT and OHT for GFDL-MOM5 experiments and select CESM coupled experiments vi

11 ACKNOWLEDGEMENTS There are many individuals I would like to thank for their support and guidance in completing this research. First and foremost, I would like to thank my advisor Dargan Frierson for taking me on as his student and providing guidance, encouragement and ideas throughout the course of my graduate studies. I am immensely grateful for your support and for making my graduate career easier by being an excellent advisor. I would also like to thank Daehyun Kim and Chris Bretherton for providing thought-provoking feedback and guidance throughout the course of my research. Thank you both for agreeing to being on my committee, reading and commenting on my thesis and for all your support with me research. I would also like to thank Yen-Ting Hwang, Elizabeth Maroon and Rachel White. Without the support of Yen-Ting, I would have struggled with my analysis of the Global Climate Model output. Thank you to Elizabeth for helping me learn how to run a global climate model and spending many hours helping me with my analysis. I am grateful for the time Rachel White spent with me discussing research ideas, ocean-atmosphere interactions and providing general career advice. I also appreciate the time I have had to discuss science with Kyle Armour, LuAnne Thompson, Matthew Wolfe, Rick Russotto, Stephanie Rushley, Tom Ackerman and many others in the department. I would like to thank Marc Michelson, Harry Edmon and David Warren for all their technical support. Thank you to the department staff, especially Erica Coleman and Melissa Pritchard, for all your support while I have been a student at the University of Washington. I would also like to vii

12 thank the Department of Atmospheric Sciences faculty for challenging me and supporting me. A big thank you to grads14 for helping me with coursework, research and for being a fun, inviting class. Finally, I would like to thank my family, friends and husband Brandon. Thank you for being supportive, loving, and an amazing husband. I could not have completed this without you. viii

13 Chapter 1. INTRODUCTION The Inter Tropical Convergence Zone (ITCZ) is an area of convection encircling the earth near the equator where the northern and southern hemisphere (NH & SH) Hadley cells converge. Global climate models (GCMs) often disagree on the location and intensity of the ITCZ (defined here as both the extension over land and ocean) as well as how the ITCZ is expected to change in the 21 st century. Some GCMs show an annual mean maximum precipitation just north of the equator indicating a stronger northern hemisphere (NH) ITCZ, while others show a maximum just south of the equator. In observations, the annual mean maximum precipitation is in the northern tropics. Looking to the future, some models show the ITCZ shifting north between 2006 and 2100 while others show it shifting south with warming. Shifts in the ITCZ result in substantial changes to the annual mean precipitation in the tropics which significantly impacts the lives of the almost 3 billion people who live in the tropics and depend on the ITCZ as a source of freshwater. If the differences in how GCMs model the ITCZ were better understood our understanding of how the ITCZ will shift with warming could be improved. The location of the ITCZ is dependent on the strength of the NH and SH Hadley cells. The Hadley cell is a large scale atmospheric circulation with poleward motion in the upper branch and equatorward motion in the lower branch. The lower, equatorward half of the Hadley cell transports moisture from the subtropics towards the tropics which results in drying out the subtropics and bringing moisture to the ITCZ. The upper poleward branch of the Hadley cell transports energy from the tropics to the mid-latitudes. If one hemisphere is warmer than the other the Hadley cell will shift across the equator to transport energy from the warmer hemisphere to the cooler hemisphere. Past studies have shown there are links between the atmospheric heat transport (AHT)

14 10 across the equator and the annual mean location of the ITCZ, indicating energy fluxes well outside of the tropics can impact the location of the ITCZ (Hwang and Frierson, 2013). Previous studies have also found differences in how GCMs model the hemispheric averaged heating result in varying ITCZ shifts. The relative impact of each forcing, along with the effects of the location and magnitude, are still unclear. In addition, the relative change in cross-equatorial (c-eq) atmosphere heat transport (AHT) and ocean heat transport (OHT) in response to heating anomalies is inconsistent among the models. 1.1 BACKGROUND ITCZ shifts in observations Since the ITCZ was first identified it has been well-known that the ITCZ intra-annual location is dependent on the seasonal heating cycle. During the NH summer the ITCZ is located north of the equator and during the SH summer the ITCZ shifts south. However, the relationship between the annual mean location of the ITCZ and inter-annual hemispheric heating asymmetries was only first discovered in observations in the 1980s. Folland et al. (1986) revealed a correlation between precipitation trends in the Sahel region of Africa and global sea-surface temperature (SST). More specifically, Folland pointed out there was a relationship between the Sahel wet and dry periods and hemispheric SST asymmetries on timescales of years to tens of years. Since then several studies have shown a correlation between ITCZ shifts in the latter half of the 20 th century hemispheric heating asymmetries in observations (Folland et al., 1986; Rotstayn and Lohmann, 2002; Chang et al., 2011; Hwang et al., 2013).

15 Last glacial maximum modeling studies ITCZ shifts appear not only in 20 th century observations but also in paleoclimate data. Sediment cores records from the Cariaco Basin just off the coast of South America in the Atlantic showed a strong coupling between the ITCZ location and high-latitude climate change over long ranges of time scale, from decadal to glacial interglacial transitions (Hughen et al., 1996; Lea et al. 2003). Modelling studies that replicated the LGM conditions also found a relationship between interannual high-latitude climate variability and the position of the ITCZ. Chiang et al. (2003) used the Community Climate Model version 3 (CCM3) atmospheric general circulation model (AGCM) coupled to a slab ocean to determine the response of the Atlantic ITCZ to the Last Glacial Maximum (LGM) conditions. In the model changes in sea ice coverage and the Meridional Overturning Circulation (MOC) resulted in the NH cooling and a southward shift of the ITCZ, indicating the ITCZ shifts towards the warmer hemisphere when a hemispheric heating asymmetry exists. Chiang and Bitz (2005) built on these results by looking at the Hadley cell response and proposing a mechanism for which the high latitude heating anomaly affected the tropics. Chiang and Bitz used the same CCM3 AGCM coupled to a 50-m slab ocean with imposed ice coverage similar to the LGM. They found that in the high-latitude regions in which the ice coverage increased the air cooled and dried. The cold anomalies were transported equatorward, inducing an increase in the Hadley cell subsidence and a shift of the ITCZ away from the cooler hemisphere. The ITCZ shift resulted in an anomalous moisture transport away from the colder, drier hemisphere and into the warmer hemisphere. Chiang and Bitz hypothesized that wind-evaporation-sst (WES) feedbacks were important in the equatorward progression of SST anomalies initiated in the

16 12 extratropics. They also noted the ITCZ displacement led to an increase in incoming radiation in the cooler hemisphere which offset the increased outgoing radiation in high latitudes due to the increased ice coverage. The effect on the Hadley cell and ITCZ shift appeared to be independent of the longitudinal position of imposed ice. Building on the LGM model studies, Broccoli (2006) used two model configurations coupled to ocean models of different complexities. For one simulation the atmospheric model was coupled to a slab ocean of 50-m depth with prescribed sea ice, CO2 concentrations, sea level, and orbital parameters consistent with the LGM. The second simulation coupled the atmosphere to a dynamical ocean. Like with the previous studies Broccoli found the imposed sea ice coverage led to a hemisphere heating asymmetry and an ITCZ shift towards the warmer hemisphere. The heating anomalies led to changes in the extratropical SST anomalies which propagated equatorward through an atmospheric bridge that allowed temperature changes at high latitudes to affect ITCZ position. Broccoli also found the relationship between ITCZ displacements and asymmetric extratropical forcing were more prominent in the atmosphere-slab ocean model than in the coupled model suggesting ocean dynamics are important in determining the response of the ITCZ to hemispheric heating anomalies ITCZ shifts in CMIP simulations The previous studies have demonstrated there is a relationship between changes in c-eq heat transport and the position of the ITCZ in idealized models. The response is dependent on whether the atmosphere is coupled to a slab ocean or dynamical ocean. The idealized slab ocean studies demonstrated the longitude in which the radiative forcing is applied does not affect the ITCZ shift unless the forcing is applied in the tropics. Conversely, the latitude in which the forcing is applied matters since the feedbacks that arise in response to the forcing are latitude dependent. Moving

17 13 forward, to determine how the ITCZ will shift in the coming century, we need to understand whether these relationships appear in fully coupled models with dynamical oceans and increased carbon dioxide (CO2) concentrations. To better understand how a dynamical ocean and warming impact the ITCZ we can look at the Coupled Model Intercomparison Project (CMIP) output. Several studies in the past ten years have looked at the relationship between the c-eq heat transport changes and ITCZ shifts in the 20 th century comprehensive GCMs and in different Intergovernmental Panel on Climate Change (IPCC) representative concentration pathways (RCP). Hwang et al. (2013) performed a comprehensive energetic analysis on CMIP5 20 th century simulations and attributed biases in TOA radiative fluxes to the double-itcz problem, a bias which exists in most GCMs still used today. The double-itcz bias results in excessive precipitation is in the SH tropics and the appearance of two bands of precipitation around the equator. This feature is not seen in observations and has been theorized to be the result of cloud biases over the Southern Ocean in GCMs that introduce anomalous warming in the SH (Hwang et al., 2013; Kay et al., 2016). When the radiative fluxes in CMIP5 historical simulations were compared with satellite observations of the Earth s energy budget the authors found the GCMs tended to overestimate the incoming SW radiation in the SH. In their analysis Hwang et al. found the c-eq AHT, precipitation asymmetry relationship was remarkably linear in the 20 th century CMIP5 GCMs. They found changes in SW scattering led to a southward heat transport for all the models. Hwang et al attributed an observed southward shift of tropical precipitation in the late 20th century to the radiative cooling effect of anthropogenic sulfate aerosols predominantly in the NH. Changes in aerosol concentrations in the 20 th century have also been linked to droughts in the Sahel in North Africa (Rotstayn and Lohmann 2002, Cowan and Cai 2011). Climate model studies

18 14 and observations suggest that the observed 20th century increase in NH aerosols relative to the SH led to a southward shift in tropical precipitation in the early 20 th century. A study completed by Rotstayn, Collier, and Luo in 2015 showed that there is a correlation between the hemispheric averaged heating and ITCZ shifts in 21st century CMIP5 RCP4.5 climate simulations. However, the location of the ITCZ and how it was expected to change in a warmer climate varied significantly in the models. Rotstayn also compared the changes in c-eq heat transport for models that include indirect aerosol effects on cloud albedo and cloud lifetime as well as direct aerosol effects to models that only include direct aerosol effects. They found the projected tropical SST and precipitation changes were sensitive to indirect aerosol effects and hypothesized the southward shift of the ITCZ seen towards the latter half of the 20th century was a result of increased aerosol emissions in the NH. They also hypothesized the ITCZ would shift north in the coming 21st century due to a reduction in NH aerosols emissions as a result of stricter regulations. Allen (2015) expanded on Rotstayn et al. (2015) results by performing a similar analysis on all CMIP5 RCP experiments, as well as miscellaneous forcing experiments with the Community Atmosphere Model version (CAM3) and the Community Earth System Model version 1.2. Like Rotstayn et al., Allen hypothesized decreases in NH aerosol emissions will result in a northward shift the ITCZ in the 21 st century. Allen expanded on Rotstayn s results by looking at the ITCZ shifts by ocean basin. Allen found the largest tropical precipitation shift occurred in the Pacific basin, with minor shifts in the Atlantic basin. He related the Pacific-Atlantic ITCZ difference to the difference in the aerosol changes in the Pacific and Atlantic basins. In addition, he hypothesized changes in the Meridional Overturning Circulation (MOC) northward c-eq heat transport counteracted the aerosol clean up in the Atlantic, leading to less of an ITCZ shift, indicating the c-eq OHT for each basin is important in determining the ITCZ shift.

19 Ocean-atmosphere c-eq heat transport split In the Broccoli (2006) and Allen (2015) studies the ocean c-eq heat transport was found to be important when evaluating the ITCZ response to changes in radiative heating. This result is not surprising considering the ocean can be transport up to 5 PW across the equator on inter-annual timescales (Fasullo and Trenberth 2008). Outside of the tropics the atmosphere is responsible for most the poleward heat transport (~80% of total poleward transport in the winter hemisphere). This is partially due to the ocean eddy heat transport efficiency decreasing in the extratropics (Jayne and Marotzke 2002). Figure 1.1 shows the annual mean ocean, atmosphere and net poleward heat transport is shown below. Figure 1.1. Atmosphere, ocean and net poleward heat transport. [Fassullo and Trenberth, 2008]. Several recent modelling studies have looked at the c-eq AHT and OHT response to radiative heating anomalies. In the base state, most ocean c-eq heat transport occurs in the MOC (Trenberth and Solomon, 1994; Frierson et al 2013) with some c-eq OHT occurring in the wind-driven gyres

20 16 in all basins (e.g. Joyce 1988). However, whether the deep MOC or surface gyres respond to changes in radiative fluxes has been shown to depend on where the forcing is applied. Cheng et al. (2008) evaluated the c-eq ocean-atmosphere heat transport induced when a sudden freshwater flux was added to the North Atlantic, simulating an abrupt slowdown of the MOC. Adding freshwater in the North Atlantic resulted in a decrease in the MOC northward OHT and led to cooling in the northern high latitudes. This cooling in the NH induced a northward c-eq AHT and a southward shift of the ITCZ indicating a strong coupling between the high latitudes and tropics. Cheng et al. found the mechanism that led to a southward ITCZ shift were similar to those found when simulating the LGM. The Hadley cell shift resulted in a strengthening the surface trade winds in the NH and weakening them in the SH. The associated Ekman transport led to anomalous northward ocean circulation in both hemispheres, transporting heat from the SH into the NH counteracting the initial freshwater forcing. This response indicates the ocean and atmosphere are coupled not only by energy balance, but also by dynamics. Ferrari and Ferreira (2011) provide an explanation of how the OHT adjusts to heating anomalies. Ferrari and Ferreira ran long simulations of a fully dynamical ocean model and looked at the partitioning of OHT by wind-driven gyres and the deep circulation. They found the majority of the heat transport in the Pacific and Indian oceans was due to wind-driven gyres in the thermocline. In the Atlantic 60% of the OHT occurred below the thermocline in the deep ocean and the remaining 40% of the OHT was by gyres in the upper ocean. They also found the Atlantic OHT below the thermocline was very sensitive to the surface winds across all latitudes, suggesting that the atmosphere-ocean wind coupling at all latitudes can influence the c-eq transport. Haywood et al. (2016) looked at the c-eq response when SH albedo biases were corrected. As mentioned earlier, an issue that is still present in GCMs today is a double ITCZ in the base state.

21 17 The double-itcz has been theorized to be the result of cloud biases over the Southern Ocean in GCMs (Hwang et al., 2013; Kay et al., 2016). Haywood ran a serious of modelling experiments in the Hadley Centre Global Environment Model version 2-Earth System (HadGEM2-ES) coupled climate model in which they adjusted the SH albedo. Equilibrating the albedos induced a northward c-eq heat transport, with the ocean contributing ~35% of the increased northward c-eq heat transport. The double-itcz bias was lessoned as a result as well. Kay et al. (2016) and Hawcroft et al. (2016) ran similar experiments to Haywood but instead of changing the albedo of the entire SH they only increased the albedo of the Southern Ocean from 40S -70S. Kay modified the shallow convection detrainment to increase the supercooled cloud liquid concentration and brightened the clouds over the Southern Ocean in the CESM fully coupled GCM. Hawcroft increased the albedo of the Southern Ocean in the HadGEM2-ES through adjustments to stratospheric aerosol optical depth, ocean albedo and cloud droplet number concentration. Brightening the Southern Ocean led to the SH cooling and a southward c-eq AHT and OHT. Despite using different approaches to brighten the Southern Ocean and using different GCMs, both Kay and Hawcroft found the atmosphere transported about 20% of the southward c- eq heat transport and the ocean was responsible for the remaining 80%. The fact that Kay and Hawcroft got the same resulting change in c-eq AHT and OHT indicates the split between the atmosphere and ocean c-eq heat transport is potentially only dependent on the location in which heat is added or removed and not the method for removing the heat. In addition, the results may not be model dependent since Kay and Hawcroft got the same result using two different models Causes for hemispheric energy imbalance The previous studies mentioned show that there clearly is a relationship between hemispheric heating asymmetries (i.e. c-eq heat transport) and ITCZ shifts. It has also been shown that changes

22 18 in aerosol concentrations and upward ocean heat flux can cause shifts in the ITCZ. The relationship between heating anomalies induced by aerosol, OHT, cloud, water vapor, and other radiative forcings and the ITCZ shift is dependent on the latitude in which the heating is applied. Therefore, to understand how the ITCZ will shift in the future we need to consider how the hemisphere radiative asymmetries and upward ocean heat flux may change; what their impact on the hemispheric heating asymmetry will be; and the latitude in which the change will occur. Changes in hemispheric heating asymmetries due to changes in the aerosol concentrations in the 20 th century have been shown to cause shifts in the ITCZ (Rotstayn and Lohmann 2002, Cowan and Cai 2011, Hwang 2013, Rotstayn et al. 2015). Looking forward towards the 21 st century, aerosol emissions are projected to decrease over North America and much of Asia decreasing the total hemisphere albedo and increasing the absorbed SW radiation in the NH (Moss et al. 2010, Lamarque et al 2011). Differences in aerosol reductions in 21 st century GCMs result in hemispheric averaged heating asymmetry (i.e. one hemisphere is warmer than the other). These differences result in varying shifts of the ITCZ but most likely a northward shift due to a larger clean-up of aerosols occurring in the NH than SH. ITCZ shifts have also been attributed to changes in the Atlantic Meridional Overturning Circulation (AMOC). The AMOC transports heat from the southern and northern tropics to the north Atlantic, warming the NH. In the base state the annual mean ITCZ location has been attributed the greater northward than southward heat transport by the AMOC (Kang et al. 2015). The IPCC report states it is very likely that the AMOC will weaken in the 21 st century which would result in less energy being transported to the north Atlantic. If the heat transport towards the north Atlantic decreases the NH would cool resulting in a southward shift of the ITCZ (Cheng et al.

23 ). If the rate at which the AMOC slows down is inconsistent between the models the change in the location of the ITCZ from 2006 to 2100 in the 21 st century would vary from model to model. The ITCZ response to radiative heating asymmetries is not only dependent on changes in the OHT but also on changes in the ocean heat storage. From 2005 to 2014, 90% of the ocean heat uptake occurred in the SH (Llovel and Terray 2016). The ocean heat content in the SH increased linearly over the ten-year period whereas the ocean heat content in the NH exhibited relatively minor changes. The largest ocean heat uptake occurred between 25 S and 50 S in the Indian and Pacific Ocean. If the distribution of ocean heat content changes and the NH begins taking up more heat in the 21 st century the AHT would change to compensate for the new spatial distribution and the ITCZ may shift was a result. The ice-albedo feedback or changes in ice and snow coverage can also affect the location of the ITCZ. Ice and snow in the high latitudes melt as the climate warms. When the snow and ice melts the reflectivity (or albedo) decreases, resulting in more SW radiation being absorbed in the NH. In an initial analysis of the change in SW radiation in CMIP5 RCP8.5 models, it was found that the Arctic SW radiation change more than the SW radiation over Antarctica. This led to the average NH minus SH energy flux asymmetry increasing and the Hadley cell shifting north to bring energy to the SH. A fourth factor that may impact the location of the ITCZ in GCMs is the longwave (LW) radiative feedback. As surface temperatures increase we expect the outgoing LW (OLR) to increase resulting in a cooling. Since there is more land in the NH we expect the mean surface temperature in the NH to increase more than the mean surface temperature in the SH leading to more OLR in the NH. Greater OLR in the NH will induce a northward c-eq AHT and southward ITCZ shift.

24 20 Finally, clouds can significantly affect the cooling and heating of the atmosphere. How clouds will change in a warmer climate is not fully understood, therefore it is not clear whether cloud changes will cause a northward or southward shift. In the analysis proposed it will be hard to treat cloud changes as solely feedbacks or forcings. For example, if the ITCZ shifts towards one hemisphere the reflected SW radiation for that hemisphere will increase. Alternatively, cloud changes due to sea surface temperatures can be either a feedback or a forcing. If the AMOC slows down the water in the north Atlantic may cool which increases low cloud coverage. In this case the cloud changes could be grouped with the AMOC forcing or be considered a positive feedback triggered by the north Atlantic cooling. Each forcing will be considered separately but may be too complex to clearly state whether the change is a forcing or feedback. 1.2 GUIDING QUESTIONS As demonstrated by the studies discussed above, the location of the ITCZ is highly dependent on the c-eq AHT. The dependence of the ITCZ on the c-eq AHT has appeared in 20 th century observations and is expected to persist into the 21 st century. The southward shift of the ITCZ in the latter half of the 20 th century has been attributed to changes in aerosol concentrations (Hwang, 2013). In the 21 st century aerosol concentrations are expected to continue to decline in the NH. Studies looking at the moderate emissions pathway, RCP4.5, have also attributed shifts in the ITCZ projected in the 21 st century CMIP5 models to changes in aerosol concentrations. Therefore, we hypothesized changes in SW clear-sky scattering as a result of aerosol clean up to be the leading cause of ITCZ shifts in the RCP8.5 models. Idealized studies have shown the zonal structure of the ITCZ is dependent on the latitude of the thermal forcing (Seo et al., 2014a). However, it is unclear how the latitude of the forcing impacts the total ITCZ shift in a fully coupled model. Seo, Kang, and Frierson used an idealized

25 21 model to show the zonal structure of the ITCZ is more sensitive to thermal forcings applied at lower latitudes than those applied at higher latitudes. If this relationship holds in a fully coupled system, we may also expect differences in the zonal extent of the forcing at higher latitudes to not have a significant impact on the structure of the ITCZ since the thermal forcing is well mixed at higher latitudes. As the forcing gets closer to the tropics, we hypothesize that the structure of the ITCZ will become more sensitive to the zonal extent of the forcing. Whether the position of heating anomalies outside of the tropics impact the c-eq OHT-AHT split has not yet been investigated. Despite using different approaches to cool the entire Southern Ocean in two different models, Kay and Haywood found the same split in the atmosphere and ocean c-eq heat transports. However, when the albedo of the entire SH was modified in the Hawcroft study the authors found a different split between the atmosphere and ocean northward c-eq heat transport indicating the method for inducing a cooling may not impact the c-eq OHT- AHT split but the locations cooled may matter. Preliminary analysis on the CMIP5 RCP8.5 GCMs show varying ITCZ shifts. To better understand the cause for the model differences we investigated three primary questions: (1) What are the primary thermal forcings that affect the location and structure of the ITCZ in the RCP8.5 scenario and what are their relative effects? (2) How does the location of the thermal forcing affect the ITCZ? (3) How does the ratio of ocean to atmosphere c-eq heat transport depend on the location of the heating anomaly?

26 22 Chapter 2. CMIP5 RCP8.5 MOIST STATIC ENERGY BUDGET 2.1 THE GLOBAL ATMOSPHERE ENERGY BUDGET Sixteen CMIP5 models forced with the business as usual emissions pathway, RCP8.5, were included in the analysis. In the RCP8.5 scenarios, the greenhouse gas (GHG) emissions continue to increase throughout the 21 st century resulting in a forcing of 8.5 W m -2 in the year Model output was available for a total of 23 models, however, seven models were removed from the analysis due to imbalances in the atmospheric energy budget. If kinetic energy is neglected, the poleward moist static energy (MSE) transport can be calculated by integrating the net radiative, latent and sensible heat flux zonally and 90 S or 90 N to a give latitude (λ). The equation for poleward MSE transport (F A ) at a given latitude is shown below. λ π 2 2π 0 F A (λ) = (F R + F SH + L M Sn + L V E + L S Su) a 2 cosλ dφ dλ (2.1) Where F R is the net radiative flux into the atmosphere, F SH is the net sensible heat flux into the atmosphere, L M is the latent heat of fusion/melting (3.35e5 J kg -1 ), L V is the latent heat of evaporation (2.5e6 J kg -1 ), E is the evaporation rate, LS is the latent heat of sublimation and Su is the sublimation rate. In Equation 2.2 the sum of these terms will be referred to as Q A, or the net radiative, latent and sensible heat flux into the atmosphere column. In Equation 2.1 and 2.2 a is the Earth s radius, λ is the latitude and φ is the longitude. If integrated from pole to pole, equation 2.1 gives the global atmospheric heat content. The globally integrated atmospheric heat content for the periods of 2006 through 2025 and 2081 through 2100 for all 23 models in which model output was available in the CMIP archive is shown in Table 2.1. The change in the atmospheric heat content calculated as the difference between the

27 23 last and first 20 years is also shown in Table 2.1. For the RCP8.5 scenarios we would expect the globally integrated atmospheric heat content to increase slightly due to the increase in moisture in the atmosphere. However, the increase in the atmospheric heat content due to the increase in moisture would not be as large as the change in globally integrated MSE seen in the RCP8.5 GCMs. If we assume the global atmospheric energy imbalance is distributed evenly between the northern and southern hemisphere, we can calculate the change in c-eq AHT by comparing the total MSE transport into each hemisphere separately. Using this approach, the atmospheric energy imbalance is removed from both hemispheres and essentially neglected from the change in c-eq AHT calculation. The resulting equation, which is commonly used in ITCZ studies, is (Frierson and Hwang, 2012; Donohoe and Battisti, 2012; Hwang et al, 2013; Allen 2015): 0 π 2 2π 0 π 2 0 AHT = 1 [ Q 2 Aa 2 cosλ dφ dλ + Q A a 2 cosλ dφ dλ] (2.2) 2π 0 For the models in which the model grid area was available, the net energy into the column was converted from W m -2 flux to W by multiplying by the grid area instead of integrating using Equation 2.2. The mean for the first 20 years, last 20 years and the change in c-eq AHT using Equation 2.2 are shown in Table 2.1. The change in c-eq AHT is on the order of 0.1 PW and for some models is less than the change in globally integrated MSE.

28 Model Table 2.1. Global atmosphere (ATM) energy imbalance and c-eq AHT values for beginning and end of 21st century RCP8.5 CMIP5 GCMs to 2025 mean 2081 to 2100 mean (2081 to 2100 mean) (2006 to 2025 mean) Δ ATM ATM ATM Δ ATM energy energy energy c-eq energy Δc-eq imbalance imbalance c-eq AHT imbalance AHT imbalance AHT / Δc-eq (PW) (PW) (PW) (PW) (PW) (PW) AHT Exclude ACCESS No ACCESS No bcc-csm Yes CanESM Yes CCSM No CSIRO-Mk No FGOALS-s No GFDL-CM No GFDL- ESM2G No GFDL- ESM2M Yes GISS-E2-R No GISS-E2-R- CC No IPSL-CM5A- LR No IPSL-CM5A- MR No IPSL-CM5B- LR No MIROC No MIROC-ESM- CHEM Yes MPI-ESM-LR Yes MPI-ESM-MR No MRI-CGCM Yes MRI-ESM Yes NorESM1-M No NorESM1-ME No 24

29 ERRORS IN THE GLOBAL CLIMATE MODELS The changes in the global atmospheric energy imbalance seen in the CMIP5 models and shown in Table 2.1 may be a result of errors in the models. Lucarini and Ragone (2010) calculated the global atmospheric heat content for 22 CMIP3 GCMs with preindustrial conditions with fixed CO2 concentrations at 280 ppm. The globally integrated atmosphere heat content was calculated for a 100-year period when the models were considered in steady state. Lucarini found most models featured biases of the order of 1 W m 2 for the net global, atmospheric, oceanic and land energy balances. The imbalances were not a result of transient effects but of the imperfect closure of the energy cycle in the fluid components and inconsistencies over land. Whether these errors were fixed for all the models in the updated CMIP5 is unknown. Imbalances in the atmosphere heat content may also be a result of moisture biases. Liepert and Lo (2012) compared the atmospheric moisture budget for 30 CMIP5 model control runs. Of the 30 models, six showed significant biases in the global moisture cycle (BNU-ESM, FGOALSs2, MIROC5, MPI-ESM-P, MPI-ESM-MR, and MPI-ESM-LR) that resulted in non-radiative forcings of about 0.15 W m 2. Globally this equates to about a PW bias in the atmosphere energy budget. In our analysis of the RCP8.5 model output we discovered errors in the model latent heat outputs. Latent heat enters the atmosphere through evaporation and sublimation. Sublimation is the physical process in which a solid (snow or ice) directly converts to a gaseous (water vapor) state without going through a liquid state. The IPCC requires that the CMIP5 surface upward latent heat flux output (labeled hfls in CMIP5 archive) to include both the latent heat from evaporation and sublimation (Taylor, 2013). The net latent heat flux at the surface can be calculated by summing the evaporation and sublimation fluxes.

30 26 LE = L v E + L S S (2.3) When the latent heat output is compared to the latent heat release from evaporation nonzero values should appear in regions in which we would expect sublimation to occur. Sublimation occurs in locations where the relative humidity is low, there is a lot of sunlight, dry winds are present and the pressure is lower (e.g. mountain ranges). Figure 2.1 shows the mean sublimation for a 20-year period for RCP8.5 runs in which sublimation data is available (6 of the 23 models). Figure 2.1. CMIP5 RCP8.5 sublimation in W m -2. Sublimation values in kg m -2 s -1 were converted by multiplying by the latent heat of sublimation (2.83e6 J kg-1). Means shown are for 2006 through For the models shown in Figure 2.1, the sublimation rate ranges from 1 W m -2 to over 50 W m -2 (GISS-E2-R and GISS-E2-RCC). The high sublimation rates for the GISS models indicate an error in the sublimation output. The output is either labeled incorrectly or in the incorrect units.

31 27 Not all models appear to include sublimation in the latent heat output field. In difference plots comparing the latent heat flux to the evaporation flux (converted to W m -2 using the latent heat of vaporization), several models show no difference between the two values (Figure 2.2). If the latent heat output for the CMIP5 models includes sublimation, plots of L V E minus LE should be nonzero in locations in which sublimation occurs. Figure 2.2. Difference between the latent heat flux and evaporative heat flux for CMIP5 RCP8.5. Mean upward latent heat flux at the surface minus the evaporation rate (kg m-2 s-1) x 2.3e6 J kg-1 for 2006 through 2025.

32 28 As shown in Figure 2.2, The spatial pattern and magnitude of the difference between the evaporative flux and latent heat flux is not uniform across the models. This could be due to a few reasons: 1. Models in which the difference between evaporative flux and latent heat flux is zero everywhere do not include sublimation in the latent heat flux calculation, 2. the amount of sublimation in the models varies significantly with very little sublimation in the models with zero difference between the evaporative and latent heat flux, or 3. the models shown do not use the same latent heat of sublimation constants when calculating latent heat flux from sublimation. Several models show a difference in the evaporative and latent heat flux terms (ACCESS1-0, ACCESS1-3, bcc, CanESM2, CCSM4, FGOALS, IPSL, MPI and NorESM). The differences range from 1 to 10 W m -2, indicating little model agreement on the mean sublimation for 2006 through For the MRI models, the evaporative and latent heat flux are different over most of the tropics. This indicates the MRI models use a different latent heat of vaporization constant. The CSIRO and GFDL models show no difference between the latent heat flux and evaporative flux. This indicates the sublimation is either omitted from the latent heat output or the sublimation is anomalously small in these models. For the models in which sublimation output is available (MIROC and MRI) plots of the difference between the latent heat from sublimation and the surface upward latent heat flux output (Figure 2.3) show the latent heat flux output does not include sublimation. Figure 2.3 shows a plot of the latent heat flux minus evaporation and sublimation, converted from kg m -2 s -1 to W m -2. All four models show nonzero values in regions in which we would expect sublimation to occur, indicating none of the four models include sublimation in the surface upward latent heat flux term.

33 29 Figure 2.3. Upward latent heat flux at the surface minus the evaporation rate (kg m-2 s-1) x 2.3e6 J kg-1 and sublimation (kg m-2 s-1) x 2.83e6 J kg-1 for CMIP5 RCP8.5 models. Mean shown for 2006 through From the checks listed above we can infer that the following models do not include sublimation in the surface latent heat flux (hfls) output: CSIRO-Mk3-6-0, GFDL-CM3, GFDL-ESM2G, GFDL- ESM2M, GISS-E2-R, GISS-E2-R-CC, MIROC5, MIROC-ESM-CHEM, MRI-CGCM3, MRI- ESM1. The following models include latent heat release from sublimation in the latent heat (hfls) output and the sublimation contribution is very small (< 0.5 W m -2 ): ACCESS1-3, CCSM4, MPI-ESM- LR, MPI-ESM-MR, NorESM-ME. The following models include latent heat release from sublimation in the latent heat (hfls) output and the sublimation contribution is greater than 2 W m -2 in some regions: ACCESS1-0, bcc-csm1-1, CanESM2, FGOALS-s2, IPSL-CM5A-LR, IPSL-CM5A-MR, IPSL-CM5B-LR

34 30 Changes in sublimation can account for up to 26% of the change in c-eq AHT, indicating excluding sublimation from the c-eq AHT calculation can result in significant errors in the change in c-eq AHT estimates. Table 2.2 shows the change in c-eq AHT due to sublimation changes alone and the change in c-eq AHT excluding sublimation for the four models in which the sublimation output is available. The change in c-eq AHT due to sublimation changes was calculated by substituting in the sublimation latent heat contribution in W m -2 in for QA in Equation 2.2. Table 2.2. Change in global latent heat flux due to sublimation and c-eq AHT due to sublimation for RCP8.5 models in which sublimation data is available. Δ c-eq AHT due to sbl / Δ c-eq AHT without sbl Model Δ global latent heat due to sublimation changes (PW) Δc-eq AHT due to sublimation changes (PW) Δ c-eq AHT (without sbl) MIROC MIROC-ESM CHEM MRI-CGCM MRI-ESM The imbalance in the RCP8.5 models could be a result of neglecting sublimation, transient effects, numerical modelling errors, increases in the global moisture budget, or other causes not yet identified. Since we are unable to identify the cause of the change in the globally integrated atmospheric heat content for each model or determine which hemisphere is impacted we excluded models in which the change in globally integrated atmospheric heat content was larger than the change in c-eq AHT or models in which the globally integrated heat content was anomalously large (MIROC-ESM-CHEM) from the analysis. Models that were excluded from the analysis discussed in the rest of this study are shown in yellow in Table 2.1, leaving 16 models for the remainder of the analysis.

35 31 Chapter 3. ATTRIBUTION OF ITCZ SHIFTS 3.1 QUANTIFYING ITCZ SHIFTS IN THE 21 ST CENTURY Previous studies have quantified ITCZ shifts by using a simple index that compares the area weighted mean precipitation from the equator to 20 N to the area weighted mean precipitation from 20 S to the equator (Hwang et al., 2013; Allen, 2015; Rotstayn et al., 2015). For this study, we used the same metric. PR ASYM = PR 0 to 20N PR 0 to 20S (3.1) In Equation 3.1 PR 0 to 20N is the area weighted mean precipitation from the equator to 20 N, PR 0 to 20S is the area weighted precipitation from 20 S to the equator. The change ( ) is calculated as the difference between the last 20-year mean asymmetry (2081 through 2100) and first 20-year mean (2006 through 2025). If the precipitation from the equator to 20 N increases more than the precipitation from 20 S to the equator we can infer that the annual and zonal mean location (over land and the ocean) of the ITCZ shifted north. A southward shift is represented by a negative value (i.e. the precipitation from 20 S to the equator increases more than the precipitation from the equator to 20 N). The change in c-eq AHT over the 21 st century is also calculated by comparing the last 20-year mean to the first 20-year mean. A plot of the change in c-eq AHT versus change in precipitation asymmetry is shown in Figure 3.1. Only the 16 models that showed a smaller change in c-eq AHT than global atmosphere heat content and did not have anomalously large atmosphere energy imbalances are shown. For models in which multiple ensemble members are available, each ensemble member was plotted separately.

36 Δ PR 0-20N - Δ PR 0-20S (mm day -1 ) 32 (2081 to 2100 mean) (2006 to 2025 mean) R = -0.77; p = 6.11e-5 Δ c-eq atmosphere heat transport (PW) Figure 3.1. Change in c-eq atmospheric heat transport versus change in tropical precipitation asymmetry defined as the precipitation from 0 to 20N minus precipitation from 0 to 20S. There is a strong linear correlation between the change in c-eq AHT and the change in the precipitation asymmetry (Figure 3.1). When thirty year periods were used to calculate the change in c-eq AHT and precipitation asymmetry the linear relationship persisted. The correlation coefficient is -0.7 indicating an increase in northward AHT (shown as a positive AHT value) results in a southward ITCZ shift. This matches the findings of studies that used 20 th century GCM output (Hwang et al., 2013), that have looked at RCP4.5 GCM simulations (Allen, 2015; Rotstayn et al., 2015) and observational data from the 20 th century (Folland et al., 1986; Rotstayn and Lohmann, 2002; Chang et al., 2011; Hwang et al., 2013). This result also aligns with our understanding of relationship between the ITCZ, northern and southern Hadley cells and poleward atmospheric heat transport. If a hemispheric heating asymmetry exists, the Hadley cells (and thus

37 33 ITCZ) shift across the equator to transport energy from the hemisphere with more heating to the hemisphere with less heating. This results in the location of the ITCZ to be dependent on the c-eq AHT as is shown in Figure 3.1. Models that include direct effects of aerosols, indirect effects aerosols on cloud albedo and indirect effects of aerosols on cloud lifetime (red) show the most significant ITCZ shifts. Models that include direct effects of aerosols and only indirect effects of aerosols on cloud albedo are shown in green. Models that only include direct effects of aerosols are shown in blue. The ITCZ shift in the green and blue models is minimal compared to the red models, indicating aerosol indirect effects on cloud lifetime potentially have a greater impact on the hemispheric heating asymmetry than direct effects. An interesting feature of the plot of the change in c-eq AHT and precipitation asymmetry is the positive intercept. The positive intercept implies a northward shift of the ITCZ in absence of a change in c-eq AHT. However, when a 95% confidence interval is added to the plot the positive intercept is shown to be insignificant. The equations used to add a 95% confidence interval to the plot are shown in Appendix A. Figure 3.2 shows the change in c-eq AHT versus the precipitation asymmetry over the Pacific, Atlantic and Indian ocean basins. The change in c-eq AHT and precipitation asymmetry is calculated only using values over the ocean. The values are also weighted by the area of each basin. Over the Indian Ocean, all the models show a northward ITCZ shift. Over the Atlantic and Pacific basins about half of the models show a northward shift while the other half show a southward shift.

38 34 Figure 3.2. Change in precipitation asymmetry over each ocean basin versus the total change in c-eq AHT. Values over land excluded when calculating change in precipitation asymmetry over ocean basin. Change calculated as mean for 2081 through 2100 minus the mean for 2006 through Figure 3.3 shows the change in precipitation (2081 through 2100 mean minus 2006 through 2025 mean) in the tropics. Almost all the models show a larger increase in precipitation in the Pacific and Indian Oceans than the Atlantic. Most models also show an increase in precipitation in the northern part of the Indian Ocean and in some cases a drying of the southern tropics over the Indian Ocean. This indicates the Monsoon over southern Asia may strengthen with warming. In the Pacific, some models show an increase in precipitation right on the equator from about 5 S

39 to 5 N and some show more of a precipitation increase in the northern part of the tropical Pacific than the southern part. None of the models show a significant southward ITCZ shift in the Pacific. 35 mm day -1 Figure 3.3. Change in tropical precipitation (mm day -1 ). Change calculated by subtracting the mean for 2006 through 2025 from the mean for 2081 through 2100.

40 ATTRIBUTION OF THE CHANGE IN C-EQ HEAT TRANSPORT Figure 3.1 shows changes in c-eq AHT affect the location of the ITCZ, represented as the precipitation asymmetry defined in Equation 3.1. In Figure 3.1 there is a wide spread in how the c-eq AHT, and thus precipitation asymmetry, is expected to change in the business as usual emissions scenario. To improve how the change in c-eq AHT is modeled we need to understand what radiative, latent and sensible heat fluxes affect the c-eq AHT in the RCP8.5 GCMs and how they differ across the models. The GCMs used in this analysis provide separate outputs for shortwave (SW) and longwave (LW) radiative flux at the TOA and surface. Using Equation 2.2 and taking the difference between the first and last 20 years we can calculate c-eq AHT induced by the change in net TOA (LW + SW), net TOA LW and net TOA SW. To determine the factors that caused the net incoming SW radiation to change we use the approximate partial radiative perturbation (APRP) method (Taylor et al., 2007). Using the APRP method the change in c-eq AHT due to changes in surface albedo (i.e., ice/snow melting), clear-sky scattering, clear-sky absorption, and cloud effects are determined. The c-eq AHT induced by the change in net upward ocean (surface) flux is calculated by summing the sensible, latent and radiative fluxes at the surface. For 12 of the 16 models OHT data was available in the CMIP5 archive (labeled hfy in the archive). By zonally integrating the northward heat transport at the equator we can determine the net northward OHT for a given period. Adding the change in northward OHT to the change in c-eq AHT induced by the upward ocean heat fluxes we can determine the impact of ocean heat uptake on the AHT. The c-eq heat transport induced by changes in incoming radiative and upward surface fluxes are shown in Figure 3.4. Panel A in Figure 3.4 shows the total change in northward c-eq AHT

41 37 (same values as shown in Figure 3.1). Panel A also shows the breakdown between surface and net TOA fluxes (LW + SW). Panel B shows the c-eq AHT induced by TOA SW radiation changes. Panel C shows the c-eq AHT induced by TOA LW radiation changes. Panel D shows the c-eq AHT induced by changes in surface flux and the split between northward OHT and storage changes that sum to the total surface flux change. All panels show the change in northward c-eq AHT induced by each term represented as the last minus first 20-year mean.

42 c-eq AHT induced (PW) 38 A Δ c-eq AHT TOA LW TOA SW Ocean flux B C TOA SW TOA SW clouds TOA SW from 15 S to clouds 15 N excluding 15 S to 15 N TOA SW clear-sky changes D TOA SW surface albedo changes TOA SW noncloud scattering TOA SW due to noncloud absorption TOA LW TOA LW from 15 S to 15 N TOA LW excluding 15 S to 15 N Figure 3.4. Attribution of the change in c-eq AHT. For all terms a northward c-eq AHT (positive value) is induced by an increase in heat entering the southern hemisphere atmosphere. For the OHT term, a positive value indicates the northward OHT decreases, inducing a northward AHT. For the ocean storage term a positive value indicates the heat uptake in the northern hemisphere is greater than the heat uptake in the southern hemisphere, inducing a northward c-eq AHT. Change is calculated as the mean c-eq AHT for 2081 through 2100 minus the mean c-eq AHT for 2006 through Ocean flux Ocean heat content -OHT -Pacific and Indian OHT -Atlantic OHT Effects of the AMOC slowdown As discussed in Chapter 2, we expect the Atlantic Meridional Overturning Circulation (AMOC) to slow down in the 21 st century which will result in less heat being transported from the

43 39 tropics to the north Atlantic. This results in the northern hemisphere (NH) cooling and induces a northward c-eq AHT and southward ITCZ shift. Figure 3.4 shows changes in the upward surface flux induce either no change in AHT or a northward AHT for all 16 models. Figure 3.5 shows the spatial change in upward surface flux. All the models show a cooling in the north Atlantic, which is consistent with what we would expect if the AMOC northward OHT slowed. Figure 3.5. Change in upward surface flux (LW + SW + sensible + latent heat fluxes). Change calculated as the mean for 2081 through 2100 minus the mean for 2006 through 2025.

44 c-eq AHT induced (PW) 40 The sum of the upward surface flux and the northward OHT terms in panel D of Figure 3.4 gives the impact of changes in the ocean heat content on the AHT. Changes in ocean heat content counteract the decrease in northward OHT slightly, inducing a southward AHT. The negative sign on the change in ocean heat storage in panel D indicates the ocean in the southern hemisphere (SH) uptakes up more heat than the NH. Figure 3.6 shows the change in c-eq AHT induced by OHT changes in the Pacific and Indian Ocean combined and by OHT changes in the Atlantic. All the models show the northward OHT in the Atlantic decreasing, which is consistent with an AMOC slowdown. If the AMOC slows down in the 21 st century less heat would be transported from the tropics to the north Atlantic by the ocean resulting in the NH cooling. This would result in an increase in northward AHT as is shown by the positive values in Figure 3.6. All basins Pacific & Indian Atlantic Figure 3.6. Change in c-eq AHT induced by Pacific/Indian and Atlantic OHT. For all terms a northward c-eq AHT (positive value) is induced by an increase in heat entering the southern hemisphere atmosphere. Change is calculated as the mean c-eq AHT for 2081 through 2100 minus the mean c-eq AHT for 2006 through 2025.

45 Effects of the aerosol clean-up and ice-albedo feedback For all the models, changes in TOA SW radiation induce either no change in c-eq AHT or a southward c-eq AHT (Figure 3.4). This is consistent with what we would expect if aerosol concentrations in the NH reduce in the 21 st century. Aerosols scatter SW radiation, decreasing the SW radiation that reaches the Earth s surface. Figure 3.7 shows the spatial change in incoming SW radiation due to changes in clear-sky scattering. For most of the models the largest increase in SW radiation occurs in the western Pacific off the eastern coast of Asia. This is consistent with what we might expect if aerosol emissions decrease in Asian countries. The incoming SW radiation appears to increase more in the NH than SH due to clear-sky scattering changes. If the concentration of aerosols decreases in the NH, the reflected SW radiation would decrease resulting in a warming and a southward AHT. The change in c-eq AHT induced by surface, cloud and noncloud SW changes using the APRP method is shown in panel B of Figure 3.4. The northward c-eq AHT decreases due to changes in the SW radiation. However, noncloud scattering is not the main driver of the southward c-eq AHT. Surface albedo and cloud changes both induce a southward c-eq AHT and in some cases, have a larger impact on the TOA SW radiation than noncloud scattering. The sign of the c-eq AHT induced by surface albedo changes is consistent with what we would expect to occur as a result of the ice-albedo feedback. We expect the area covered by snow and ice in the NH to decrease more than the SH. This reduces the albedo of the NH and results in an increase in incoming SW radiation. The increase in SW radiation warms the NH relative to the SH, inducing a southward c-eq AHT and northward ITCZ shift.

46 42 Figure 3.7. Change in incoming SW radiation due to clear-sky scattering changes calculated using the APRP method. Change calculated as the mean for 2081 through 2100 minus the mean for 2006 through Models with an asterisk (*) include the indirect effects of aerosols on cloud albedo and lifetime as well as direct aerosol radiative. Models with a plus sign (+) only include indirect effects of aerosols on cloud albedo in addition to direct radiative effects. Models without a symbol only include direct aerosol radiative effects. Figure 3.8 shows the change in downward SW radiation at the TOA due to cloud changes. Models with an asterisk (*) by their name include the indirect effects of aerosols on cloud albedo and lifetime as well as direct aerosol radiative effects (ACCESS1-0, ACCESS1-3, CSIRO-Mk3-6-0, GFDL-CM3, MIROC5, NorESM1-M and NorESM1-ME). Models with a plus sign (+) only include indirect effects of aerosols on cloud albedo in addition to direct radiative effects (GISS-

47 E2-R, GISS-E2-R-CC, IPSL-CM5A-LR, IPSL-CM5A-MR and IPSL-CM5B-LR). The remaining four models only include direct aerosol radiative effects (CCSM4, FGOALS-s2, GFDL-ESM2G). 43 Figure 3.8. Change in incoming SW radiation due to cloud changes calculated using the APRP method. Change calculated as the mean for 2081 through 2100 minus the mean for 2006 through Models with an asterisk (*) include the indirect effects of aerosols on cloud albedo and lifetime as well as direct aerosol radiative. Models with a plus sign (+) only include indirect effects of aerosols on cloud albedo in addition to direct radiative effects. Models without a symbol only include direct aerosol radiative effects. In the tropics changes in SW radiation due to cloud changes could be a result of ITCZ shifts. To determine the effect of SW cloud changes on the ITCZ we only want to look at cloud changes outside of the tropics. Figure 3.9 shows the c-eq AHT attribution due to cloud changes in the tropics, extratropics and clear-sky SW changes.

48 c-eq AHT induced (PW) 44 Cloud changes outside of the tropics tend to induce a southward c-eq AHT. This is consistent with what we would expect if aerosol emissions reduce and clouds are altered due to aerosol indirect effects. However, the models with red labels that include effects of aerosols on cloud albedo and lifetime do not stand out from the rest of the models. The MPI-ESM-MR model does not include aerosol indirect effects, yet it shows cloud changes result in a greater southward c-eq AHT than the NorESM1-M and NorESM1-ME models which do include aerosol indirect effects. Therefore, the southward c-eq AHT and northward ITCZ shift induced by SW cloud changes is likely not solely a result of aerosol clean-up. TOA SW TOA SW from 15 S to 15 N TOA SW excluding 15 S to 15 N TOA SW clear-sky changes TOA SW due to surface albedo changes TOA SW due to noncloud scattering TOA SW due to noncloud absorption Figure 3.9. Change in c-eq AHT induced by SW radiation changes. For all terms a northward c- eq AHT (positive value) is induced by an increase in heat entering the SH atmosphere. Change is calculated as the mean c-eq AHT for 2081 through 2100 minus the mean c-eq AHT for 2006 through 2025.

49 Effects of the longwave radiative feedback For almost all the models, changes in TOA net LW result in a northward AHT, indicating the outgoing LW radiation (OLR) in the NH increases more than the OLR in the SH. This is what we would expect if the LW radiative feedback cools the NH more than the SH. The LW radiative feedback results in increased OLR cooling when surface temperatures increase. The NH surface temperatures are expected to increase more than the SH surface temperatures due to there being more land in the NH, resulting in a greater increase in OLR in the NH than SH. If the OLR in the NH increases more than the OLR in the SH this would induce a northward c-eq AHT and southward ITCZ shift. Changes in the TOA LW in the tropics may be a result of ITCZ shifts. The ITCZ is characterized by deep convection and increase OLR. If the precipitation asymmetry increased and ITCZ shifted north we would expect the OLR in the northern tropics to increase, resulting in a NH cooling. To remove the effects of the tropics we can look at the c-eq AHT induced by tropical LW changes and extratropical LW changes. Figure 3.10 shows the c-eq AHT induced by LW changes in the tropics (15 S to 15 N) and in the extratropics. All models show extratropical LW changes resulting in a northward AHT which is what is expected if the OLR cooling in the NH increases more than the OLR cooling in the SH.

50 c-eq AHT induced (PW) 46 TOA LW TOA LW from 15 S to 15 N TOA LW excluding 15 S to 15 N Figure Change in c-eq AHT induced by LW radiation changes. For all terms a northward c- eq AHT (positive value) is induced by an increase in heat entering the SH atmosphere. Change is calculated as the mean c-eq AHT for 2081 through 2100 minus the mean c-eq AHT for 2006 through Result of competing forcings As Figure 3.1 shows, an increase in northward c-eq AHT results in a southward ITCZ shift. When more heat enters one hemisphere, the Hadley cells will shift towards the hemisphere that is heated more to transport the heat across the equator. The ITCZ is located at the convergence of the northern and southern Hadley cells, so when the Hadley cells shift across the equator to transport heat northward (southward) the ITCZ will shift south (north). Figure 3.11 shows surface albedo (ice-albedo feedback) and noncloud scattering (aerosol cleanup) changes result in a southward c-eq AHT. The southward c-eq AHT induced by changes in TOA SW radiation is counteracted by OHT in the Atlantic and the LW cooling in the NH. Comparing the multi-model mean of the c-eq AHT induced by aerosol direct effects, ice-albedo feedback, AMOC slowdown and LW radiative feedbacks, the forcings that result in a northward

51 cross-eq AHT induced(pw) 47 c-eq AHT appear to be larger than the forcings that result in a southward c-eq AHT. However, the multi-model mean c-eq AHT is approximately zero indicating aerosols and the ice-albedo feedback are not the only forcings that result in a southward c-eq AHT. Aerosol direct effects Icealbedo feedback AMOC slowdow n Ocean heat content Longwave feedback Δ c-eq AHT TOA SW noncloud scattering TOA SW due to surface albedo changes -AMOC OHT Ocean TOA LW heat excluding content 15 S to 15 N Remainder Figure C-eq AHT induced by aerosol Change in precipitation asymmetry versus change in c-eq AHT induced by clear-sky scattering, albedo, northward AMOC heat flux, ocean heat content changes and longwave changes. To determine whether the forcings shown in Figure 3.11 can explain the ITCZ shift we plot the c-eq AHT induced by each forcing against the total change in the precipitation asymmetry which represents the ITCZ shift (Figure 3.12). An interesting feature of the plot is the positive correlation between changes in northward c-eq AHT induced by the AMOC slowdown and LW cooling. This contradicts our understanding of how the Hadley cells transport heat across the equator and results in an ITCZ shift when a hemispheric heating asymmetry exists. The positive correlation for the AMOC slowdown and LW cooling could indicate the AMOC and LW cooling

52 Δ Pr 0-20N - Δ Pr 0-20S (mm/day) 48 are responding to other forcings that result in the NH warming. In other words, when the heat flux into the NH increases due to clear-sky scattering, surface albedo, etc. changes the AMOC and LW cooling adjust to counteract the NH warming. A Δ TOA clear-sky SW scattering, r = B Δ surface albedo r = C Δ OHT in the Atlantic r = 0.48 D Δ TOA incoming LW, excluding 15S to 15N r = 0.34 cross-equatorial atmosphere heat transport induced (PW) Figure Change in precipitation asymmetry versus change in c-eq AHT induced by clear-sky scattering, albedo, upward ocean heat flux and northward AMOC heat flux changes. The low correlation coefficient for surface albedo changes, which is representative of the icealbedo feedback, versus the change in precipitation asymmetry indicates the ice-albedo feedback does not fully explain the ITCZ shift or the spread in the models. The forcing with both the highest correlation coefficient and correct sign on the slope in Figure 3.12 is the direct effects of aerosol clean-up. However, the correlation coefficient for aerosol direct effects versus changes in the precipitation asymmetry is lower than the correlation coefficient for the total c-eq AHT versus

53 change in precipitation asymmetry, indicating aerosol direct effects alone cannot explain the change in c-eq AHT and ITCZ shifts EXTRATROPICAL FORCING AND FEEDBACKS We separated out the impacts of extratropical feedbacks and forcings from tropical feedbacks and forcings by looking at the c-eq AHT induced by radiative changes from 90 S to 15 S and from 15 N to 90 N. Equation 3.2 was used to calculate the c-eq AHT induced by extratropical radiative changes. π 12 2π π 2 0 π 2 π 12 AHT Extratrop = 1 [ Q 2 Aa 2 cosλ dφ dλ + Q A a 2 cosλ dφ dλ] (3.2) The c-eq AHT due to extratropical TOA LW, TOA SW, surface albedo, cloud, noncloud scattering and noncloud absorption changes are shown in Figure π 0

54 Excludes 15 S to 15 N 50 A B C D E F G H I J Includes 15 S to 15 N Figure Change in c-eq AHT induced by radiative changes calculated using hemisphere totals versus change in c-eq AHT calculated with excluding 15 S to 15 N. A positive value indicates the radiative change results in a greater increase in heat entering the SH than NH inducing a northward AHT. The x-axis in Figure 3.13 lists the change in c-eq AHT calculated by integrating over the entire hemisphere (Equation 2.2). The heat flux that induced the c-eq AHT is listed in the title. The y-axis lists the change in c-eq AHT calculated by integrating from 90 to 15 (Equation 3.2). Each plot includes a 1-to-1 line. If a model falls on the line, the tropical feedbacks and forcings do not impact the hemisphere heating asymmetry. For example, panel D shows the change in c-eq AHT

55 51 induced by changes in the surface albedo. All the models fall on the 1-to-1 line indicating changes in the surface albedo occur almost entirely in the extratropics. For terms that do not fall on the 1-to-1 line, the change in the radiative forcing in the tropics either leads to an enhancement of the hemisphere heating asymmetry (positive feedback) or counteracts the asymmetry (negative feedback). For example, panel H shows a comparison of the c-eq AHT induced by changes in SW noncloud absorption. For all the models the change in c-eq AHT calculated by integrating over the entire hemisphere results in less of a c-eq heat transport than when the tropics are excluded and in some cases, excluding the tropics changes the sign of the heat flux induced. This indicates changes in the SW noncloud absorption in the tropics dampen the heating asymmetry and lead to a negative feedback. Panel J list the sign associated with each section of the figure. The models included in this analysis show changes in the tropical net TOA radiation, SW clouds and SW noncloud scattering lead to both an enhancement and dampening of the c-eq AHT. Changes in tropical LW radiation lead to a dampening of the c-eq AHT for most models. Tropical SW noncloud (absorption) enhance the c-eq AHT for all models. The c-eq AHT induced by surface flux changes in the extratropics and the entire hemisphere are shown in Panel I. For 11 of the 16 models, including the tropical surface flux changes results in a greater c-eq AHT than when the tropics are excluded. For the remaining six models the change in surface flux in the tropics has little impact on the c-eq AHT. 3.4 OCEAN-ATMOSPHERE C-EQ HEAT TRANSPORT SPLIT As demonstrated by Figure 3.4, the RCP8.5 models show a range of changes in c-eq AHT and OHT and the ratio between AHT and OHT does not seem to be consistent across the models. Table

56 3.1 shows the c-eq AHT and OHT in PW along with the fraction of total c-eq heat transport due to the atmosphere and the ocean. The table is sorted based on the fourth column, which shows the fraction of total northward heat transport by the atmosphere. The fraction of heat transport by the atmosphere and ocean varies significantly. The fraction of total c-eq heat transport by the atmosphere ranges from -21% to 140%. When the fraction is greater than 100% the ocean and atmosphere are transporting heat in opposite directions. The fraction of total c-eq heat transported by the ocean ranges from -261% to -5%. For all the models the northward OHT decreases. The northward AHT decreases for six of the models that OHT data is available and increases for the remaining six. Table 3.1. Change in c-eq AHT and c-eq OHT. A positive Δc-eq AHT indicates the increase in heat entering the southern hemisphere is greater than the increase in heat entering the northern hemisphere, inducing a northward c-eq AHT. A negative Δc-eq OHT induces a northward AHT. AHT fraction of northward transport OHT fraction of northward transport Model Δc-eq AHT (PW) Δc-eq OHT (hfy, PW) GFDL-CM % -79% GFDL-ESM2G % -65% ACCESS % -80% IPSL-CM5B-LR % -29% NorESM1-M % -67% ACCESS % -93% NorESM1-ME % -145% GISS-E2-R % -168% IPSL-CM5A-LR % -127% GISS-E2-R-CC % -261% IPSL-CM5A-MR % -5% MPI-ESM-MR % -240% 52 Figure 3.14 shows the change in zonal mean column integrated heat flux into the atmosphere (downward TOA + upward surface flux) for the models listed in Table 3.1. Figure 3.14 shows the change in zonally integrated northward OHT for all basins and Figure 3.16 shows the change in northward OHT in the Atlantic basin alone. The models are sorted from the smallest to largest c-

57 53 eq AHT fraction in both plots (same as Table 3.1), with the northward c-eq AHT decreasing for the models shown in red and increasing for the models shown in blue. The three models with the largest decrease in northward c-eq AHT also appear to have the largest increase in zonal mean column integrated heat flux in the northern tropics. The models with the largest increase in northward c-eq AHT and decrease in northward c-eq OHT show the largest increase in zonal mean column integrated heat flux around 42 S.

58 54 Δ zonal mean column integrated heat flux (W m -2 ) Sorted from least to most AHT fraction of total northward heat transport - = northward AHT increases; -- = northward AHT decreases Figure Change in zonal mean column integrated heat flux (TOA down + surface flux up) into the atmosphere. Change calculated as the mean for 2081 through 2100 minus the mean for 2006 through 2025.

59 55 Δ zonally integrated northward ocean heat transport (PW) Sorted from least to most AHT fraction of total northward heat transport - = northward AHT increases; -- = northward AHT decreases Figure Change in zonally integrated northward OHT. Change calculated as the mean for 2081 through 2100 minus the mean for 2006 through 2025.

60 Figure Change in zonally integrated northward OHT in the Atlantic basin. Change calculated as the mean for 2081 through 2100 minus the mean for 2006 through

61 57 Figure Northward AHT induced by changes in the upward ocean (surface) flux. Table 3.2 shows the c-eq AHT induced by changes in northward OHT and differences in the SH and NH ocean heat uptake. Comparing the c-eq AHT induced by the northward OHT plus storage changes and the total c-eq AHT gives a smaller range of ocean-atmosphere c-eq compensation than when the c-eq OHT without storage changes and c-eq AHT are compared. Changes in ocean storage counteract the decrease in northward OHT for many of the models. Table 3.3 shows the upward surface flux breakdown into northward OHT and storage changes.

62 Table 3.2. Total change in c-eq AHT and the change in AHT induced by changes in the upward surface (ocean) fluxes. A positive Δ c-eq AHT due to upward surface fluxes value indicates the change in upward heat flux in the SH is greater than. The upward surface heat flux includes ocean storage and northward OHT effects. Δc-eq AHT due to upward surface fluxes (PW) AHT fraction based on surface flux OHT OHT fraction based on surface flux Δc-eq Model AHT (PW) ACCESS % 37.70% ACCESS % 34.20% GFDL-CM % 37.80% GFDL-ESM2G % 16.90% GISS-E2-R % -8.80% GISS-E2-R-CC % -6.10% IPSL-CM5A-LR % 56.80% IPSL-CM5A-MR % 76.20% IPSL-CM5B-LR % 82.50% MPI-ESM-MR % 34.40% NorESM1-M % 48.40% NorESM1-ME % 51.80% 58 Table 3.3. C-eq OHT due to changes in northward heat transport and due to hemispheric storage asymmetries. Model Δc-eq AHT due to surface flux changes (PW) Δc-eq OHT due to hfy (PW) Δc-eq AHT due to ocean storage (PW) GFDL-CM ACCESS ACCESS MPI-ESM-MR GFDL-ESM2G GISS-E2-R-CC GISS-E2-R NorESM1-ME NorESM1-M IPSL-CM5A-LR IPSL-CM5B-LR IPSL-CM5A-MR

63 59 Kay et al. (2016) found most of the southward OHT induced by brightening the southern hemisphere low clouds occurred in the Pacific Ocean. They found the easterly wind stress over the SH subtropics increased, especially in the Pacific basin. The increased wind stress strengthened the poleward Ekman transport, leading to increased poleward heat transport. If this same mechanism appears in the CMIP5 RCP8.5 models we may expect a correlation to appear between changes in heat transport in the Pacific Ocean and the change in c-eq AHT. Figure 3.6 shows the change in c-eq OHT by basin. Figure 3.18 shows the change in c-eq OHT by basin plotted against the total change in c-eq AHT. Table 3.4 shows the c-eq OHT in the Atlantic and Pacific basins along with the total c-eq AHT. The northward heat transport in the Atlantic basin either decreases or shows no change in all 12 models. This is consistent with what we would expect if the AMOC slowed with warming. However, the plot of change in northward OHT in the Atlantic basin versus change in c-eq AHT shows there is a low correlation between the two fluxes (Figure 3.18). The change in northward heat transport in the Pacific basin varies the most of the three basins. Of the three basins, changes in c-eq OHT in the Pacific basin shows the highest correlation coefficient (0.57) with the change in c-eq AHT, indicating heat transport in the Pacific basin may counteract or drive the c-eq AHT.

64 60 A B Figure Change in c-eq AHT versus change in c-eq OHT by basin. Values over land are omitted. Change calculated as the mean for 2081 through 2100 minus the mean for 2006 through Table 3.4. Change in c-eq OHT by basin and change in c-eq AHT. Model Δc-eq AHT (PW) Δc-eq OHT (hfy, PW) Δc-eq Pacific Heat Transport (PW) Δc-eq Atlantic Heat Transport (PW) GFDL-CM GFDL-ESM2G ACCESS IPSL-CM5B-LR NorESM1-M ACCESS NorESM1-ME GISS-E2-R IPSL-CM5A-LR GISS-E2-R-CC IPSL-CM5A-MR MPI-ESM-MR

65 61 Chapter 4. GFDL-MOM EXPERIMENTS Two separate studies by Kay and Hawcroft in 2016 found the same split between the fraction of the increase in northward OHT and AHT when the Southern Ocean albedo was increased. Kay and Hawcroft used different methods to brighten the Southern Ocean in two different fully coupled GCMs. Both found the ocean transported 80% of the induced c-eq heat transport southward and the atmosphere transported 20%. A third study published by Haywood et al. (2016) that used the same model as Hawcroft et al. resulted in a different ocean-atmosphere c-eq heat transport split when the entire southern hemisphere was brightened in the HadGEM2-ES model. Haywood found 34% of the c-eq heat transport was transported by the ocean and 66% was transported by the atmosphere. Based on the results of these studies, we hypothesized the compensation between the ocean and atmosphere in response to a thermal forcing may be dependent on the location in which a thermal forcing is added. Our analysis of the RCP8.5 GCMs showed the c-eq heat transport split between the ocean and atmosphere varied significantly across the models. In the CMIP5 RCP8.5 simulations the spatial structure of the thermal forcings (cloud changes, water vapor concentrations, aerosol concentrations, etc.) is not consistent. Thus, we cannot determine the impact of the spatial pattern of the energetic forcings on the c-eq ocean and atmosphere heat transport split. To determine the dependence of the c-eq heat transport on the location of the forcing we would need to keep all variables constant except for the location of the forcing. Since this has not been done with the CMIP archive, we ran a series of simulations in a fully coupled GCM.

66 MODELS AND SIMULATIONS To evaluate the impact of the location of a thermal forcing on the c-eq heat transport we used the Geophysical Fluid Dynamics Laboratory Climate Model version 2 (GFDL CM2) with Modular Ocean Model version 5 (MOM5) at coarse resolution (CM2Mc). The atmosphere has a resolution of 3 x 3.75 with 24 vertical levels. The ocean has a resolution of 2.25 x 3 with 28 vertical levels. The ocean and atmosphere composition, and temperature were set to pre-industrial levels. The model was fully coupled with full geography. The control simulation was run for 1300 years but only the last 300 years were used for comparison with the simulations in which a heat flux was applied. The heat flux was applied in the form of a q-flux at the ocean-atmosphere interface, meaning a heat source was added. Heat was not removed from the atmosphere and added to the ocean or vice versa. Five simulations with different q-fluxes applied were run for 100 years. The location and magnitude of the heat fluxes added are shown in Table 4.1. Table 4.1. GFDL-MOM5 simulations with the total heat added and region in which the heat was added. Experiment Name q-flux applied (W m -2 ) Ocean basin heat was added to q-flux latitude boundaries Total heat added NPac10 10 Pacific 30 N to 50 N PW NH NAtl Atlantic 30 N to 50 N PW NH TPac Pacific 10 N to 30 N PW NH SOcn Southern Ocean 75 S to 45 S PW SH SOcn10 10 Southern Ocean 75 S to 45 S PW SH

67 63 To keep the model energetics in balance, heat was removed uniformly from the remainder of the ocean-atmosphere boundary. For each simulation setup, three separate ensemble members were run for 100 years starting from the control run year 1000, 1005, and Figure 4.1 shows the locations in which q-fluxes were added to the model. TPac8 Q-flux = W m to 30 N NPac10 Q-flux = 10 W m to 50 N NAtl11 Q-flux = W m to 50 N SOcn2 Q-flux = W m to 45 S SOcn10 Q-flux = 10 W m to 45 S Figure 4.1. Heat fluxes applied at the ocean surface in the GFDL-MOM5 simulations. For each simulation, except for SOcn10, the total PW added was constant at PW. For the SOcn10 run a total forcing of 0.6 PW was added. Three ensembles were run with each setup for 100 years. For each experiment, we evaluated the change in surface flux, TOA LW, TOA SW, poleward MSE transport and poleward OHT. Comparing the change in c-eq AHT and OHT for the different ensemble members for each experiment we found the results of each experiment did

68 not vary significantly across the three ensemble runs. Therefore, only three ensemble members were run for each experiment C-EQ ATMOSPHERE AND OCEAN HEAT FLUX RESPONSE Northward MSE transport for each GFDL-MOM5 simulation was calculated using Equation 2.1. The c-eq AHT was calculated for each run using Equation 2.2. The northward c-eq OHT was calculated by summing the meridional heat transport due to advection, subgrid scale flux components from the Gent and McWilliams (1990) scheme, subgrid scale flux components arising from the mixed-layer submesoscale parameterization scheme of Fox-Kemper et al. (2008) and meridional diffusion. In Figure 4.2 Panels A and B show the experiment minus control 100-year mean zonally integrated poleward MSE and OHT. In the figures the values at the equator indicate the change in c-eq AHT and OHT. The difference between the experiment and control ensemble mean c-eq AHT and OHT (exp ctl) for each experiment are also shown in Table 4.2. The similarity in the change in c-eq AHT and OHT for each ensemble member for each experiment indicate three ensembles is sufficient.

69 PW PW 65 experiment - control northward atmosphere heat transport northward ocean heat transport Latitude Figure 4.2. Change in northward atmosphere heat transport and ocean heat transport (PW, experiment control). In plots of the zonally integrated poleward OHT (panel B) as we would expect when heat is added to the SH the ocean transports heat northward and when heat is added to the NH the ocean transports heat southward. However, this is not always the case for the atmosphere. When 10 W m -2 was added to the Southern Ocean, as expected, the atmosphere transports heat northward. However, when 2.36 W m -2 is added to the Southern Ocean the atmosphere transports heat southward. The differences between these two experiments indicates the c-eq AHT and OHT split may be dependent on the magnitude of the heating.

70 66 Comparing the NH experiments, when heat was added to the Pacific Ocean the atmosphere transports almost twice as much heat southward than when heat is added to the Atlantic. When heat was added to the tropical Pacific (TPac8) the c-eq AHT is similar to the experiment in which heat was added to the north Pacific (NPac10). However, the c-eq OHT differs for these experiments. The c-eq OHT for the TPac8 and Natl11 experiments are very similar. When heat was added in the North Pacific (NPac10) the ocean transported more heat southward. The difference in the southward c-eq OHT between the Pacific experiments (NPac10 and TPac8) and c-eq AHT in the NH extratropic experiments (NPac10 and NAtl11) could be due to the cloud changes which amplify the heating added in the North Pacific (Figure 4.8). Figure 4.2 also appears to indicate the atmosphere will be responsible for a larger portion of the c-eq total heat transport if heat is added to the NH. Comparing the change in c-eq AHT for the five experiments, the atmosphere transports a larger fraction of the total southward heat transport in response to heating in the NPac10, Natl11 and TPac8 experiments than when heat is added to the Southern Ocean in the SOcn2 and SOcn10 experiments. This supports our hypothesis that the c-eq heat transport split between the atmosphere and ocean is dependent on the location of the heating. Figure 4.3 shows the ensemble mean change in precipitation for each experiment and Figure 4.4 shows the change in the precipitation asymmetry calculated using Equation 3.1 versus the change in c-eq AHT and c-eq OHT for each experiment. When heat was added to the NH the precipitation in the northern tropics generally increased and the precipitation in the southern tropics decreased, indicating a northward ITCZ shift. When heat was added to the SH, the precipitation on the equator increased but it is not clear if there was a northward or southward ITCZ shift. As was found with the CMIP5 RCP8.5 GCMs, the change in precipitation asymmetry is anticorrelated

71 67 to the c-eq AHT. The change in precipitation asymmetry also appears to be anticorrelated with the change in c-eq OHT, which is expected if the ocean transports a fraction of the c-eq heat transport in response to hemispheric heating asymmetries. The most significant ITCZ shifts occurred when heat was added to the NH with little to no change in the precipitation asymmetry when 10 W m -2 was added to the Southern Ocean. NPac10 Q-flux = 10 W m to 50 N Natl11 Q-flux = W m to 50 N TPac8 Q-flux = W m to 30 N mm day -1 SOcn2 Q-flux = W m to 45 S SOcn10 Q-flux = 10 W m to 45 S Figure 4.3. Experiment minus control precipitation (mm/day).

72 Δ P 0-20N - P 0-20S (mm/day) 68 experiment - control r = -0.97; p = 5.28e-9 r = -0.88; p = 1.53e-5 Δ northward AHT (PW) Δ northward OHT (PW) Figure 4.4. Experiment minus control precipitation asymmetry versus northward c-eq AHT and northward c-eq OHT. Table 4.2 provides a summary of the heating added to each run and the change in c-eq OHT and AHT for each experiment. The values shown for the GFDL-MOM5 runs in Table 4.2 are ensemble means. The last three rows show the results of experiments run in the CESM coarse resolution model by Rachel White provided through personal correspondence. White ran experiments using the CAM4-POP2 version of the CESM model at low resolution (f45-g37: ~4-degree atmosphere, ~3-degree ocean) as a comparison to the Haywood (2016), Hawcroft (2016) and Kay (2016) studies. In the simulations, White imposed ocean-surface heating of 10 W m -2 in three locations. The area of the heating was changed but the q-flux applied was held constant. Therefore, the total heat added in PW changed for each run. The heating was added by increasing the radiation that reached the surface of the ocean. The first simulation added the forcing between 75 S and the equator over all longitudes (10SH) in a similar setup as Haywood et

73 69 al (2016); the second simulation added heating between 45 S and 75 S over all longitudes (10SO) in a similar setup as Hawcroft et al. (2016) and Kay et al. (2016); and the third between 25 N and 55 N over the Atlantic region only (10NA). The third simulation was run to evaluate the impacts of aerosol clean-up off the northeast coast of North America. For the first two simulations (10SH and 10SO) only one ensemble member was run for 100 years. For the 10NA experiments, an ensemble of 4 members was created using different initial conditions and run for 100 years. Comparing the results of White s CESM coarse resolution simulations in a fully coupled model to our GFDL-MOM5 results (Table 4.2) a few similarities and differences arise. First, in both the GFDL and CESM 10 W m -2 Southern Ocean simulations the c-eq ocean-atmosphere heat transport split is approximately 80:20, resembling the results of Hawcroft et al. (2016) and Kay et al. (2016). The response to heating over just the Southern Ocean region (10SO) is a c-eq heat transport largely (80%) in the ocean. In addition, White found a 50:50 split between the atmosphere and ocean c-eq heat transport was added to the entire SO, which is consistent with the results of Hawcroft et al. (2016). The consistencies between these results and the Hawcroft, Haywood and Kay studies indicates the results may not be dependent on the model, method for heating/cooling or whether a heating or cooling force is added. When heat was added to the northern extratropics in the GFDL NPac10, Natl11 and CESM 10NA the ocean-atmosphere c-eq heat transport split is approximately 40:60 with most the heat transport occurring in the atmosphere. The results of the northern extratropic simulations are also within 5% of each other despite the differences in the total heat added, basin in which the heat was added and the method used to add the heat flux. The heat was added to relatively similar latitude bins, with the NPac10 simulation adding heat between 30 N and 50 N and the 10NA simulation adding heat over a slightly larger latitude band from 25 N and 55 N.

74 Comparing the GFDL SOcn2 simulation run to the GFDL SOcn10 and CESM 10SO simulations we find the atmosphere c-eq heat transport varies significantly. In the simulations heat was added to the same region from 75 S to 45 S in the Southern Ocean. In the simulation in which less heat was added (SOcn2) the AHT is in the opposite direction as we would expect. This result indicates the magnitude of the forcing may impact the resulting ocean-atmosphere c-eq heat transport. Table 4.2. Change in c-eq AHT and OHT for GFDL-MOM5 experiments and select CESM coupled experiments. 70 GFDL simulations Total heat added (PW) Δ c-eq OHT (PW) Δ c-eq AHT (PW) ocean % atm % NPac10: 10 W m -2 Pacific, 30 N to 50 N NAtl11: W m -2 Atlantic, 30 N to 50 N TPac8: 8.37W m -2 Pacific, 10 N to 30 N SOcn2: W m -2 Southern Ocean, 75 S to 45 S SOcn2: 10 W m -2 Southern Ocean, 75 S to 45 S % 62.1% % 62.3% % 72.4% % -114% % 19.1% CESM simulations (Rachel White) 10NA: 10 W m -2 Atlantic, 25 N to 55 N 10SH: 10 W m -2 Southern Ocean, 75 S to 0 S 10SO: 10 W m -2 Southern Ocean, 75 S to 45 S % % % 57% 50% 16%

75 SURFACE AND TOA FLUX RESPONSE To understand the atmosphere response further we can break down the forcing on the atmosphere into the surface and TOA heat fluxes. The upward surface flux was calculated by summing the following terms at the ocean surface. F S = R S + R L + F SH + L V E + L M Sn + F PME + F River (4.1) Where F S is the net surface flux, R S is the net upward SW flux, R L is the net upward LW flux, F SH is the net upward sensible heat flux, L V is the latent heat of evaporation (2.5e6 J kg -1 ), E is the evaporation rate (kg m -3 m s -1 ), L M is the latent heat of fusion/melting (3.35e5 J kg -1 ), Sn is the snowfall rate (kg m -3 m s -1 ), F PME is the net heat flux from the flux of freshwater across the ocean surface and F River is the heating associated with the introduction of river runoff over the uppermost model grid cells. The difference between the experiment and control runs (exp - ctl) upward surface flux are shown in Figure 4.5. The heating added at the ocean surface appears in the upward ocean surface heat flux for four of the five simulations. For the SOcn2 simulation the added q-flux does not appear in the surface flux deviation plots. The q-flux applied in the SOcn2 runs may either be too small to appear in the exp ctl or dynamics in the Southern Ocean may mix the added heat quickly preventing it from appearing in the surface flux plots.

76 72 NPac10 Q-flux = 10 W m to 50 N NAtl11 Q-flux = W m to 50 N TPac8 Q-flux = W m to 30 N SOcn2 Q-flux = W m to 45 S SOcn10 Q-flux = 10 W m to 45 S Figure 4.5. Ensemble mean experiment minus control surface flux response for GFDL- MOM5 experiments.

77 73 NPac10 Q-flux = 10 W m to 50 N NAtl11 Q-flux = W m to 50 N TPac8 Q-flux = W m to 30 N SOcn2 Q-flux = W m to 45 S SOcn10 Q-flux = 10 W m to 45 S Figure 4.6. Ensemble mean experiment minus control sea surface temperature response for GFDL-MOM5 experiments. The experiment minus control zonal mean column integrated heat flux (incoming TOA + upward surface flux), surface flux and net TOA (LW + SW) incoming radiative flux are shown in Figure 4.7. The added heat is visible in plots of the zonal mean column integrated heat flux and surface flux. The runs in which heat was added show an increased column integrated heat flux and upward surface flux in the region in which the heat was added. Plots of the zonal mean TOA net (LW + SW) radiation were not as straight forward. For the NPac10 and TPac8 runs there appears to be an increase in the TOA radiation where the heating was added. For the TPac8 run, the change

78 74 in TOA radiation appears to indicate a shift of the Intertropical Convergence Zone (ITCZ) leading to a positive and negative TOA radiation anomaly on either side of the equator. Looking at the change in TOA radiation and low cloud fraction for the NPac10 run (Figure 4.8), where heating is added the low cloud fraction appears to decrease leading to an increase in downward TOA radiation. This could indicate heating added in the high latitudes is amplified due to positive cloud feedbacks.

79 Figure 4.7. Experiment minus control zonal mean column integrated energy flux into the atmosphere, upward surface flux and net downward TOA (LW + SW) radiation (W m -2 ). 75

80 76 Figure 4.8. Change in net downward TOA (LW + SW) radiation and low cloud fraction (experiment control). Values shown are mean values for the 100 period the experiments were run. Figure 4.9 shows the difference between the experiment and control run northward c-eq OHT by basin for each simulation and ensemble member. When heat was added to the Southern Ocean from 75 to 45 S almost all the northward OHT occurred in the Pacific basin. Kay et al. (2016)

81 PW 77 also found most the northward c-eq heat transport when the albedo of the clouds over the Southern Ocean was altered between 75 to 45 S occurred in the Pacific due to wind driven gyre changes. For our NH experiments we found the c-eq OHT did not occur exclusively or even mostly in the Pacific Ocean. When heat was added to the NH tropics, the c-eq OHT was almost evenly split between the Pacific and Atlantic basins. When heat was added to the NH extratropics the largest change in OHT occurred in the Atlantic. experiment control northward c-eq OHT Total Pacific & Indian Atlantic Figure 4.9. Experiment minus control northward c-eq OHT (PW) split up by ocean basin. Kay hypothesized the heat transport in the Pacific was due to an increase in the easterly winds in the tropics which enhanced the poleward Ekman transport. In the NH tropics the Ekman mass and heat transport is 90 to the right of the wind stress. Since the winds come from the east in the tropics the result is heat being transported away from the equator towards the pole. In the SH tropics, the Ekman mass and heat transport is 90 to the left resulting in a southward (poleward) heat transport. Looking at the change in the wind stress compared to the control run when 10 W m -2 was added to the Southern Ocean the easterly winds did increase. However, the easterlies

82 78 increased both north and south of the equator, implying an increase in poleward heat transport but not necessarily northward heat transport. In the SOcn2 experiment there was little to no change in the easterlies in the tropics. In the NH experiments, the easterlies north of the equator decreased implying a decrease in northward Ekman heat transport and the easterlies south of the tropics increased implying an increase in southward Ekman heat transport. The decrease in the easterlies in the northern tropics and increase in the easterlies in the southern tropics imply a net southward heat transport in the Pacific when heat was added to the NH. However, in the simulations in which heat was added to the NH most the c-eq OHT occurred in the Atlantic and not the Pacific. NPac10 Q-flux = 10 W m to 50 N NAtl11 Q-flux = W m to 50 N TPac8 Q-flux = W m to 30 N N m -2 SOcn2 Q-flux = W m to 45 S SOcn10 Q-flux = 10 W m to 45 S Figure Change in zonal wind stress (experiment control). Values shown are mean values for the 100 period the experiments were run.