Tropical and Extratropical Air-Sea Interactions: Modes of Climate Variations

Tropical and Extratropical Air-Sea Interactions: Modes of Climate Variations provides a thorough introduction to global atmospheric and oceanic processes, as well as tropical, subtropical and mid-latitude ocean-atmosphere interactions. Written by leading experts in the field, each chapter is dedicated to a specific topic of air-sea interactions (such as ENSO, IOD, Atlantic Nino, ENSO Modoki, and newly discovered coastal Niños/Niñas) and their teleconnections. As the first book to cover all topics of tropical and extra-tropical air-sea interactions and new modes of climate variations, this book is an excellent resource for researchers and students of ocean, atmospheric and climate sciences.

• Presents case studies on the ocean-atmosphere phenomena, including El Nino Southern Oscillation (ENSO), Indian Ocean Dipole and different Nino/Nina phenomena
• Provides a clear description of air-sea relationships across the world’s ocean with an analysis of air-sea relations in different time scales and a focus on climate change
• Includes prospects for air-sea interaction research, thus benefiting young researchers and students

• Language: English
• Publisher: Elsevier Science
• Release date: Nov 18, 2020
• ISBN: 9780128181577

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http://www.jamstec.go.jp/res/ress/yamagata/

Preface

Climate variability and change have huge impacts on the global socioeconomic conditions. From agriculture to human health, the human society is now facing enormous challenges owing to the extreme climate events that frequently appear with higher intensities under the stress of global warming. It is not the first time that the planet has seen such changes. The world has experienced the vagaries of climate extremes and climate change in all its pasts. The only difference now, unlike our ancestors, is that we are in an opportune time when the climate science has advanced rapidly and the scope of its scientific exploration has increased manifold in the past few decades. Routine weather observations including satellite observations and advancement in telecommunication made it easier for the development of effective weather prediction systems. Those together with the progresses in ocean observations have also helped us to monitor and understand modes of climate variations like the El Niño/Southern Oscillation (ENSO). We have also developed better insights on the behavior of mean climate system and processes that are helping to maintain the mean ocean and atmospheric heat budget, global circulations, etc. In the meantime, advances in computational sciences have helped us to develop state-of-the-art numerical models and reliable climate prediction systems. The present generations of global climate models are able to reliably predict climate variations, especially the tropical climate variations like ENSO and Indian Ocean Dipole, several seasons ahead with skills not far behind that of the weather forecasts that are done with a lead time of a few days.

Recent studies have also helped us in discovering new modes of climate variations in subtropical and coastal regions as mentioned in the foreword of Prof. Toshio Yamagata. Those are shown to be extremely important not only for the climate but also for the marine and terrestrial ecosystems of those regions. An attempt is made in this book to review the present status of all those research studies. Links to available resources are also provided at the end of the book for further research in these areas. While we have developed a lot of understanding on the air–sea interactions of tropical and subtropical climate phenomena, the research in mid-latitude air–sea interactions is not that advanced. Nevertheless, we have tried to bring one such topic for the discussions in the book to provide a flavor of what is happening at this frontier. I hope the studies made by the leading experts in those areas of climate research will help us to establish a base for understanding and predicting the present climate. A better understanding of the present climate system will also help us to reduce model biases and associated errors in the projections of future climate.

Swadhin Kumar Behera

Yokohama, May 19, 2020

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Introduction to atmosphere and ocean variability and air–sea interactions

Swadhin Kumar Behera,    Application Laboratory, Research Institute for Value-Added-Information Generation, Japan Agency for Marine-Earth Science and Technology, Yokohama, Japan

Abstract

Atmosphere and oceans are the major components of the Earth’s climate system and play important roles in determining the intra-seasonal to inter-decadal climate variations discussed in this book. This introductory chapter broadly discusses the internal physical and dynamical processes in both systems in addition to the air–sea interactions that form the basis for basin-scale climate variations such as the El Niño/Southern Oscillation and the Indian Ocean Dipole discussed in the subsequent chapters. The variations in coastal processes are also important for the regional marine ecosystems and regional climate variations. These are also briefly introduced here and covered in some of the chapters of this book.

Keywords

Air–sea interaction; teleconnection; Pacific; Indian Ocean; tropics; extratropics

1.1 Introduction

Earth’s climate system is an interactive system consisting of five major components: the atmosphere, the hydrosphere, the cryosphere, the land surface, and the biosphere. The interaction is very complex as is the internal processes within each of those five components. Atmosphere is one of the major components of the system that has been observed for centuries even before the inventions of weather instruments. In fact, the science of it, known as the meteorology, said to come from the observations of meteors in evening skies. One of the early documentations of the atmospheric science was that of the Aristotle’s treatise Meteoroligca said to be written around 340 BCE. It was notably one of the early documents that brought the subject of meteorology closer to the scientific discussions. Actual scientific observations and analysis of the atmospheric elements started much later in time, that is, after the inventions of Galileo’s air thermometer and Torricelli’s barometer in 17th century. The observations taken with those early instruments helped to understand the vertical distribution of temperature and pressure in the atmosphere, for example, the observation of the temperature and pressure decrease with height (in the troposphere, the lower atmosphere). Besides, those instruments also helped in picking up early signs of local day-to-day atmospheric variations such as dreary-weather or fair-weather conditions associated with lower and higher than normal barometric pressure, respectively. The observation of weather phenomena progressed rapidly after launching of the World Weather Watch program led by the World Meteorological Organization, which brought the satellite observations besides the many other atmospheric parameters observed in weather stations. At the same time, advancement in computational field helped in the development of numerical modeling for weather predictions using advanced computers.

Climate, the topic of the discussions in the present book, is a compilation of the weather states over a place for quite long period of time, for example, over a period of 30 years. Since most regions over Earth experience changes from season to season (Fig. 1–1), an average must be taken, for instance, over January of many different years to observe a climatological condition for January and so on. Also, in addition to these climatologies of, for example, rainfall and temperature averages for a particular month or even a day (the 30-year average of daily data), the climate of a region could also describe the probability of extreme events, such as a major rainfall event occurring in July to August in India, the range of variations of temperature that typically occur during January to March in Tokyo, or the number of hurricanes that typically hit the American coasts in the hurricane season.

Figure 1–1 Schematic representation of the annual cycle of Earth’s rotation around the Sun causing the seasons. The seasons mentioned in the diagram refer to Northern Hemisphere ones.

Classical climatology provides classifications and descriptions of the various climate regimes found in different parts of the world. For example, the Mediterranean climate that is a very pleasant climate with warm and dry summers and cool but not so wet winters as is seen in some of the regions (though not limited to) around the Mediterranean Sea from which it takes its name. Similarly, there are extreme climate regimes like the Desert climate, which refers to very dry climate, and the Tropical climate (discussed in length in this book) that refers to extremely hot and wet conditions. So, climate varies from place to place, depending on latitude, distance to the sea, vegetation, presence or absence of mountains, or other geographical factors. Climate varies also in time, within a season, season to season, year to year, decade to decade, or on much longer timescales like the Ice age and the global warming (referred to as climate change) that we are experiencing now. The natural climate change that Earth has experienced from Ice age to global warming has taken a new turn now since human activities are inducing such changes, which is referred to as anthropogenic climate change. However, all the chapters in this book discuss about the climate variations of timescales ranging from a few months to a decade.

This book discusses those variations with a particular emphasis on the interaction between two of the major components of the climate system, namely, the atmosphere and the ocean. Both components are extremely complex in their own respects with many internal processes that maintain their inherent variability while contributing to the maintenance of the Earth’s heat budget as discussed in the following section.

1.2 Atmospheric heat budget

The variation in the atmosphere is mainly driven by the thermal gradient. The major source of this thermal energy is Sun, which directly and indirectly (e.g., through the oceans) influences the atmospheric heat budget. First, the radiative energy-related heat budget is discussed here. The other form of the energy, called the turbulent energy is discussed in Section 1.4. The incoming solar radiation (in short wave) from the Sun goes through several processes after entering the Earth’s atmosphere and only about 51% of the total solar radiation reaches the Earth’s surface (Fig. 1–2). Diffusion and absorption by atmospheric molecules and clouds reduce the incoming energy. In addition, a considerable amount of energy is also reflected back from the ocean and land surfaces. In total, 30% of the energy is reflected from Earth’s surface and atmosphere back to the outer space. The actual energy absorbed in the upper layers of the ocean may vary within wide limits according to the nature of the surface, season, time of the day, and cloudiness. Since Earth is a black body, it also radiates back the energy albeit in infrared radiation. This happens after reaching an equilibrium between thermal emission emitted from the sea surface (ground surface as well) and from that of the atmosphere above it (Fig. 1–2) through a complex exchange of radiations.

Figure 1–2 Typical pathways of energy transfer in the global energy budget. Solar radiation is either reflected by the Earth’s surface or clouds or absorbed by the atmosphere or surface. The surface exchange (mostly turbulent) includes sensible heat and latent heat associated with conduction and evaporation. Some part of terrestrial (infrared) radiation emitted from the Earth’s surface is absorbed by the atmosphere and clouds, some part escapes to outer space, and some other radiated back to the surface by clouds and greenhouse gases.

Let us estimate the temperature of Earth considering the black body radiation and Stefan–Boltzmann law. In this theory, the power emitted per unit area is given by σT⁴, where T is the uniform absolute temperature of Earth and σ is the Stefan–Boltzmann constant. Since the power is emitted in all directions from a total surface area 4πa² ("a" being the radius of Earth), the total power emitted by Earth is 4πa²σT⁴.

This outgoing radiation should be equal to the incoming radiation from Sun if Earth is in a thermal equilibrium. The incoming solar radiation can be estimated by the formula (1−A)Fsπa² based on the total solar irradiance (or the solar constant) Fs on a disk and the energy reflected back (albedo) A. So, the balance between incoming and outgoing radiation will be used to estimate Earth’s equilibrium temperature T:

By inserting the values of A (0.3), Fs (1370 W m−2), and σ (5.67×10−8 W m−2 K−4), we will obtain T≈255K (around −18°C). This is obviously lower than the observed global mean temperature, which is about 288K (~15°C). It is clear that the above black body radiation model is missing out some aspect that is inherent in the Earth’s atmosphere. In fact, similar differences (actually, a large difference) in estimates and actual temperature can also be found for Venus.

The layer of atmosphere is what we have missed out in the above calculation using the radiative equilibrium assumptions. What if we have a layer of gases, which allow Sun’s shortwave radiations to pass through but selectively restrict the Earth’s longwave radiation to escape? And in fact, that is what is seen in case of Earth and Venus. The Earth’s dry atmosphere is composed mainly of nitrogen (78.1%), oxygen (20.9%), and argon (0.93%). These gases have only limited interaction with the incoming solar radiation and have almost no interaction with the outgoing longwave radiation emitted by the Earth. However, the dry atmosphere also has a number of trace gases, such as carbon dioxide, methane, nitrous oxide, and ozone. These gases, with a total of less than 0.1% by volume, play an essential role in the Earth’s energy budget. This is because of their ability to absorb and emit the Earth’s longwave radiation, unlike the nitrogen and the oxygen. Because of their ability to absorb and emit longwave radiation, while allowing solar radiation to pass through, these gases are called the greenhouse gases (since they just act like greenhouses in farming). It may be noted here that the water vapor in a moist atmosphere also acts like a greenhouse gas. By choosing right ratio of transmission (of solar radiation) and absorption (of longwave radiation) we can reach a radiative equilibrium that is just right to yield a global mean surface temperature closer to what is observed. But any change in that well-balanced ratio, for example, an increase or decrease in the amount of greenhouse gases, is bound to disrupt Earth’s thermal equilibrium and we will end up having a warmer or a cooler global temperature.

At the moment, it is believed that our anthropogenic activities are adding up more greenhouse gases leading to the global warming depicted in the SST anomalies (Fig. 1–3). On the other hand, some dust particles (usually from volcanic eruptions) and aerosols released to the atmosphere could act in an opposite way to cut down the solar radiation leading to Ice age, as it has happened in the past. Coming back to the recent historical records of SST anomalies, shown in Fig. 1–3, we find low-frequency variations in the global monthly mean time series on interannual-to-decadal timescales on top of the long-term global warming trend. We will focus on those interannual to decadal variations for most part of the discussions in this book in addition to the air–sea interactions on intra-seasonal timescales as reviewed by Li and Wang (2020) in Chapter 2, Impact of Atmosphere–Ocean Interactions on Propagation and Initiation of Boreal Winter and Summer Intraseasonal Oscillations, of this book.

Figure 1–3 Monthly sea surface temperature anomaly averaged over the globe relative to a 1982–2011 climatology derived from Characteristics of Global Sea Surface Temperature Analysis Data (COBE SST; Japan Meteorological Agency, 2006). The dark solid line is a 7-year running mean applied on the monthly anomalies.

The net incoming solar energy (i.e., after subtracting the part of the solar energy that is reflected back to space from the incoming solar radiation) varies between equator and poles. Tropical regions, being closer to the Sun, receive much of the solar radiations throughout the year, whereas the mid- and high-latitude regions (beyond 30°N and 30°S) receive much less energy compared to tropics with a large seasonal variation. On the other hand, the outgoing longwave radiation (Earth’s radiation) varies much less as a function of latitude. This is because polar regions though receive very little incoming solar radiation because of the atmospheric and oceanic transports they remain warm enough to emit quite large amounts of outgoing longwave radiation. As a result of this meridional difference in the distribution of the incoming and outgoing radiations, tropics become the net gainer, whereas the mid- and high-latitudes become net looser of the radiative energy.

The tropics receive about 60 W m−2 more incoming solar radiation than outgoing longwave radiation and the higher latitudes almost loose about 100 W m−2. This is reflected in the equator-to-pole distribution of zonally averaged global SST (Fig. 1–4). We see that the 24°C isoline comes closer to 30°N in August and closer to 30°S in February with a 2 months lag from the solar cycle, and peak radiations in June and December, respectively. Fig. 1–4 also shows the seasonal migration of the warm temperatures to the summer hemisphere (albeit less pronounced in the Southern Hemisphere owing to more ocean areas with large heat capacity). It may be noted here that the ocean currents and associated meridional overturning, especially in the North Atlantic Ocean, explain the northward displacement of the intertropical convergence zone from the equator during most part of the year. The tropical regions are considerably warmer than the mid- and high-latitudes and that difference in heating basically drives the atmospheric and oceanic circulations that in turn help in maintaining the Earth’s energy balance. Because of those circulations and associated energy transports, the Earth is approximately in energetic balance in annual average and the net energy loss in the polar regions is roughly balanced with the net energy gain in the tropical regions. In fact, if it were not those heat transports and heat storage in the ocean the gradients between poles and equator would be so huge that we cannot imagine a stable climate regime as we have now.

Figure 1–4 Zonally averaged global sea surface temperature (SST) climatology for the period 1982–2011 (°C) derived from Characteristics of Global SST Analysis Data (COBE SST; Japan Meteorological Agency, 2006).

1.3 Atmosphere and ocean circulations

Like any other fluid, atmosphere and ocean follow the general fluid dynamics; air/water flows from higher pressure/height to lower pressure/height. The net gain of heat in the equatorial region discussed in the previous section causes air to expand and generate a higher height at the top of the troposphere (lower part of the atmosphere; please refer to Andrews, 2010; Marshall and Plumb, 2008) in that region. Similarly, net loss in heat at the polar regions compresses the air and to generate lower height in those regions. Keeping everything else aside, this difference in the tropospheric height would create a pressure gradient favoring air to blow (water to flow) from equatorial region to the polar regions. The displaced air at the top of the troposphere in equatorial region must be replenished by rising air from the ground to keep a stable circulation.

As Earth’s gravity pulls the air and water toward the center of the Earth, the highest pressure in air and water is seen at the ground/bottom level. As a result, pressure decreases upward in both atmosphere and ocean. Like pressure, density (a function of pressure, temperature, and to some extent moisture for air and salinity for water) also decreases with height and in a stable stratification lighter air/water stays above denser air/water. This stratification can be disturbed by heating at the ground level and/or turbulence. The heating at the ground (through the net radiative fluxes discussed earlier) will lighten (and hence make the air buoyant) the otherwise denser air at the ground compared to air above. The buoyant air will rise adiabatically until it becomes stable with respect to its environment.

The stability of the rising air with respect to its environment will be determined by the lapse rate of rising air parcel’s temperature compared to the lapse rate of the environment. Depending on the amount of water vapor in the air parcel, either dry or moist (by including the effect of water vapor) adiabatic lapse rates (cf. Andrews, 2010; Marshall and Plumb, 2008) would be applied to determine the stability of the air. A typical moist adiabatic lapse rate is around 6K km−1 as opposed to a dry adiabatic lapse rate of 10K km−1 (which would usually be the environmental lapse rate where the air surrounding the parcel is dry). So, when the moist parcel is lifted in the air its temperature will be higher than the surrounding air due to this difference in the lapse rate. However, this will change with height as the moistures in the parcel get condensed out and the parcel reverts to an almost dry adiabatic lapse rate.

An adiabatic cooling of a moist air parcel will saturate the water vapor inside the parcel at some height leading to cloud formation. The latent heat released in the process will further increase the buoyancy to the parcel favoring instability. This is what is generally observed in equatorial/tropical regions. The rising air then travels to higher latitudes but because of the rotation of the Earth and the associated Coriolis force (cf. Holton, 2004; Pond and Pickard 1978; Vallis, 2006), the air traveling to higher latitudes in the Northern (Southern) Hemisphere turn to its right (left). Air heated in the tropics rises and moves poleward with high potential temperature (the temperature the air would have if it were brought back down to ground level while conserving energy), and as it cools to a lower potential temperature, it sinks while releasing the heat to the environment in its path. So, instead of traveling all the way to pole, the air will descend around the subtropical region and the descending air will then travel back to the equator at the surface level. A zonally averaged circulation will look like a meridional overturning cell known as the Hadley cell (Fig. 1–5).

Figure 1–5 Schematic of major features of the atmospheric circulation without the considerations of the effects of zonal variations due to land-sea contrasts and asymmetries around the equator and other latitudes. Hadley cell, Ferrel cell, and Polar cell are the schematic representations of zonally averaged meridional overturning circulations.

The poleward traveling air in the Hadley cell will experience greater rotation as it moves to higher latitudes due to Earth’s sphericity. Therefore the traveling air must spin faster relative to the Earth’s surface. Further owing to the rise in Coriolis effect, the winds become more zonal leading to the development of subtropical/mid-latitude westerly jets in the upper troposphere (Fig. 1–6). These jets are pronounced in winter seasons in both hemispheres and are very important for weather and climate variations. Especially, several chapters (Behera et al., 2020a, Chapter 3: Air–Sea Interaction in Tropical Pacific: The dynamics of El Niño/Southern Oscillation (ENSO); Marathe and Ashok, 2020, Chapter 4: The El Niño Modoki; Behera et al., 2020b, Chapter 5: Air–Sea Interactions in Tropical Indian Ocean: The Indian Ocean Dipole; Kosaka et al., 2020, Chapter 6: The Indo-Western Pacific Ocean Capacitor Effect; Richter and Tokinaga, 2020, Chapter 7: The Atlantic Zonal Mode: Dynamics, Thermodynamics, and Teleconnections) in this book discuss their role in atmospheric teleconnections in which the signal from tropical region is projected onto the jets and gets trapped in the associated Rossby wave path for the tropical signal to travel all around the globe.

Figure 1–6 A 200-hPa zonal wind climatologies for June to August (upper) and December to February (lower) derived from National Centers for Environmental Prediction (NCEP) reanalyses (Kalnay et al., 1996). Wind speed exceeding 30 m s−1 is shaded to demarcate the upper troposphere jet stream.

The Northern Hemisphere jet is not as continuous as its southern counterpart but in some extreme years they are uninterrupted enough to carry the signal around the globe (e.g., Behera et al., 2020b, Chapter 5: Air–Sea Interactions in Tropical Indian Ocean: The Indian Ocean Dipole of this book). The sinking air in the subtropical region heats the atmosphere and gets dry, depriving the region of moisture. Hence, most of the deserts are found in these regions. Also, the accumulation of air mass leads to the formation of surface high-pressure systems in those regions, some of which are important for subtropical climate variation over oceans (Morioka et al., 2020; Chapter 9: Interannual-to-decadal variability and predictability in South Atlantic and southern Indian Oceans, of this book). When the air returns toward the equator from subtropics at the lower level they start turning to their right/left owing to the Coriolis and surface frictional forces leading to easterly trade winds in both hemispheres (Fig. 1–5).

On the poleward side the traveling air from the subtropics reaches the poles through two other weak meridional overturning circulations, namely Ferrel cell and Polar cell (Fig. 1–5). Those are somewhat related to the residual circulations resulting from averaging many transient weather disturbances seen in the mid- and high-latitudes as weather fronts (Fig. 1–5). Those fronts help to defuse meridional temperature gradient between the tropics and the polar regions by transporting cold air equatorward and warm air poleward, generally accomplishing poleward heat transports. Those transients also help in transporting moisture from the tropics to higher latitudes besides the heat and the momentum. The rising air around 60°N and 60°S travels poleward in the Polar cell to subside near the polar regions. The return flow at the surface meets the westerlies of the Ferrel cell to complete the circulation.

1.4 Ocean circulation, upwelling, and climate variations

The situation is very different in the oceans as compared to the atmosphere. Since oceans are heated from the top (in addition to freshening by rainwater in some regions), the lighter water stays at top, that is, opposite of the atmosphere. This is referred to as stable stratification and a lot of energy will be needed to break such a stratification and to generate convections. The buoyancy-driven circulations are mainly in the areas of deep-water formation near the Arctic and the Antarctic where cold dense (and hence heavier) waters sink sometime all the way to the bottom and then spread toward the equator. This global circulation meanders through ocean basins interacting with topography and eventually get upwelled to surface via ocean mixing/diffusion as part of a global thermohaline circulation. But the timescale of that kind of circulation is very long with the order of a thousand-year and is not covered in this book. In the subtropical regions where there is a strong atmospheric subsidence at the descending branch of the Hadley cell, discussed in the previous section, the dry surface air and the cloud-free conditions promote strong evaporations and the evaporated water vapors are transported to the equator by the surface easterlies (Fig. 1–5). However, due to the loss of the fresh water at the ocean surface, the upper oceans get denser with heavier saline water. These regions then tend to produce intermediate waters where the salty water sinks to intermediate levels of about 700 m (e.g., pronounced in northeast Pacific). Those intermediate waters then travel to the equator in the subsurface to get upwelled and provide a mechanism for the decadal variation of the ENSO discussed in Chapter 3, Air–Sea Interaction in Tropical Pacific: The dynamics of ENSO, of this book (Behera et al., 2020a).

Most surface circulations in the oceans are wind-driven, the winds in the tropics and the subtropics (Fig. 1–5) give rise to the large upper ocean circulations called the subtropical gyres (Fig. 1–7). The gyres have the equatorial westward currents driven by easterly winds and strong western boundary currents like the Kuroshio in the northwest Pacific and the Gulf Stream in the northwest Atlantic. In the interior, currents are strongly influenced by the rate of change of the zonal wind with latitude and the currents are set by the rate of change of the wind and Coriolis force with latitude through a process called the Sverdrup balance (Pond and Pickard, 1978; Gill, 1982; Vallis 2006). Hence, in the subtropical gyres, currents flow slowly equatorward in most parts of the basins (Fig. 1–5). These gyres help to carry heat on the western boundaries by the strong western boundary currents in both hemispheres and bring cold waters from the higher latitudes to tropics in the interior and eastern boundaries of the oceans (Fig. 1–7).

Figure 1–7 Schematic representation of major features of the ocean circulations as depicted in the streamlines of ocean currents at 55 m for August (upper) and February (lower) climatologies derived from NCEP ocean reanalysis ( Behringer and Xue, 2004). Red and blue arrows indicate warm and cold currents respectively carrying heat from tropics to extratropics and vice versa.

Oceans also play an important role in the heat budget of the atmosphere and its heat transport. In the long term, the convergence and divergence produced by oceanic heat transports provide sources and sinks of heat for the atmosphere and partly shape the mean climate of the Earth. The oceans also exchange gases, water vapor, suspended particles, momentum, and energy with the atmosphere at the air–sea interface. These exchanges are often associated with the friction at the sea surface and the associated turbulences caused by the surface winds. Hence, these kinds of heat exchanges are often categorized as turbulent energy to differentiate them from the category of radiative energy discussed in Section 1.2 of this