Meteorology for Coastal Scientists [1 ed.] 9783030730925, 9783030730932

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Robert V. Rohli Chunyan Li

Meteorology for Coastal Scientists

Meteorology for Coastal Scientists

Robert V. Rohli • Chunyan Li

Meteorology for Coastal Scientists

Robert V. Rohli Department of Oceanography & Coastal Sciences, College of the Coast & Environment Louisiana State University Baton Rouge, LA, USA

Chunyan Li Department of Oceanography & Coastal Sciences, College of the Coast & Environment Louisiana State University Baton Rouge, LA, USA

ISBN 978-3-030-73092-5    ISBN 978-3-030-73093-2 (eBook) https://doi.org/10.1007/978-3-030-73093-2 © Springer Nature Switzerland AG 2021 This work is subject to copyright. All rights are reserved by the Publisher, whether the whole or part of the material is concerned, specifically the rights of translation, reprinting, reuse of illustrations, recitation, broadcasting, reproduction on microfilms or in any other physical way, and transmission or information storage and retrieval, electronic adaptation, computer software, or by similar or dissimilar methodology now known or hereafter developed. The use of general descriptive names, registered names, trademarks, service marks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use. The publisher, the authors, and the editors are safe to assume that the advice and information in this book are believed to be true and accurate at the date of publication. Neither the publisher nor the authors or the editors give a warranty, expressed or implied, with respect to the material contained herein or for any errors or omissions that may have been made. The publisher remains neutral with regard to jurisdictional claims in published maps and institutional affiliations. Cover illustration: NASA, Geocolor Image from NOAA’s GOES-16 Satellite of Powerful East Coast Nor’easter, 1/4/2018. This Springer imprint is published by the registered company Springer Nature Switzerland AG The registered company address is: Gewerbestrasse 11, 6330 Cham, Switzerland

Foreword

The coastal environment is among the most intensively used and chronically abused components of the Earth-ocean-atmosphere system. It is also home to an ever-­ increasing proportion of humanity with increasing developmental, recreational, industrial, transportation, and trade activities, amid ever-increasing impacts of natural hazards. Meteorology is a vital component of the coastal environment. Winds drive waves and currents, which in turn move sediment, nutrients, and pollutants, and shape and reshape coastal landforms. These meteorological-ocean dynamical processes impact human activities such as agriculture, fisheries, recreation, energy, land use development, and of course, hazard resilience. Given the increasingly intensive activities and fragility of the coastal environment, the rapidly-growing attention of scientists and planners to the coastal zone should come as no surprise. It is an indubitably welcomed development. Much contemporary work relies on the use of computer simulations that evaluate the impact of weather on the physical and human environments. Yet many if not most of the professionals studying and protecting our coastal environment have little background in weather and climate science. This book represents an effort to assist the many scientists and planners as they become familiar with the science of weather as it pertains to their area of expertise. This is not intended to be a textbook, but instead a reference manual for professionals. Several excellent meteorology textbooks at the introductory and advanced level exist. However, there remains no comprehensive meteorology primer for professionals that starts at the beginner’s level and proceeds through more specialized topics relating to the meteorology involved in modeling coastal processes. Simply stated, the introductory-level books ignore too many of the mathematical and physical processes to make them comprehensive enough to be useful to professionals, and the advanced books are too specialized and quantitative for most scientists who are not specialists in atmospheric processes. In this day in which coastal processes are increasingly being modeled numerically, it has become increasingly important for coastal scientists and planners to have an understanding of the fundamental physical processes that shape the coastal atmosphere, how these processes can be modeled, and the philosophy, assumptions, strengths, and shortcomings of such v

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Foreword

models. This text is designed to fill that niche. Today’s professional would likely lack the time to consult multiple texts to cover the breadth and depth necessary to become familiar with the content that would be presented in the proposed text. And even if multiple texts would be consulted, the interested professional would need to navigate through different symbology, notation, and terminology used by the authors of the various books on the market that would need to be consulted. In Meteorology for Coastal Scientists, the mathematics and physics are presented comprehensively, and in a way that allows non-specialist professionals to follow the physical explanations and equations as each step is shown and annotated. The result is a “one-stop-shop” – a comprehensive, step-by-step guide to both qualitative and quantitative understanding the atmospheric processes that govern coastal atmospheric processes. Thirty-three years have passed since Dr. S.A.  Hsu’s seminal work, Coastal Meteorology, was written. That textbook, along with Dr. Hsu’s other fine work, including the 2013 update of Coastal Meteorology, provided Louisiana State University (LSU) recognition as a leader in coastal atmospheric research. Eleven years after Dr. Hsu’s “official” retirement (though he maintains a very active research agenda), LSU is undergoing a resurgence in its coastal meteorology program. This text is a tribute to the work begun and inspired by Dr. Hsu. A strong case could be made that it is logical for Louisiana State University to be a leader in coastal atmospheric research, as Louisiana is perhaps affected more than any other U.S. state by the vagaries of coastal weather and climate. With the largest river system of North America crossing Louisiana into the Gulf of Mexico, the U.S.A.’s largest wetland, the world’s second-largest seasonal hypoxic zone off Louisiana’s coast, the most rapid rate of land loss in the U.S.A., constant vulnerability to oil spills and other environmental hazards wrought by the presence of a “working coast,” and perhaps the greatest vulnerability to hurricanes and other coastal storms in the nation, the effects of atmospheric forcing on waves, storm surge, coastal erosion, sediment transport, and Mississippi River and Atchafalaya River plumes should never be underestimated.

Content A new reference book in this area is necessary to accommodate a need for modern attention to the atmosphere of the coastal zone, where a large and increasing proportion of humanity dwells. The book is broader than just a “Coastal Meteorology” book per se, because its first half covers content that would be important for any scientist or planner who needs to understand atmospheric processes. The second half of the book puts the principles into practice in the coastal atmosphere, with heavy emphasis on processes that occur within the air-sea and air-sea-land interface. The book consists of seven parts: Introduction, Thermodynamics, Dynamic Processes, Coastal Weather Systems, Atmospheric Boundary Layers and Air-Sea Interaction, Air-Sea-Land Interactions, and Dispersion and Engineering

Foreword

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Applications. The concepts flow logically from the basic laws and equations that govern atmospheric energy exchanges and motion, but placed in a coastal zone framework. Scientific rigor is maintained, but not at the expense of comprehensiveness. This is not a climate change book per se, as climate change is treated as an implicit part of each discussion rather than as a topic unto itself. “Climate” implies not just “averages” but also “variability,” “extremes,” and “changes.” Thus, potential and/or likely effects of global climatic change will be incorporated into the presentation. For example, any discussion of the urban coastal atmosphere must include comments about impacts of rising sea levels. Our philosophy here is that climatic change is not to be separated from the rest of the discussion, as climatic change is an inherent component of each of the atmospheric properties discussed throughout the book.

Style Most professionals have little time to read detailed explanations. Therefore, the 52 chapters are kept as short and succinct as possible. Although the book is not designed as a university textbook, its abundant, short chapters would appeal to today’s university students as well, as students increasingly appreciate smaller “chunks” of information. Cognitive theory also supports the notion that learning is enhanced with smaller “packets” of information with abundant “shingling,” review of content, and presence of images on the same page that they are presented in textual form. A strong effort is made to design the book with these pedagogical principles in mind. When introducing new, complicated terms and processes, the phenomenon is described first, followed by its name, whenever possible. This approach connects readers to something familiar – the process that they may have observed themselves, before introducing the unfamiliar – the term, thereby reducing the likelihood that the readers will be intimidated by or uninterested in the phenomenon. The readers will also be more likely to remember the term. Terms used for the first time in the book are placed in boldface. Italics is used on the first use in a chapter of a term that has already been introduced in a previous chapter. This strategy reminds the readers that they may need to brush up on that term if it is unfamiliar. However, to avoid unnecessary distractions, a few common terms, specifically “meteorology,” “climatology,” “air,” “water vapor,” and “atmosphere,” are not italicized in subsequent chapters after the chapter in which they are first used. We always welcome comments from readers on how the work can be improved. Baton Rouge, LA, USA August 2021 

Robert V. Rohli Chunyan Li

Acknowledgements

We wish to thank the many people who have supported us in this endeavor. First, we are grateful for the faith and trust that our partners at Springer have placed in us. We hope that we have returned their confidence in us by producing work of which they can be proud, on time and on budget. We are especially grateful to Neelofar Yasmeen, Robert Doe, Dewika Lachman, Corina van der Giessen, Pavitra Arulmurugan, Carmen Spelbos, and Tiruptirekha Das Mahapatra. In addition, reviewers spent substantial time sharing their professional insights with us, resulting in an improved product. Our many fine colleagues, especially in the LSU Department of Oceanography & Coastal Sciences and College of the Coast & Environment, have inspired us with their research and offered ideas to us that have contributed to the effectiveness of this book. Our students, past and present, have challenged us at least as much as we have challenged them. And our families have shown patient willingness to give us the time needed to complete this project. We thank them all.

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Contents

Part I Introduction to Our Coastal Atmosphere 1 Scope, Uniqueness, and Importance of Our Coastal Atmosphere������    3 2 Atmospheric Composition, Structure, and Evolution��������������������������    9 Part II Thermodynamics in Our Coastal Atmosphere 3 Energy Transfer and Electromagnetic Radiation ��������������������������������   27 4 Temperature ��������������������������������������������������������������������������������������������   49 5 Application of the Gas Laws in Meteorology����������������������������������������   57 6 The Hydrostatic Equation and Adiabatic Processes����������������������������   61 7 Atmospheric Moisture ����������������������������������������������������������������������������   67 8 Atmospheric Stability and Potential Temperature ������������������������������   77 9 Measuring and Estimating Atmospheric Stability��������������������������������   91 10 Using Thermodynamic Diagrams in Meteorology��������������������������������   95 11 Clouds�������������������������������������������������������������������������������������������������������  105 12 Precipitation Processes and Types����������������������������������������������������������  125 Part III Dynamic Processes in Our Coastal Atmosphere 13 Pressure and Winds ��������������������������������������������������������������������������������  139 14 Coriolis Effect������������������������������������������������������������������������������������������  147 15 Effect of Friction��������������������������������������������������������������������������������������  151 16 Centripetal Acceleration and the Gradient Wind ��������������������������������  157 17 Gravitation������������������������������������������������������������������������������������������������  167 xi

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18 The Seven Basic Equations in Weather Forecasting Models ��������������  171 19 Comparison of Weather Forecasting Models and General Circulation Models����������������������������������������������������������������������������������  187 20 General Circulation of the Atmosphere ������������������������������������������������  193 Part IV Weather Systems in the Coastal Zone 21 Air Masses������������������������������������������������������������������������������������������������  213 22 Atmospheric Lifting Mechanisms����������������������������������������������������������  219 23 Fronts and the Mid-Latitude Wave Cyclone ����������������������������������������  225 24 Thunderstorms����������������������������������������������������������������������������������������  243 25 Lightning��������������������������������������������������������������������������������������������������  259 26 Tornadoes and Waterspouts��������������������������������������������������������������������  267 27 Advising the Public About the Severe Weather Risk����������������������������  283 28 Tropical Cyclones������������������������������������������������������������������������������������  291 29 Coastal Flooding��������������������������������������������������������������������������������������  309 30 Coastal Drought ��������������������������������������������������������������������������������������  317 31 Winter Storms������������������������������������������������������������������������������������������  321 32 Sea Ice and Weather Systems������������������������������������������������������������������  331 33 Summary of Energy Transfer by Atmospheric and Oceanic Motion ������������������������������������������������������������������������������������������������������  341 Part V Atmospheric Boundary Layers and Air-­Sea Interaction 34 Introduction to Near-Surface Atmospheric Dynamics ������������������������  347 35 The Logarithmic Wind Profile in Neutral Stability Conditions����������  355 36 Non-neutral or Diabatic Wind Profile����������������������������������������������������  369 37 Introduction to the Transition (or Ekman) Layer��������������������������������  377 38 The Classical Solution to the Atmospheric Ekman Spiral ������������������  381 39 Fundamentals of Air-Sea Interactions ��������������������������������������������������  393 40 Weather Effects on the Coastal Ocean��������������������������������������������������  399 41 Wind Stress and Turbulent Flux Drag Coefficients Over Water��������  415

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Part VI Air-Sea-Land Interaction 42 Surface Fluxes of Energy, Moisture, and Momentum��������������������������  421 43 Sea and Land Breezes������������������������������������������������������������������������������  427 44 Coastal Fog ����������������������������������������������������������������������������������������������  437 45 Coastal Upwelling and Weather ������������������������������������������������������������  443 46 Atmospheric Impacts on Lake Processes����������������������������������������������  451 47 Coastal Jets ����������������������������������������������������������������������������������������������  455 48 Atmospheric Optical Effects in the Coastal Zone��������������������������������  469 49 Solar Radiation in Aquatic Systems ������������������������������������������������������  485 Part VII Dispersion and Engineering Applications 50 Meteorology and Climatology of Coastal Cities������������������������������������  493 51 Atmospheric Dispersion in the Coastal Zone����������������������������������������  497 52 Engineering Aspects of the Wind Profile ����������������������������������������������  505 Appendices��������������������������������������������������������������������������������������������������������  513 References ��������������������������������������������������������������������������������������������������������  517 Index������������������������������������������������������������������������������������������������������������������  519

Part I

Introduction to Our Coastal Atmosphere

Sunrise. (Source: UCAR Image Library)

Chapter 1

Scope, Uniqueness, and Importance of Our Coastal Atmosphere

Abstract  This chapter provides an overview of the scope, uniqueness, and importance of the coastal atmosphere. The distinction between meteorology and climatology is stressed, along with the basic functions of the atmosphere and its interactions with the other components of the surface–ocean–atmosphere system. Then, the structural features and the functional processes occurring in the layer of the atmosphere immediately adjacent to the surface, in which energy, matter, and momentum are exchanged vigorously with the surface, are described. Scale in coastal meteorology and the importance of coastal meteorology in our everyday lives are emphasized. Keywords  Atmosphere · Climate · Meteorology · Coastal meteorology · Planetary boundary layer · Atmospheric boundary layer · Surface boundary layer · Temperature inversion · Scale in meteorology

1.1  What Is Coastal Meteorology? Meteorology is the study of the instantaneous state of the lower atmosphere, mostly in the troposphere of roughly within 10  km above the surface in which significant vertical convection of air occurs. These features contrast with those of climatology, which is the study of the long-term state of the atmosphere, including not only average conditions but also extremes and variability of those conditions. Because instantaneous conditions tend to result from processes occurring solely in the atmosphere, such as the absorption of heat energy from the Sun, which heats the air or surface, meteorology tends to focus only on the lower atmosphere. By contrast, the long-term state of the atmosphere—the climate—is affected by exchanges of energy, water, and other matter, and momentum (the product of mass and velocity), all carried by winds in three-dimensional directions, between the atmosphere and the land surface (lithosphere), water surface (hydrosphere), ice and snow-covered surface (cryosphere), and life zones (biosphere—which crosscuts the lithosphere, hydrosphere, and atmosphere). Thus, climatology includes the © Springer Nature Switzerland AG 2021 R. V. Rohli, C. Li, Meteorology for Coastal Scientists, https://doi.org/10.1007/978-3-030-73093-2_1

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Fig. 1.1  The complexity of interactions in the natural climate system that contributes to climate in the coastal zone. (Source: UCAR image library)

interactions between the atmosphere and the other “spheres.” These interactions take the form of exchanges of energy, water (in all its phases), and other matter, and momentum (Fig.  1.1). In other words, climatology includes the study of the climate system. However, some studies in meteorology must, by nature, examine interactions between the atmosphere and the lithosphere, hydrosphere, cryosphere, and biosphere, because the interactions occur on instantaneous time scales. This branch of meteorology is called boundary-layer meteorology. The “boundary” refers to the edge of the atmosphere that interfaces with the surface, whether that surface involves land (lithosphere), sea (hydrosphere), ice or snow (cryosphere), or vegetation (biosphere). Boundary-layer meteorology, then, is very different from the perception that most people have of what meteorologists do. Not all meteorologists are weather forecasters!

1.2  Boundary Layer Structure and Function

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One of the most complicated parts of the boundary layer is the coastal zone, where the hydrosphere, lithosphere, cryosphere (in some cases), biosphere, and atmosphere all meet. The different surfaces all have different thermal properties, different amounts of water and other matter, and different amounts of momentum, in the form of wind, to exchange with the atmosphere. And the exchanges from each of those surfaces vary through time, both in the short term and long term. Some of those exchanges are instantaneous, and others occur over slightly longer time scales, such as within the day–night cycle, and still others occur on much longer time scales, such as across seasons or across many years. And each of those surfaces can exchange energy and matter with each other, independent of the atmosphere too. Thus, the coastal atmosphere deserves close attention as a specialized form of boundary-layer atmospheric science. It involves an inherently complicated array of interactions in the lower atmosphere.

1.2  Boundary Layer Structure and Function The exchanges of energy, matter, and momentum originate in and are dominated by the shallow layer of the atmosphere adjacent to the surface, termed the planetary boundary layer (PBL) or atmospheric boundary layer (ABL). The term “boundary layer” is borrowed from fluid mechanics and denotes that part of a fluid (liquid or gas—the atmosphere in this case) adjacent to a solid surface that is significantly influenced by the presence of that boundary. The PBL is of critical importance in the coastal atmosphere because the atmospheric properties of a surface that change drastically over short distances also change drastically over short distances. This makes the PBL change in extent, properties, and impact on the environment abruptly across space and time. The PBL extends upward perhaps to about 0.5–2  km (0.3–1.2  mi) above the surface. It is often, especially during daytime, capped by a short vertical zone in which temperature increases with height. This capping inversion, a form of temperature inversion—an increase in temperature with height—prevents significant mixing from taking place across it, and it traps pollutants, salt crystals from waves, dust, soot, pollen, and other matter beneath it. If you have ever looked carefully out of the airplane window when taking off or landing, you have probably noticed this capping inversion layer at the top of the PBL as a thin, “dirty” layer of air. The pollutants cannot get dispersed because of the inversion above it. The inversion means that there is a thin layer of colder air beneath warmer air. Since the colder air is denser than the warmer air, it cannot rise over it. Thus, the pollutants and other solid aerosols (soil particles, salt crystals from the ocean, volcanic soot, pollen, and other similar solids) carried by the air are “stuck” in that thin layer, with the inversion layer acting like the lid on a pot. The pollutants and other aerosols cannot fall to the surface because the rising motion near the surface (i.e., convection) keeps them suspended, like a snow globe when turbulence is introduced to it by shaking it. The main reason for the capping inversion is that the highest solid aerosols stirred up by

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convection absorb the most radiant energy from the Sun, making them warmer than the air below them. Thus, the thickness of the capping inversion depends on the thickness of the solid aerosols and intensity of solar radiation. Because the Sun’s radiant energy heats the surface more effectively than the atmosphere, the surface is usually warmer than the atmosphere immediately above it, and therefore convection currents are directed upward, pushing those aerosols as high as the surface-based turbulence can push them, with the inversion layer trapping them there. At night, when the convection quiets down, the PBL shrinks. When the surface heats up again during the daytime, the convection increases and the PBL gets thicker. If the convection gets intense enough, it will punch through the PBL’s inversion and you will be in for a bumpy takeoff or landing and maybe a thunderstorm. The inversion layer exists because the aerosols absorb solar radiation more effectively than the rest of the atmosphere. This absorption of solar radiant energy makes the topmost aerosols in the capping inversion warmer than those at the bottom of the thin, polluted layer, creating the inversion that prevents the turbulent convection currents in the PBL from interacting much with the air just above the PBL. Thus, the air just above the PBL has much less turbulence than the air within the PBL. This lack of turbulence allows for a smooth, fast, nonturbulent environment for coastal low-level winds called coastal jets to form just above the PBL. The lowest 10% or so of the PBL is called the surface boundary layer (SBL) or sometimes the Prandtl layer. The SBL is defined as the layer in which the atmosphere is influenced so much by the characteristics that, at a given instant in time, the rates of flow (or flux) of energy, matter, and momentum from the surface upward remain constant with height. This is not to say that the fluxes of energy, matter, and momentum are constant across time, however. They vary greatly from day to day and even minute to minute. Internal boundary layers can even exist within SBLs, if local surfaces can influence the air so directly that the fluxes are constant with height at a given instant up to a given height, while above that level the broader-scale surface holds sway to govern fluxes that are constant with height at a given instant of time. For example, in March, the Mississippi River in Louisiana is transporting snowmelt from the northern USA, so its water is very cold compared to the air above it. This cold water chills the air immediately above it, and that chilled air has very different properties from the “usual” March SBL air over Louisiana that is a product of the locally warm March weather. So, an internal boundary layer that might not even extend over the ten-foot levees adjacent to the river may form beneath the “regular” SBL, which is in turn beneath the PBL.  There are usually sharp discontinuities at each of the boundary layers’ boundaries. One ramification of this is that the Mississippi River could be very foggy when the fog is not noticeable beyond the levees. These (and other) atmospheric boundary layers differ by not just structure (i.e., where they are, how thick they are, what they are composed of), but by function (what processes occur within them). Two features might be noted of boundary layers in general. First, significant exchanges of energy, matter, and momentum take place between the surface and the fluid within a boundary layer, with the momentum exchanges influencing

1.3  Scale in Coastal Meteorology

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circulation (i.e., wind velocity). And second, sharp variations in the related properties (i.e., temperature, humidity, and fluid velocity) occur in the boundary layer, even though the rate of change of those properties in the vertical direction remains constant within the SBL. The “variations” take the form of both fluctuations in temporal and spatial gradients.

1.3  Scale in Coastal Meteorology The Earth–ocean–air system is even more complex in the near-surface coastal zone than elsewhere, because of the complexity of surfaces at the land–air–sea interface. Both in the coastal zone and elsewhere, atmospheric motions and processes resulting from them occur on a great variety of scales, from the order of a millimeter, such as the flow of air across a leaf’s pore, up to the circumference of the Earth (horizontal, such as the steering wind belts) or the depth of the atmosphere (vertical, such as a violent thunderstorm). Associated with these spatial scales is an equivalent range of time scales, from fractions of a second up to months, years, or even longer (climatic variation). In general, features at large spatial scales also have longer life spans (Fig. 1.2).

Fig. 1.2  Space-time correlation in meteorology. (Source: RVR modified from UCAR MetEd)

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1  Scope, Uniqueness, and Importance of Our Coastal Atmosphere

1.4  The Importance of Coastal Meteorology Coastal meteorology and the intricately related coastal climatology are particularly important because of the large and increasing population that lives near the coast. Coastal dwellers, and the industries and activities that support and are supported by them, including fisheries, industrial plants, and shipping facilities, are affected by serious meteorological hazards and by everyday weather events alike. The complications of the coastal environment make the relationship between people and their environment especially complex in the coastal zone. The principles of coastal meteorology and climatology are important from both science and policy perspectives. Basic science research on the coastal atmosphere can inform other sciences that are influenced directly by the coastal atmosphere, such as fisheries, wetland biogeochemistry, ecology, barrier island geomorphology, and many others. But coastal atmospheric science also has much broader implications. Science is needed to inform and enact effective policy, which in turn impacts planning and culture, with the hope of improved protection of life, property, and the environment.

Chapter 2

Atmospheric Composition, Structure, and Evolution

Abstract  This chapter introduces the most fundamental entities of the atmosphere that cause it to behave as it does, including the concepts of pressure, density, and volume. The composition of air, which is dominated by diatomic nitrogen and diatomic oxygen, along with widely varying (across space and time) concentrations of water vapor, carbon dioxide, which is increasing in concentration over time, and other constituents is then described. The conventionally defined layers of the atmosphere—troposphere, stratosphere, mesosphere, and thermosphere—are then introduced, along with the characteristic features and processes that take place within each layer. The ozone layer in the stratosphere is given particular attention, because of its essential role in absorbing harmful ultraviolet radiation. Finally, the general evolution of the atmosphere since the formation of the Earth is described, along with the reasons for the temporal changes in the concentrations of its major constituents. Keywords  Atmospheric layers · Atmospheric composition · Greenhouse gas · Troposphere · Electromagnetic radiation · Stratosphere · Ozone layer · Mesosphere · Thermosphere · Evolution of atmosphere

2.1  Definitions and Introduction The atmosphere is a thin envelope of gases surrounding the Earth (Fig. 2.1). These gases cannot escape to space because gravity pulls them to the surface and because the Earth’s magnetic field keeps them “trapped” near the surface. In meteorology, the Système international (SI) or fundamental metric unit of pressure (in Newtons of force per square meter, or N m−2) is the Pascal (Pa). The SI units of most direct relevance in coastal meteorology and their derivations are explained in Appendix A. Because the atmosphere exerts on the order of 100,000 Pa of pressure at sea level, atmospheric pressure is usually expressed in units of kiloPascals (kPa; 1000 Pa) or hectoPascals (hPa, or 100 Pa), to make the numbers more manageable. One percent of a Pascal is defined as a millibar (mb). Conveniently, © Springer Nature Switzerland AG 2021 R. V. Rohli, C. Li, Meteorology for Coastal Scientists, https://doi.org/10.1007/978-3-030-73093-2_2

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2  Atmospheric Composition, Structure, and Evolution

Fig. 2.1  The atmosphere as viewed from space. (Source: Australian BOM; http://www.bom.gov. au/info/GreenhouseEffectAndClimateChange.pdf)

the mb and hPa are equivalent. The average atmospheric pressure at sea level is 101,325 Pa, 1013.25 mb, 1013.25 hPa, or 101.325 kPa, all of which are equivalent to 14.7 pounds inch−2. Any value higher than these at sea level represents higher than normal pressure, and any value lower than these at sea level represents lower than normal pressure. Places with lower elevation will have higher atmospheric pressure than adjacent places at higher elevations, because there is more atmospheric weight (i.e., force) exerted on the lower-elevation site. So, a map of ground level pressure on a given day might resemble an elevation map, and that map would look the same day after day because the variation of ground level air pressure due to weather is usually smaller than that due to elevation variations from place to place. Meteorologists want to observe differences in pressure between different locations and from day to day from the same reference level so that the effect of differing elevations can be removed, so they adjust the pressures to their sea level equivalents with some extrapolation. This allows a comparison of daily weather changes both at a place and relative to other places. Meteorologists have identified a convenient way of measuring atmospheric pressure—the mercury barometer (Fig. 2.2). Air pressure pushes down on the mercury in the pool, forcing mercury up the glass tube, where its level is calibrated to the amount of air pressure exerted. At sea level, the mercury would normally go up 29.92 inches, which is equivalent to 1013.25 mb or any of the other measurements described above. Mercury is used because it is relatively immune to changes in properties at the range of temperatures experienced on Earth, and because it slides a convenient distance up the tube so that differences can be observed easily. The atmosphere’s density, represented by rho (ρ), is its mass per unit volume (kg m−3) at a given position. Because the atmosphere is compressible under pressure, its density changes vertically and laterally. The weight of the atmosphere mass

2.1  Definitions and Introduction

11

Fig. 2.2  The mercury barometer. (Source: RVR)

resulting from gravity contributes to the pressure, which compresses the atmosphere. The thicker the atmosphere, the greater the weight and the greater the pressure. As a result, lower levels in the atmosphere are generally more compressed and have greater density. The maximum atmospheric density occurs at the Earth’s surface at or even below sea level. At sea level, the density averages approximately 1.225 kg m−3, or about 1/800 that of water. Because the pressure force associated with the average sea level density amounts to about 14.7 pounds per square inch, your head is supporting hundreds of pounds of weight right now. But do not worry, the atmosphere does not crush us with all of that weight because a balanced force is being exerted outward from our tissue. Even ants are not crushed by the weight of the atmosphere! This balance of forces, however, is not always going to protect us if we dive into the ocean deeper than about 30 m, where it would have four times the sea level pressure. The evolution of our bodies restricts our capability to balance out relatively little extra pressure beyond that exerted by the atmosphere at sea level. Atmospheric density decreases very drastically (exponentially in fact) with increasing altitude, as is shown in Fig. 2.3—so much so, in fact, that the difference is noticeable even at the range of elevations where people live. At 1 km (3100 ft), the density is only 91% of that at sea level. At the top of Mount Everest (8.8 km), the density is nearly 40% of the density at sea level. Because much less mass of air can be inhaled within a fixed volume of human lungs, it is difficult to get enough oxygen at high altitudes, especially when the body is overexerted. This decreased density also reduces the lift of aircraft, and it limits commercial aviation to about 12 km. Technically, air is different from the atmosphere, even though the terms are often used interchangeably. Air is a mechanical mixture of gases (atmosphere) and aerosols—tiny solid or liquid particles that float in the atmosphere. A typical range of diameters of aerosols is about one nanometer (10−9  m) to 10,000 nanometers (10−5 m); most viruses, including coronavirus, fall within this range of diameters. These aerosols include water droplets in clouds, salt crystals lifted above the surface by waves crashing at shorelines, soil particles stirred up by winds, soot from volcanic eruptions, pollen, human-generated pollutants, and other particles. They are “stirred up” within the planetary boundary layer (PBL) and suspended. Even though most stay within the PBL, some escape much higher in the atmosphere. For instance,

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2  Atmospheric Composition, Structure, and Evolution

Fig. 2.3  Schematic of the relationship between atmospheric pressure, density, and altitude. (Source: UCAR; https:// scied.ucar.edu/learning-­ zone/how-­weather-­works/ change-­atmosphere-­ altitude)

a volcanic eruption can eject particles so high in the air that they are above the level where they can readily precipitate back down to the surface. Because of the waves crashing on shore, increased likelihood of intense human activities, and more luxuriant vegetative cover in general, the coastal atmosphere tends to have a greater concentration of aerosols than other areas.

2.2  Composition of the Air The gases that make up the atmosphere can be divided into two categories: uniform gases and variable gases. The uniform gases are those that are essentially constant in their concentration worldwide, at least in the first three layers of the atmosphere closest to the surface—the homosphere. So in this regard, the coastal zone is no different than any other place. Diatomic nitrogen (N2) makes up about 78% of the mass of the dry atmosphere. Diatomic oxygen (O2) comprises nearly all of the rest, at 21%. The remaining 1% of the dry atmosphere is mostly argon (Ar), with trace gases, such as neon, helium, hydrogen, krypton, and xenon. Ironically, the most abundant gases are not very important climatologically. Variable gases have concentrations that differ over space and/or upward in the atmosphere. By far the most abundant variable gas is water vapor, with a concentration that varies from nearly 0% in the Antarctic winter to 4% by mass near the surface over the tropical oceans. Of course, in general, coastal areas have a larger share of water vapor in their local atmosphere than inland locations. Water vapor is a variable gas because it can change phase easily to liquid or solid forms; in some

2.2  Composition of the Air

13

places and on some days there is more water vapor in the atmosphere than at other places and times. The amount of water vapor is not negligible! For example, at sea level with a temperature of 20 °C (68 °F) at a time when the air has half as much water vapor in it as it could possibly have at that temperature, there is 8.7 grams (g) of water vapor for every m3 of air. So, a 1000 ft2 apartment with an 8-foot-high ceiling under such conditions would have 1975 g, or a little less than 2  kg (or 4.3 pounds), of water vapor. Although water vapor is by far the most important greenhouse gas—an atmospheric gas that keeps the planet warmer and more amenable to life as we know it than it would be otherwise—carbon dioxide (CO2) is the next most important greenhouse gas. CO2, on average, occupies 0.042% of the mass of the dry atmosphere as of June 2021. Despite the coronavirus pandemic at the time, which limited human emission of this greenhouse gas, this value exceeded that of May 2019. The actual concentration of CO2 may vary slightly from place to place but it is quite uniform globally in general. As implied above, the concentration of CO2 changes over time. It has been increasing abruptly since even before the Industrial Revolution, mainly because of fossil fuel burning. For example, 30 years ago the concentration was 0.033% (Fig. 2.4). In fact, atmospheric CO2 also has varied quite significantly through geologic time, with smaller concentrations during times of colder Earth and higher

Fig. 2.4  Keeling curve showing atmospheric CO2 concentration by year at Mauna Loa Observatory, in parts per million by volume. (Source: NOAA; https://www.esrl.noaa.gov/gmd/ccgg/trends/)

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concentrations (even much higher than today’s values) during warmer periods in Earth history. Thus, it is apparent that CO2 has an important role in regulating Earth’s temperature, including in the coastal zone. The global CO2 concentration also varies in the intra-annual time scale, as is shown by the squiggles in the graph of Fig.  2.4, which is known as the Keeling curve. The local maxima in CO2 each year corresponds to times when the vegetation is been assimilating less atmospheric CO2. This happens just after the winter ends and before the leaves are restored as Northern Hemisphere spring begins, because the Northern Hemisphere has over two-thirds of the land on Earth. Likewise, the annual minima in atmospheric CO2 occurs in September, after a full growing season for the Northern Hemisphere’s vegetation. Another variable component of the atmosphere is ozone (O3), which comprises only 0.000007% of atmosphere, but is nevertheless vital to terrestrial life on the planet because of its absorption of harmful ultraviolet radiation, as will be described in more detail in the next section. Most O3 is in the layer of the atmosphere above the weather and climate layer, for reasons that will also be explained in the next section. But O3 also exists near the surface, where it is a pollutant released mainly by automobile exhaust. The phrase “good up high, bad nearby” is a way to think of ozone. Its concentration is variable across space, both vertically and horizontally in the atmosphere. Because coastal zones are inhabited more densely and have more intense human activities than other parts of the planet, they tend to have a greater concentration of surface O3 than other areas. But the “good” O3 up high in the atmosphere is no more concentrated in the coastal areas than elsewhere. In fact, it is less abundant in polar areas. Methane (CH4) is another variable gas that exists in trace concentrations. Only two of every billion molecules in the atmosphere are CH4, but these are very effective greenhouse gases. Coastal areas have more local CH4 concentrations than other areas, both because it is released as a pollutant and because natural processes common in coastal wetlands release CH4 to the atmosphere. Aerosols are another important constituent of the atmosphere. Liquid aerosols are primarily the water droplets that comprise clouds and fog. Solid aerosols include the types of particles described previously—salt crystals, soil particles, soot, pollen, and volcanic dust, along with ice crystals, which are often found in high clouds. Solid aerosols often act as condensation nuclei or freezing nuclei, because they encourage the conversion of water vapor to liquid water (i.e., condensation) or liquid water to ice (i.e., freezing) around them, forming a droplet large enough that can eventually be pulled down to the surface as precipitation by gravity. In the coastal zone, salt crystals from the ocean would seem to be the most important cloud condensation nuclei or freezing nuclei. However, recent research shows that salt crystals are not as common in cloud droplets as soot, soil particles, and volcanic dust. In addition, recent research suggests that vertically oriented clouds have droplet densities of about 100 cm−3 over oceans but 1000 cm−3 over continents. So, the coast appears to be a transition zone between very different types of precipitationformation regimes. As will be shown in the next chapter, both solid and liquid aerosols interact with solar radiation that is incident upon them.

2.3  Vertical Structure of the Air

15

2.3  Vertical Structure of the Air The air has several distinct layers, which can be defined based on the vertical profile of temperature that is characteristic of the layer (Fig. 2.5). The troposphere is the lowest layer of the atmosphere. It contains about 75% of the atmosphere’s mass and nearly all of its water. Thus, all weather of significance occurs in the troposphere. The troposphere is characterized by a typical decrease in temperature with height and significant vertical motion, or convection. It exists from the surface up to an average of about 13 km (8 mi) but varying in thickness across time and space. It expands when the air is warm; therefore, its extent depends on latitude and time of year. At the poles in winter, it may be about 8 km (5 mi) thick, and at the Equator it is about 20 km (12.5 mi) thick. The characteristic decrease in temperature with height in the troposphere amounts to an average of about 6.4 C° km−1 (18.5 F° mi−1), as shown in Fig. 2.5. This rate of decrease in temperature with height is known as the environmental lapse rate (γ). A lapse rate is simply a rate of change in temperature with height. A primary reason for the average γ of a cooling with height by 6.4 C° km−1 is that less of the Sun’s energy is absorbed in the air than reaches the surface. Thus, the surface is warmed more effectively by the Sun than is the air, which leaves the surface to radiate energy upward to the air. This radiant energy continues even into the night, without the Sun’s help, because it is stored by the surface during daytime hours. Thus, the part of the air closest to the surface—the troposphere, and even more specifically, the PBL within the troposphere and the surface boundary layer (SBL) within the PBL—is warmest. Another way of saying this is that 30

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Tropical tropopause 18 15

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Planetary Boundary Layer Surface Boundary Layer

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Fig. 2.5  Atmospheric layers. (Source: UCAR MetEd)

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heated air is transferred upward from the surface to warm the troposphere via convection, so the part closest to the surface gets warmed the most. As the air warms, it becomes less dense, because density and temperature are inversely proportional to each other, as long as pressure remains constant, according to the Ideal Gas Law:



p =ρ

R∗ T m

(2.1)

In Eq. 2.1, p is pressure (in N m−2, or Pascals [Pa]), ρ represents density (in kilograms per cubic meter [kg  m−3]), R∗ is the universal gas constant (8.314 × 103 Joules [J] of energy per kilomole [kmol] per Kelvin [K] in the Kelvin temperature scale [J kmol−1 K−1], which always remains the same for the same gas), m represents the molecular weight (kg kmol−1) of the gas, also unchanging for the same gas, and T is temperature measured (K). The warmed, less dense air then rises by the buoyancy force. Thus, the troposphere is characterized by convection and turbulence—especially in the part closest to the surface, the surface boundary layer. “Tropo” means “turn,” so the name “troposphere” is appropriate. While it is tempting to think that the highest parts of the troposphere might be warmest because they are closer to the Sun than is the surface, this is not an important factor, at least in the troposphere. Instead, the proximity to the Sun-warmed surface and the higher pressures of the lower troposphere are far more important factors. Thus, the lower troposphere is much warmer than the upper troposphere. The PBL and especially the SBL are the most extreme cases of these features of the troposphere. It should be noted that the γ in the troposphere over a local place on a local day may vary widely from the worldwide average of 6.4C° km−1. The rate even changes with height in the same place on the same day. It can even be negative, which means that the temperature increases with height, a temperature inversion, in some small layers within the troposphere, with those layers changing in vertical extent and lapse rate, across space and time. The troposphere is a complicated place. The stratosphere lies above the troposphere, at a global average height of 13–50 km (8–31 mi) above the surface. Unlike the troposphere, the characteristic feature of the stratosphere is an increase in temperature with increasing height, a temperature inversion, with a global average temperature of 0 °C (32 °F) at its top. The lower stratosphere, from about 15–35 km (9.3–21.7 mi) is the ozone layer, which protects terrestrial life from the Sun’s harmful ultraviolet radiation. Because the absorption of ultraviolet radiation occurs most effectively by the ozone molecules that are higher in the stratosphere, the temperature of the matter in the stratosphere increases with increasing height, creating the characteristic temperature inversion. To understand how the ozone layer works and how it causes the stratospheric temperature inversion, it is important to understand the nature of the Sun’s radiant energy. The Sun emits radiant energy at a range of wavelengths of electromagnetic

2.3  Vertical Structure of the Air

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radiation, but in general the Sun emits shortwave radiation. The shorter the wavelength, the more intense the energy. Electromagnetic radiation having wavelengths shorter than 0.40 micrometers (μm) can harm animal skin tissue and plant tissue. While the Sun’s peak amount of emitted radiant energy has a wavelength of about 0.50 μm, it also emits radiation at shorter wavelengths, including at wavelengths shorter than 0.40 μm. So, there must be some processes that can protect us from such dangerous radiation. These processes are illustrated in Fig.  2.6, which suggests that, in the upper atmosphere far above the stratosphere, the most abundant atmospheric gases, N2 and O2, absorb the solar radiation of shortest wavelengths well before it reaches us. Specifically, N2 molecules absorb energy of wavelengths less than 0.12 μm. This is very helpful, because this is the most harmful radiation from the Sun. However, the amount of radiation at such wavelengths is small. But fortunately, O2 absorbs energy of wavelengths between 0.12 and 0.18 μm—the window of wavelengths in which the next most dangerous solar radiation falls. But that still leaves the dangerous wavelengths of electromagnetic radiation between 0.18 and 0.40 μm. And, while not as dangerous as the radiation of wavelengths shorter than 0.18 μm, that window of wavelengths contains a much greater quantity of solar radiant energy. Without a process to remove most of this energy before it reaches the surface, life on Earth would be browbeaten by such energy.

Fig. 2.6  The process of ozone formation and photodissociation.JPG. (Source: RVR)

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Conveniently, a process is in place, primarily in the stratosphere, that comes to our rescue. In absorbing the shortwave radiation, O2 gets split (through photodissociation) into two monatomic oxygens (O and O; Fig. 2.6). As you are likely to have learned in basic physical science or chemistry courses, monatomic oxygen is unstable and has a propensity to bond with other elements and compounds. If monatomic oxygen bonds with a diatomic oxygen (O2, the second-most abundant atmospheric gas), it forms O3. The O3 can then absorb solar radiant energy with wavelengths between 0.18 and 0.34 μm. This process protects life on the surface from the majority of the most harmful remaining radiation from the Sun that is not absorbed by N2 and O2. But in the absorption process, the O3 molecule is itself photodissociated into O2 and O again, with the O2 again being liberated to absorb radiant energy of wavelengths from 0.12 to 0.18 μm, and the O seeking another O2 for the formation of another O3 molecule. The process continues zillions of times per second. A problem occurs if people put pollutants into the atmosphere that make its way up to the stratosphere and might attract an O3 molecule or O atom for bonding. In such a case, the O3 is participating in other chemical reactions that photodissociate it, and O cannot participate in chemical reactions that form O3. In such cases, O3 is photodissociated at a greater rate than it is re-formed. Chlorine and bromine atoms, including a family of pollutants called chlorofluorocarbons (CFCs), are the most likely culprits for these kinds of reactions. A simplified version of a chemical reaction involving chlorine is:

Cl + O3 → ClO + O 2

(2.2)

In other words, a pollutant chlorine atom bonds with O3 and breaks it down, leaving one fewer O3 molecule to absorb radiation of wavelengths between 0.18 and 0.34 μm. To make matters worse, the ClO then reacts with O to produce Cl + O2:

ClO + O → Cl + O 2

(2.3)

In other words, the chlorine monoxide then goes on to capture a monatomic oxygen, making it incapable of finding an O2 to bond with to form O3. The Cl is now liberated to break up another O3!

Cl + O3 → ClO + O 2

(2.4)

The process repeats continually, with up to 100,000 O3 molecules destroyed by every Cl atom! Treaties and policies over the last few decades limiting the use of CFCs have caused an amazingly effective “healing” of the ozone holes that had been observed in previous decades using remote sensing techniques. This lesson provides hope

2.3  Vertical Structure of the Air

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that investment in sound science, enactment of effective policy, and international cooperation can team up to solve other complicated environmental problems. You may be wondering what happens to solar radiant energy of wavelengths between 0.34 and 0.40 μm. These reach the surface by the process of transmission through the air, even without a “hole” in the ozone layer. Implications are that diseases and ailment such as skin cancer, wrinkling, premature aging, cataracts, and reduced health immunity remain concerns from too much exposure to the Sun. Likewise, dangers to ecological balance and destruction to crops and phytoplankton also remain. You may also be wondering why the “ozone hole” problem is more of a concern over the poles than in other parts of the Earth’s atmosphere. The polar (and especially the Antarctic) stratosphere is so cold that it contains icy particles in clouds that are not normally found in warmer areas. These icy particles transform chlorine (and bromine) pollution into more destructive chemicals. Moreover, atmospheric circulation around Antarctica isolates it from the rest of the world, hindering these dangerous chemicals from dispersing around the world. And finally, the sunlight is needed to energize the chemical cycle that destroys O3, and polar areas have no sunlight for many months (with long periods of sunlight for the other part of the year), so there are long periods when new ozone cannot form. So just before Arctic or Antarctic spring (March and September, respectively), O3 concentrations would be expected to be minimized. Recent research, however, has revealed that the seasonality of stratospheric O3 is far more complicated. Current research is investigating the reasons for the observed O3 climatology. The “ozone layer” is in the stratosphere because this is the highest level in the atmosphere at which oxygen concentration is dense enough to form ozone, which can then absorb ultraviolet radiation, quickly enough to allow it to prevent enough solar radiant energy at wavelengths between 0.18 and 0.34 μm from reaching the surface. The absorption of energy by ozone molecules (or by any molecules) causes a temperature increase in the molecules that do the absorbing. The highest molecules of ozone get the first chance to absorb the radiation and therefore are most likely to have a temperature increase. This process explains why temperature increases with height (i.e., the γ is generally negative) in the stratosphere. A simple analogy may be that the fish swimming near the top of the aquarium have the first opportunity to eat, while the bottom-dwellers only have the detritus that settles downward uneaten. The mesosphere is the layer above the stratosphere and extends from about 50 to 80 km (31–50 mi) above the surface (Fig. 2.7). The matter making up the mesosphere is too sparse to allow effective ozone absorption, and more of the miniscule amount of heat that reaches the mesosphere is received as radiant energy from the surface than from direct absorption of solar radiant energy. Therefore, the “closer to the surface is warmer” rule of the troposphere prevails again. At the top of the mesosphere, temperatures might approach –1406 °F. The first three layers (troposphere, stratosphere, and mesosphere) are sometimes collectively referred to as the homosphere, because the gases in the atmosphere are

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Fig. 2.7  Atmospheric layers. (Source: UCAR image library)

so well mixed, with the obvious exception of O3, which is most concentrated in the lower stratosphere. The fourth layer is different, however, because the gases become layered according to their density, due to the lack of enough matter in the form of winds to keep the atmosphere “stirred” effectively. This “layered layer” is sometimes called the heterosphere, but more often is known as the thermosphere, and begins at about 80 km (50 mi) above the surface. There is no real place where it “ends,” as the molecules and atoms just gradually and eventually give way to space. Temperatures in the thermosphere increase with height (i.e., a temperature inversion is present) because it is too far above the surface to be warmed by the surface; instead, the only source of heat is direct absorption of solar radiant energy, primarily by the sparse N2 and O2. Therefore, locations in the thermosphere that are closer to the Sun are heated more. However, there are so few molecules that if by some miracle you could survive there under such miniscule atmosphere pressures, you would freeze to death instantly, even though the temperature might be 1500 °F because the matter that contains the heat is so sparse. The boundaries between the various layers are not clear-cut—instead there are transition zones. The tropopause separates the troposphere and stratosphere. At the tropopause, pressure is only about 25% of its value at the surface. The stratopause separates the stratosphere from the mesosphere. The mesopause separates the mesosphere from the thermosphere. Layers in the atmosphere may also be defined by chemistry rather than by their temperature profiles (as in the troposphere, stratosphere, mesosphere, and thermosphere). In addition to the ozone layer, another chemistry-defined layer is the ionosphere. In the upper mesosphere and thermosphere, there are large numbers of

2.3  Vertical Structure of the Air

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positively and negatively charged particles called ions. In the upper part of the ionosphere lies the magnetosphere, where Earth’s magnetic field interacts with the ions. The primary function of the ionosphere is to absorb very dangerous short wavelengths of energy that are emitted from the Sun, like cosmic rays, gamma rays, X-rays, and the shortest-wavelength types of ultraviolet radiation. These waves are relatively rare, but the ones that do get radiated into the Earth’s atmosphere would be very harmful to terrestrial life on the planet if they were not absorbed in the ionosphere. Reactions in the ionosphere also produce the brilliant light shows called the aurora borealis and aurora australis (Figs. 2.8 and 2.9) when the charged particles return to their “normal” states after being “excited” by the absorption of such energetic radiation, emitting radiation, and subatomic particles ejected by the Sun that are captured by the Earth’s magnetic field. The aurorae are more visible in the polar atmosphere than elsewhere because the atmosphere is thinner there, due to the constriction of the atmosphere under such low temperatures. Of course, the light shows are far easier to see at night than by day, but they occur both during the day and night. The ionosphere is composed of three sublayers. The D-layer is closest to Earth, absorbs amplitude modulation (AM) radio waves, and disappears at night. The E-layer is higher and weakens at night. The top-most layer of the ionosphere, the F-layer, reflects AM radio waves rather than absorbing them. As a result, it is possible to hear radio stations from hundreds of miles away, and most stations are required to reduce the power of their transmissions to avoid interfering with other broadcasts on the same frequency.

Fig. 2.8  Aurora Borealis. (Source: UCAR image library)

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Fig. 2.9  Another image of the Aurora Borealis. (Source: UCAR image library)

2.4  Evolution of the Air The best, prevailing, current consensus of scientific thinking suggests that today’s atmosphere came into existence after the expulsion of volatile substances from the Earth’s interior in association with volcanic activity. Volcanoes release NH3 (ammonia), CH4, H2O, N2, and other gases, but no O2. It is the same lack of oxygen that makes volcanic eruptions smell unfavorably today. Without oxygen in the atmosphere, life as we know could not have existed during the primordial Earth. And if there was no O2, there could not have been O3, because, as was shown previously, O3 forms from O2. Without O3, ultraviolet rays would have reached Earth’s surface, so life could not have begun on land. Thus, life must have begun in shallow water, as water is an effective absorber of ultraviolet rays. Amazingly, when an electric current is run through the gases emitted by volcanic eruptions under optimal circumstances, the result is the formation of amino acids, the building blocks of life! The most logical “next question” is: “Then how did oxygen and ozone get into our atmosphere?” Through a process that has yet to be fully explained by scientific research, subsurface-dwelling, one-celled organisms came to be and used fermentation to break down food. Fermentation releases CO2 into the air. The green curve in Fig.  2.10 suggests that life began on Earth at the time when CO2 entered the atmosphere. More complicated aquatic plants eventually came to be and took in the CO2 and released O2 into the atmosphere through photosynthesis. Notice on Fig. 2.10 that the O2 curve begins to increase at the time when photosynthesis began. And soon (in geologic time) afterward, the curve for CO2 began to decrease. Once O2 was released

2.4  Evolution of the Air

23

Fig. 2.10  Evolution of the atmosphere. (Source: RVR)

into the atmosphere, O3 could form, but its concentration remains so small that it does not even appear on a graph at this scale. Billions of years of geologic time have changed the composition of the atmosphere. CO2 came to be stored up in the shells and decaying plant matter, so its share of the air’s composition decreased over time, but in the last couple of centuries, we have been releasing it again. Notice the tiny little upward bend at the tail end of the CO2 curve. At the time, there was apparently nothing to take N2 out of the atmosphere, so it accumulated over time. The original atmospheric constituents participated in other processes or just got lost in the increasing population of molecules, so their percentages have decreased.

Part II

Thermodynamics in Our Coastal Atmosphere

Coastal Fog at Valdez, Alaska. (Source: NOAA Photo Library)

Chapter 3

Energy Transfer and Electromagnetic Radiation

Abstract  This chapter describes the mechanisms by which energy is transferred across space, and how these mechanisms are important in meteorology. Of particular importance is electromagnetic radiation—the mechanism by which energy from the Sun crosses through space and enters Earth’s atmosphere. Laws governing the intensity of this solar (shortwave) radiation can tell us the wavelength of peak intensity, the energy emitted at a particular wavelength, and the total energy intensity across all wavelengths, for hypothetical ideal radiators called blackbodies; the Sun behaves similarly to a blackbody. Thus, it is possible to know the amount of radiation entering the top of Earth’s atmosphere, at a given place and time—the solar constant. Once this energy enters the atmosphere, the amount that is neither absorbed, scattered, nor reflected by the air reaches the surface as transmission. At the same time, the air and the surface are also emitting radiant energy, albeit at longer wavelengths. The amount of energy involved in each process, at either short wavelengths (from the Sun) or longer wavelengths (from Earth and atmosphere), is described on a global, annual, average basis in the form of the energy balance. Convective and conductive fluxes help to “balance” the budget of radiation by moving warmer air within the Earth–ocean–atmosphere system. Keywords  Conduction · Advection · Electromagnetic spectrum · Blackbody radiation laws · Shortwave and longwave radiation exchanges · Atmospheric energy balance · Greenhouse effect · Sensible heat flux · Latent heat flux · Net radiation

3.1  Energy Transfer The Sun is the source of virtually all weather and climate, including in the coastal zone, because it provides the energy that is needed to drive motion in the atmosphere and to evaporate the water that becomes precipitation. In order to have energy to be used on Earth to drive weather and climate, it must be brought here from the Sun. There are three mechanisms by which energy can be transferred from one body to another: conduction, convection, and radiation. © Springer Nature Switzerland AG 2021 R. V. Rohli, C. Li, Meteorology for Coastal Scientists, https://doi.org/10.1007/978-3-030-73093-2_3

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Conduction is the transfer of energy from one molecule to the one next to it. An example of conduction is the heating of a glass pot handle even though it is not directly on the fire. Conduction is a slow, inefficient process that “wastes” a large percentage of the energy without it being transferred, and it does not transfer energy long distances. It is much more effective in solids than in liquids or gases because the molecules are much closer together in solids, but even in solids it is not very efficient. Because there is no matter between Sun and Earth for nearly all of the 93 million miles of distance separating them, conduction cannot get energy to Earth from the Sun. However, conduction can transfer energy within the Earth–ocean system that arrives here from other means, albeit very inefficiently. Convection is the vertical transfer of energy, often but not entirely in the form of heat, carried by moving air. The transfer of energy in the lateral direction—advection—is usually much smaller in magnitude, since hotter air rises, but when it occurs, it is known as advection. Like conduction, convection and advection require matter to be present to “carry” the energy. But unlike conduction, the matter that “carries” the energy in convection and advection is liquid or gaseous, because it needs to be capable of moving the energy rather than simply transferring it by bumping into the adjacent molecule. Thus, like conduction, neither convection nor advection can bring solar energy to the Earth–ocean–air system, but each can redistribute energy once it enters the system. Radiation involves the transfer of energy via electromagnetic waves. Radiation allows the transfer even when the radiating body is not in contact with the receiver of energy. No matter is necessary to serve as a medium of transfer—radiant energy can travel through a vacuum. Thus, radiation is the only mechanism that can transfer solar energy into the Earth–ocean–air system.

3.2  Electromagnetic Radiation The electromagnetic waves that carry energy are characterized by a wavelength (λ), usually expressed in millionths of a meter (μm). Wavelength is the distance between successive crests on the wave. Some forms of electromagnetic radiation have long wavelengths, and some have short wavelengths—the wavelength is a continuum. The shorter the wavelength the higher the energy associated with it. This implies, then, that hotter bodies emit their peak energy at shorter wavelengths than cooler bodies. All bodies have a curve associated with their emitted energy, such as the one shown in Fig. 3.1. Wien’s Law describes the wavelength of maximum emission of a hypothetical, ideal, maximally efficient radiator and absorber of radiation, a blackbody, given its temperature in Kelvins (TK, which will be covered in the next chapter). It is

29

3.2  Electromagnetic Radiation

Fig. 3.1  Sun’s electromagnetic emission spectrum. (Source: NASA)



max 

2897  m K TK

(3.1)

Wien’s Law shows that objects with a higher temperature have a shorter λ of peak emission. Because both the Sun and Earth behave nearly as blackbodies, we can use the radiation laws, including Wien’s Law, to describe their radiation properties. Since the Sun’s Kelvin temperature near its surface is about 6000 K, its wavelength of maximum emission as shown by the yellow curve in Fig. 3.2 is at 0.5 μm, a wavelength that falls in the visible part of the spectrum. It should come as no surprise that the Sun’s radiation is manifested as brightness to our eyes, and the peak wavelength of energy from the Sun makes it appear yellow to us, at least most of the time. The range of visible light is from about 0.4 (violet) to 0.7 (red) μm. At a temperature of 300 K for Earth, Wien’s Law suggests that the peak wavelength of radiant energy emitted by Earth is near 10 μm, which corresponds to the red curve on Fig. 3.2 in greatly exaggerated size, so that it is distinguishable. Thus, solar emitted radiation is often called shortwave radiation or shortwave energy, and Earth’s emitted radiation is referred to as longwave radiation or longwave energy or terrestrial radiation. As suggested previously, energy at shorter wavelengths is more harmful than that from longer waves—short waves are about the same size as

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3  Energy Transfer and Electromagnetic Radiation

Fig. 3.2  Sun and Earth emission spectra. (Source: UCAR MetEd)

Fig. 3.3  The electromagnetic spectrum. (Source: NASA/Goddard Space Flight Center)

human cells and can damage them. So sunburn is much more of a problem for living things than “earthburn.” Fortunately, the atmosphere—especially the Earth’s magnetic field and ozone in the stratosphere in the case of ultraviolet rays—absorbs most of the dangerous waves before they can reach us at the surface. The electromagnetic spectrum is the full range of all possible wavelengths (Fig. 3.3). From short to long, these waves include cosmic rays, gamma rays, X-rays, ultraviolet rays, visible wavelengths (violet to red), infrared, microwaves, and television and radio waves.

3.3  Solar Radiant Energy

31

Notice from Fig. 3.3 that the amount of energy that we can see (i.e., the range of wavelengths of visible light within the spectrum) is very small—from about 0.4–0.7 μm—centered on where the Sun’s peak emission is. Although the longer-­ wavelength thermal infrared radiation is generally not visible to humans, we can sense it as heat. This includes the wavelength of peak emission from Earth (about 10 μm). Even though Wien’s Law describes the wavelength of peak emission of a blackbody, bodies (like the Earth and Sun) still emit some energy at other wavelengths. The radiant energy emittance at any wavelength (Eλ) of a blackbody at a certain temperature can be determined by using Planck’s Law: c1

E 

 exp 5

 c2   1  T 

(3.2)

where c1= 3.74 × 108 W m−2 and c2= 1.44 × 104 K m. Planck’s Law provides all the points on the yellow and red curves in Fig. 3.2, with the curves just connecting the dots. Planck’s Law says, “Tell me what the Kelvin temperature of some surface is, and what wavelength you want to know about, and, if that surface is a blackbody, I will tell you the intensity of energy that the surface emits at that wavelength.” This equation is useful, especially when the equation is solved for Kelvin temperature, using remote sensing technology, which can capture the total energy emitted at a specified wavelength. For instance, scientists can measure the sea surface temperature from satellites, hundreds or thousands of miles away from any surface-based weather instruments, with slight adjustments made for the departure of the sea surface from blackbody radiation properties and for effects caused by the interactions of the radiation with the atmosphere. Remotely sensed surface temperatures derived using Planck’s Law are also used to monitor periodic ocean warming events known as El Niño events and assess global temperature trends by including remote areas. It is also possible to calculate the full-spectrum emittance of a blackbody (i.e., total emittance at all wavelengths—the area under the yellow and red curves in Fig. 3.2). This would be done using Stefan-Boltzmann’s Law:

ETotal   TK4

(3.3)

where σ is the Stefan-Boltzmann constant (5.67 × 10–8 W m−2 K−4). Equation 3.3 implies that an increase in temperature increases the total emitted energy tremendously.

3.3  Solar Radiant Energy The Sun is the only significant source of heat energy to Earth. Other sources include inner core nuclear reactions, longwave radiation emitted by the Moon, and shortwave radiation from the Sun reflected off of the Moon. But collectively, these

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sources amount to a very small fraction of a percentage point of the energy that we get from the Sun. And yet Earth still only intercepts one two-billionth of the radiant energy emitted by the Sun! The Sun’s output of radiant energy to the “top” of Earth’s atmosphere is almost constant. There are a few reasons for the deviation from a constant receipt of solar radiant energy. First, the Earth’s orbit is not circular about the Sun. At this time in Earth’s history, Earth is closer to the Sun in January (perihelion) than July (aphelion), so there is a fluctuation of ±3% of the mean value. It may come as a surprise to people in the Northern Hemisphere to know that their winter is the time when the Earth is actually closest to the Sun. But perihelion and aphelion change over time, by about 1 day per 65 years, making 1 complete cycle in about 21,000 years. During times in Earth’s history when perihelion came during Northern Hemisphere winter, those winters would be more severe, all other factors being equal. A second reason that the Sun’s output is not perfectly constant is that sunspots— flare-ups near the Sun’s surface—undergo an 11-year cycle. Sunspots are actually darker, “cooler” areas of the Sun, but brighter spots called faculae, which surround sunspots and are associated with magnetic storms, overcompensate for the “cooling” effect of sunspots. Thus, during years of maximum sunspots and faculae, the Sun is emitting slightly more energy toward the Earth. The solar “constant” is the mean shortwave radiation from the Sun at the “top” of the atmosphere, when and where the Sun’s rays shine perpendicular to the surface. In other words, the solar constant is calculated for the point when and where the Sun is shining directly overhead at a given time of day and year. It is approximately 1366 Watts per square meter (or W m−2). However, the amount of radiation at any given point on the Earth’s surface is not constant. It varies tremendously by time of day (it is zero at night), by time of year, and because of changing conditions in the air. For instance, if a cloud passes over a place, its receipt of solar radiation will decrease. The solar constant only represents the input of energy into the “top” of Earth’s atmosphere. Units such as W m−2 are important in any science that involves measurement of physical entities, and coastal meteorology is no exception. The Watt (W) is the Système international (SI) unit of power equal to one Joule (J) per second. The Joule is the SI unit of work (or energy), equal to the product of one Newton (N) of force exerted across one meter of distance. By definition, energy or work is equal to force times distance. The Newton is the SI unit of force equal to one kilogram moved at an acceleration of one meter per second squared or a change of velocity in one second by one meter per second (kg m s−2). How do we know this? Sir Isaac Newton, the Newton’s namesake, told us that force is equal to mass times acceleration, and the unit of the Newton is mass (kg) times an acceleration (m s−2). Become comfortable with these units, because they will be important in the explanation of pressure in Chap. 13.

3.4  Shortwave Radiation Exchanges in the Earth–Ocean–Atmosphere System

33

3.4  S  hortwave Radiation Exchanges in the Earth–Ocean– Atmosphere System What happens to solar energy after it reaches the top of the atmosphere? It can have several different fates, some of which are shown in Fig. 3.4. One possibility is that it is absorbed in the air. Absorption is the process by which radiant energy is converted to internal energy; that is to say, it becomes stored up within the molecules in the air. Temperature, the subject of the next chapter, is a measure of how much internal energy molecules have, and, therefore, how much absorption has occurred by that substance. The thermosphere has a high temperature because its (sparse) molecules get the first chance to absorb the Sun’s shortwave radiation and therefore get heated up. The air absorbs only about 25% of the solar (shortwave) radiation that reaches the “top” of the atmosphere. Much of the absorption is done not by atmospheric gases but by liquid water droplets and solid aerosols. And likewise, during daylight hours, those liquid and solid aerosols serve to minimize the surface

Fig. 3.4  The fate of incoming shortwave radiation from the Sun. (Source: UCAR image library)

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heating. Coastal areas have abundant liquid and solid aerosols, and this is part of the reason that the near-surface coastal zone is likely to remain slightly cooler in summer during daylight hours than adjacent inland areas. Scattering is another process that can interact with incident solar radiant energy (see again Fig. 3.4). In scattering, the radiant energy is deflected in all directions while its wavelength remains the same. Unlike absorption, scattering does not involve an increase in temperature of the scattering particle. Rayleigh scattering is the dominant scattering process when the scattering agent is among the smallest components of the air. Thus, Rayleigh scattering is most effective on gases. In Rayleigh scattering, half of the radiant energy goes forward and half goes in opposite direction, and the percentage of energy scattered is inversely proportional to λ4; in other words, the smaller the wavelength, the greater the percent scattered. Rayleigh scattering explains why the sky is blue; violet and blue light have the smallest wavelength so they are scattered more effectively than the longer-­ wavelength light (red, orange, yellow, and green), but because the human eyes are less sensitive to violet light, we perceive the sky as blue. Of course, near sunrise and sunset the sky does not appear blue or violet. At the beginning and end of the daylight hours, the Sun’s radiant energy must cross-cut the atmosphere so far across the horizon that the blue light is either absorbed or scattered so much that all we see is the less-scattered longer wavelengths of orange and red. To understand this better, imagine dropping a bag of sugar on the floor, which scatters the grains. You can probably still see the grains of sugar on the floor. But now stand on a ladder and drop the same bag of sugar on the floor. The scattering is so effective that the sugar grains are not as visible to your eye because of the increased distance and more effective scattering. Sunrises and sunsets appear more orange than red to us for the same reason that the afternoon sky appears more like blue than violet. But why are clouds white? Clouds are made up of bigger particles than gases (i.e., water droplets) and, therefore, Rayleigh scattering is only marginally effective. Instead, these larger particles scatter solar radiant energy mostly by Mie scattering, which is largely independent of wavelength. So violet, blue, green, yellow, orange, and red light (and nonvisible wavelengths of electromagnetic radiation too) are all scattered more or less equally. White is the color that results when the various colors are mixed. Mie scattering is most effective for clouds and aerosols—i.e., for cases in which particles (though still microscopic) are bigger in relation to the wavelength than for Rayleigh scattering. In Mie scattering, most of the scattered energy goes forward. A milky blue sky suggests that both Rayleigh and Mie scattering are occurring. Without scattering, the sky would appear black! Thunderstorm clouds are dark because they have so many water droplets in the clouds that light cannot emerge effectively through the cloud without being absorbed or scattered (i.e., through the process of transmission). The combined effect of absorption and scattering is known as attenuation. Dense rain clouds attenuate light effectively. By the way, why do we need to drive with our low-beam headlights instead of our bright lights when other drivers are nearby? Because there are likely to be bigger particles, such as water droplets (i.e., fog) or pollutants or soil particles

3.4  Shortwave Radiation Exchanges in the Earth–Ocean–Atmosphere System

35

suspended in the air, between your vehicle and an approaching vehicle, Mie scattering will be occurring and deflecting most of the energy in the forward direction. So, driving with the high-beam headlights would subject a driver of an approaching vehicle with an even more intense light, compromising the safety of all on the road. Reflection is another process that involves incident solar radiation and is shown in Fig. 3.4. Reflection involves a change in direction of the radiant energy—not a change in form. In other words, unlike scattering, after the incident radiation reaches the reflecting medium, the ray of energy remains in one “line” rather than deflecting into several “lines.” Reflection in clear, dry air is usually a minor factor, however, because gas molecules are not very efficient reflectors of solar radiation. Shortwave reflectivity, or albedo, is measured as a percentage of the incident radiation. For example, an albedo of 0.80 means that 80% of the incident energy is reflected. This would be a very high albedo, but it is realistic for fresh snow. Bright, light-colored objects have higher albedos than darker, dense objects. The high albedo of fresh snow is what allows it to remain on the ground even on a day with a temperature above freezing, since so much of the incident radiation reflects from it rather than infiltrates into it and converting to internal energy of the snow, melting it. Coastal environments have several distinctive features regarding albedo. First, clear-air albedo in coastal areas is likely to be larger than in most other atmospheres because there are higher percentages of solid and liquid aerosols than in most other environments. And second, because the albedo of water varies significantly with solar elevation, surface albedo over water also varies tremendously. When the Sun is high in the sky, such as around noon in summertime, albedo is small (perhaps less than 10%), especially if the water is clear. This is because the incident radiation can penetrate through the surface water layer rather than simply being reflected from it. But when Sun angle is low (i.e., Sun is near the horizon, such as near sunrise and sunset, particularly in winter), the albedo may be as great as 60 or 70%. You have probably noticed this as glare when you were in a boat around sunrise or sunset. Solar radiant energy that is neither absorbed, scattered, nor reflected will pass through the air and reach the surface unimpeded. The ability of a medium to allow radiation to pass through it is known as transmission, and transmission can depend on wavelength. For example, glass transmits radiant energy of short wavelengths much more effectively than the lower-energy long waves. You must have undoubtedly noticed how hot the inside of a parked car can be in summer. Rock, cement, and soil are generally poor transmitters of solar radiation energy because conduction is such a slow, inefficient process. Instead, solar radiant energy is retained near them. Dogs have this instinct: they know that a surface soil warmed by the Sun will be uncomfortable for a nap, so they dig a small hole, where the unexposed soil is much cooler, for a time. But eventually as that “new” surface accumulates solar energy as conduction fails to move it downward into the surface, they find that the hole needs to be dug just a little deeper in order to provide a cool respite for them. Eventually the yard has been destroyed in the humble yet understandable pursuit of finding a natural air conditioner. Transmission in coastal environments gives the coast some distinguishing features. Water is excellent transmitter, because sunlight can penetrate down below the

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3  Energy Transfer and Electromagnetic Radiation

water surface and be diffused by convection rather than by the less efficient conduction. This feature contributes to making coastal areas rather mild in temperature compared to nearby inland locations, as will be shown in the next chapter.

3.5  L  ongwave Radiation Exchanges in the Earth–Ocean– Atmosphere System Longwave radiation emitted by the Earth and air can have fewer possible fates than shortwave incident radiation. Notice from the orange arrows in Fig. 3.5 that longwave radiation is neither reflected nor scattered, so the longwave budget is similar—it is generally either absorbed or transmitted. Air is a much better absorber of terrestrial (longwave) radiation, on the whole, than it is for shortwave radiation. However, longwave radiant energy with wavelengths between 8 and 13 μm (which straddles the wavelength of peak energy emittance by Earth’s surface; see again Fig. 3.2), is not absorbed well by atmospheric gases. This range of wavelengths is called the atmospheric window. On cloudy days, the atmospheric window “closes” because thick clouds (i.e., liquid aerosols) absorb the longwave radiation that would otherwise escape into space. Kirchhoff’s Law says that an efficient absorber is also an efficient emitter of radiation. Therefore, after experiencing an increase in internal energy from the

Fig. 3.5  The fate of longwave radiation emitted by Earth. (Source: NASA)

3.6  Radiation and Energy Balance

37

absorption of longwave energy emitted by Earth’s surface, the clouds that absorbed the longwave radiant energy emitted by the surface then remit longwave radiant energy in all directions (albeit at a slightly smaller wavelength, due to the increase in temperature that they experienced in the absorption process). Some of this emitted radiation makes its way out of the Earth–ocean–air system to space. However, because so much more of the cloud mass is closer to the surface than to space, more of this absorbed energy is re-emitted back down to the surface as counter-radiation. Thus, counter-radiation is longwave (terrestrial) energy emitted by the air (particularly by clouds) downward and absorbed at the surface. Note that even if it is emitted by the air, longwave radiation is still considered “terrestrial.” You may have observed that a cloudy night is warmer than a clear one, as the counter-radiation warms the surface. Thus, the absorption of terrestrial (longwave) radiation at the surface relies heavily on the presence of liquid water, along with water vapor and solid aerosols, in air. Because the coastal atmosphere often has these solid and (especially) liquid aerosols in abundance, coastal nights are often warmer than nights in adjacent inland locations.

3.6  Radiation and Energy Balance Just as an accountant must keep track of where financial resources are, both at the micro-scale, such as for a given client, and at the macro-scale, such as for the company as a whole, a meteorologist must understand where energy resources are. Tracking shortwave and longwave radiant energy at the micro-scale is done using sophisticated field instrumentation. Tracking shortwave and longwave radiant energy at the macro-scale can be done using satellite imagery and Planck’s Law. While tracking resources of any kind at any scale can be tedious, understanding the radiation and energy budget is important in coastal meteorology because so many coastal studies require understanding how much energy is available for vegetation, for energizing atmospheric and oceanic circulation, and for other coastal applications. Even more important is understanding how that radiation and energy input might change in the future, and how those changes might cascade to changes in vegetative distribution and growth, atmospheric and oceanic circulation, and other applications. Any coastal meteorologist may find herself measuring the radiant energy fluxes, or at least reading about them, since the Sun ultimately energizes virtually all coastal processes. A convenient way to begin tracking energy is with the shortwave radiation balance. Rather than considering the variations with season and location, this discussion will be restricted to current, global, annual average values. And instead of specifying the number of Watts per square meter involved in each process, the discussion will start with 100 units of shortwave (solar) radiant energy reaching the “top” of the atmosphere. While at some locations at certain times of day and times of year, this will amount to the solar constant, at most it will be far less than the solar constant because the incident angle of the Sun’s rays on the “top” of the atmosphere is not perpendicular.

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3  Energy Transfer and Electromagnetic Radiation

The various components of the shortwave radiation balance are depicted in Fig.  3.6. Of the 100 units input into the Earth–ocean–air system, 19 units are reflected off cloud tops. More specifically, clouds cover about half of the air, on a global, annual average, and all of the different types of clouds have a mean weighted average albedo of 38% of the shortwave radiation incident on their tops. This yields a mean of 19% “sky albedo.” Figure 3.6 also shows that another 25 units, on average, are absorbed by solid and liquid (i.e., clouds) aerosols, and atmospheric gases. Scattering of shortwave radiation involves 6 units that are scattered from air (whether solid, liquid, or gas) back out to space and 20 units that reach Earth’s surface only after being scattered down to the surface after being transmitted through the air. The latter is diffuse radiation. It should not be surprising that so much more solar (shortwave) radiation is scattered down to the surface than is scattered back out to space, because so much more of the mass of the air is closer to the surface than to the “top” of the atmosphere. A total of 25 units reaches the Earth’s surface after being transmitted down through the air without being absorbed, scattered, or reflected first. Such shortwave radiant energy is known as direct radiation. In other words, less than half of the shortwave radiation that makes it to the top of the atmosphere even gets to the surface, through either direct or diffuse radiation. The 45 units of shortwave energy representing the combined direct and diffuse radiation is known as (K↓) (pronounced “K-down”).

Fig. 3.6  Global annual average shortwave radiant energy budget. (Source: RVR)

3.6  Radiation and Energy Balance

39

The remaining 5 units are transmitted through the atmosphere to the surface, reflected by the surface, and transmitted once again through the depth of the air and back out to space. This value is quite low for two reasons. First, as has been shown earlier, the albedo (K↑) of water is relatively low when the Sun has a high elevation in the sky. The Earth is mostly water, with an even larger percentage of water surface in the tropical latitudes where the Sun has its highest elevation. And second, the value is low because it is unlikely that solar energy will be neither absorbed nor scattered on two trips through the entire atmosphere. Despite its low value on a global, annual average, K↑ can be much higher locally, such as with a surface snow cover and under clear skies. But because the snow-covered part of the Earth is small and much of that area is only covered by snow seasonally, the global, annual average of K↑ is only 5%. The sum of K↑ and the shortwave radiant energy that is reflected or scattered out to space by the air yields the average planetary albedo, which is 19 + 6 + 5, or 30%. The various components of the longwave (terrestrial) energy balance are depicted in Fig. 3.7. It is noteworthy that Fig. 3.7 depicts more than a radiation balance, since it includes convective components too. Even though there is little scattering or reflection of longwave radiation, this balance is more complicated than the shortwave balance because there is no inherently logical starting point. When the surface acquired those 45 units of shortwave radiation (K↓), it then needed to emit longwave radiation in accordance with Kirchhoff’s Law. Most of that longwave radiation was absorbed by the air, in addition to the shortwave radiant energy that the air had already absorbed directly from the Sun. This leaves the air with even more energy than it had gotten from the Sun, so it must radiate more energy in accordance with Kirchhoff’s Law. That extra longwave energy absorbed by the air then is radiated from the air in all directions, including downward, back to the surface. This, in turn,

Fig. 3.7  Global annual average longwave energy budget. (Source: RVR)

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3  Energy Transfer and Electromagnetic Radiation

energizes the surface in addition to the 45 units of shortwave radiation that the surface had already acquired from the Sun. So, the air and the surface are energizing each other beyond the radiant energy that each acquired from the Sun. This process is known as the greenhouse effect and it explains why 104 units are emitted by the surface (L↑) as longwave radiation. If Earth did not have an atmosphere, it would have no greenhouse effect, the radiation balance would involve none of this recycling of longwave radiation, and both the surface and atmosphere would be much colder than at present. It is tempting to think of this longwave radiant energy going back and forth between the surface and air as “bouncing” or being reflected or being like a blanket, but really none of those descriptions or analogies are accurate. The L↑ gets absorbed into the molecular structure of the air (or surface), raising its temperature, and then, in fulfillment of Stefan-Boltzmann’s Law, radiates extra energy because its temperature has increased, and, in fulfillment of Wien’s Law and Planck’s Law, is radiated in longwave form. As suggested earlier and as shown in Fig. 3.7, the air is a very efficient absorber of longwave radiation. Of the 104 units of L↑ from the surface, 100 units are absorbed by air and only 4 units of the 104 escape into space as outgoing longwave radiation. In turn, 88 units are emitted by the air back to the surface (L↓, or counter-­ radiation). Even fewer units are involved in counter-radiation under clear, cold, nighttime skies, which leaves the atmospheric window “open.” Conditions on these kinds of evenings can cause freezes to sensitive vegetation, like orchards, so sometimes growers will install fans, like the ones in Fig. 3.8 from California, to bring warmer air above the surface back down to the surface.

Fig. 3.8  Orange groves in California; the white posts in the background are fans. (Source: RVR)

3.7  The Convective (Turbulent) Fluxes

41

The remaining 66 units are emitted by atmosphere to space. The atmospheric emission to space is much less than the emission to the surface because more of the air’s mass and high temperature (i.e., internal energy) is near the surface.

3.7  The Convective (Turbulent) Fluxes A closer look at the amount of radiant energy absorbed and emitted at the surface in Figs. 3.6 and 3.7 suggests that a problem exists. The sum of K↓ and L↓ (45 + 88, or 133 units) does not balance with the loss of energy from the surface (L↑ is 104 units). There is no need to include reflected short wave (K↑) because those 5 units would have to be added as they strike the surface and subtracted immediately as they are reflected away from the surface. But this imbalance by 29 units begs the question of why the surface is not warming every year by 29 units. We are indeed in a period of warming, but not by 29 units per year! Likewise, another look at the amount of radiant energy absorbed and emitted by air (see again Figs. 3.6 and 3.7) yields more cause for initial alarm. A total of 125 units are transferred into the air (25 units of short waves though absorption + 100 of long waves from L↑). But 154 units are lost by the air (66 + 88 units, both long waves). This imbalance, also by 29 units, also begs the question of why the atmosphere is not cooling every year by 29 units. The logical assumption and hope for humanity, of course, is that 29 units must be transferred from the surface to the air by some energy transfer process other than radiation. And fortunately, that is exactly the case. The imbalance is balanced by convection, or turbulence, within the Earth–ocean–air system, not by radiation. In convective, or turbulent, mixing, ascending parcels of air must carry more heat energy than the parcels of air descending to the surface. This amounts to 8 units on a global annual average, from Earth to air—the “sensible heat flux” (QH)—energy that can be sensed as heat. Figure 3.9 shows the global annual mean QH. It is obviously highest in areas that receive intense solar radiation without much cloud cover. These areas then would have abundant surface heating from K↓ and relatively low absorption of K↓ in air, creating a gradient of heat that would create a need to amplify the transfer of that energy upward by the QH. The remaining 21 units on a global annual average are moved by the “latent heat flux” (QE). “Latent” means “hidden” because this energy is tied up in the molecular structure of water, not in heat that we can feel, and the energy is absorbed or released in the phase changes of water. Sensible heat can be considered the energy that causes a temperature change and latent heat can be considered the energy that causes a phase change of water. Evaporation, melting, and water’s conversion from a solid directly to a gas— known as sublimation—require energy from the surrounding environment in order to cause water molecules to move to a “higher”—a more excited and spaced-apart— energy state. These involve a change in state from liquid to vapor, solid to liquid, or solid directly to vapor, respectively. The acquisition of this energy in latent form to

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3  Energy Transfer and Electromagnetic Radiation

Fig. 3.9  Global annual average sensible heat flux. (Source: UCAR MetEd)

energize the phase change of water to a higher energy state takes an equal amount of energy (in J kg−1) away from the surrounding environment in sensible form. This process causes the surrounding environment to cool. Some familiar processes occur that involve this acquisition and absorption of latent energy. For example, when a person begins to overheat, perspiration results, and the evaporation of the perspiration from your skin provides some relief in the form of cooling. Similarly, even if the temperature is 95°F (35 °C), a person will feel cold after leaving a swimming pool, because the water begins to evaporate from her skin. Another familiar example is that evaporating a bucket of water or melting a block of ice is likely to occur more quickly in summer than in winter, because in summer more energy is available to energize the water molecules to jump out of the bucket and into the air, or liberate themselves from the crowded, more static nature of their bonds in solid form (ice). More evaporation occurs near the surface than higher in the atmosphere because the vertical gradient of heat and moisture is much stronger in the planetary (and especially the surface) boundary layer than in the higher atmosphere. Thus, there is more QE in the planetary boundary layer than aloft. Water can also change phase from a “higher” to a “lower” energy state. Water’s transition from a gas to a liquid—known as condensation—along with freezing and water’s phase change from a gas directly to a solid—known as deposition—are the three cases of this type of transition. All release energy to the surrounding environment when the water molecules “lose” energy in the process. These processes warm the surrounding environment. One implication is that clouds become more buoyant as they form, since their formation involves condensation, freezing, and/or deposition. Contrary to what some may think, clouds are not composed of water vapor. More condensation occurs aloft than in the boundary layer, because the atmosphere is cooler and water gets cooled to the dew point temperature (Td)—the

3.7  The Convective (Turbulent) Fluxes

43

temperature at which the air becomes saturated and begins to condense water vapor and release latent energy. If Td is below freezing, it is known as the frost point temperature. Clouds are usually found higher in the sky, because the air is cool enough to reach Td or frost point temperature. This release of latent heat is felt at the molecular level by tiny perturbations. Amazingly, the sum of all the perturbations is the kinetic energy that drives the winds in a storm. So, then it is logical that clouds are associated with storms and with high winds, and that clear days are usually less windy. Not surprisingly, the largest QE occurs over tropical oceans, while the smallest is over deserts and polar areas (Fig. 3.10). Coastal areas have higher QE than inland locations at comparable latitudes. The “Bowen ratio” is the ratio of sensible heating to latent heating at a given place at a given time. Deserts and other warm, dry locations have a high Bowen ratio (>10.0 over a desert). Tropical, oceanic areas have a low Bowen ratio (32 km  hr−1, or 20 mi  hr−1) before hitting the South Carolina coast because he was caught in an area of a steep pressure gradient. Different sectors have different degrees of danger, because of forward velocity and circulation. In the Northern Hemisphere, the right “forward” side of the system is the quadrant in which the effect of the storm’s propagation and the wind speed of the storm itself are superimposed in the same direction to produce a strengthened effect of wind-driven waves. Moreover, such wind-driven waves, along with the overall storm surge (see Chap. 41), and (depending on the time of landfall), the nonlinear interaction of the storm surge with the astronomical tides, combine to make this right “forward” side of the storm even more dangerous. Furthermore, winds on the right side are usually stronger because the wind and forward velocity of the storm are in the same direction, thus a superposition of the two occurs.

28.5  Tropical Cyclone Motion

303

Atlantic-Caribbean-Gulf tropical cyclones may form immediately west of Africa, near the Cabo Verde islands, if the water is warm enough and the atmosphere is unstable enough there. These are called Cabo Verde storms or Cape Verde storms. Historically, the air is not unstable enough there early in the season, because the cold Canary Current (coming from the north) supports the subsidence inversion. Therefore, early in the hurricane season, Atlantic-Caribbean-Gulf tropical cyclones are more likely to originate farther west in the Atlantic Ocean, or in the Caribbean or Gulf of Mexico. Later in the tropical cyclone season, the percentage of Cabo Verde storms increases. Tropical cyclones migrate westward around subtropical anticyclones as they are steered by the trade winds. They may curve back northeastward and eastward as they pass the subtropical anticyclones. Once they move out of the tropics, westerlies can steer them back eastward. Often, the colder waters of the mid-latitudes help to kill the storms. Western Pacific storms that form off the Mexican coast are the best example (see again Fig. 28.1). In the North Atlantic, the warm, northward-flowing Gulf Stream keeps water warm. Therefore, areas as far north as New England and Canada’s Maritime Provinces can get hit by tropical cyclones. The outflow from one tropical cyclone can interact with the circulation associated with an adjacent tropical cyclone, in a Fujiwhara effect. A Fujiwhara effect can result in the merger of two tropical cyclones, or it can result in a diversion of the track in one or both storms. Tropical cyclones die over land because of friction and loss of evaporation and latent heat of condensation which fuel them. Friction slows the winds, which reduces the Coriolis effect. From that point on, winds usually, but not always—in the case of the brown ocean phenomenon, decrease rapidly. Heavy rain is likely to continue for much farther distances inland. In some cases, inland rainfall from tropical cyclones is welcome, since it arrives near the end of the growing season. In other cases, it can flood mature crops just before going to market and/or interfere with the harvest schedule and trade industries. Despite the fact that the features supporting and suppressing tropical cyclone development and the forces governing tropical cyclone motion are well-known, it is important to recognize that no two storms are the same, and that eccentricities in storm behavior are relatively common. For example, Barry in 2019 began as a cluster of thunderstorms over Kansas, before driving in an unusual track over the northeastern Gulf of Mexico, whereupon tropical characteristics were acquired. The very weak steering currents led many experts to believe that excessive precipitation exceeding 75 cm (30 in) were possible in low-lying coastal Louisiana. These ominous predictions were even more terrifying because of the record elevation of the Mississippi River at the time, prompting many to speculate about whether the storm surge could cause river water to overtop the levees confining the Mississippi and/or possibly breaching through the levee itself, as had happened in Hurricane Katrina in 2005. Fortunately, an intrusion of dry air north of Barry contributed to his asymmetrical shape and meant that the widespread, excessive precipitation never materialized. However, as much as 58  cm (23 in) of rain did fall in isolated areas of Louisiana, with record rainfall in Arkansas.

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28  Tropical Cyclones

28.6  Categories of Tropical Cyclone Strength Once cyclonic circulation develops and there is at least one closed isobar (with isobars drawn at a contour interval of four millibars), the tropical cyclone is called a tropical depression. If winds continue to strengthen to at least 17 m s−1 (37 mi  hr−1) and the closed circulation remains, the term “tropical storm” is applied and a name is assigned. The storm is called a hurricane (a typhoon in the East Pacific or a cyclone in the Indian Ocean or near Australia) when winds reach or exceed 33 m s−1 (74 mi hr−1). Winds of hurricane velocity will usually be sufficient to develop an eye. Tropical cyclones are classified by intensity according to the Saffir–Simpson scale (Table 28.1). “Major tropical cyclones” are 3 through 5 on the scale, but 5 is a very rare occurrence, especially at landfall. The 1 through 5 wind-speed-based Saffir–Simpson scale is reminiscent of the Enhanced Fujita Scale. Category 2 hurricane’s winds correspond closely to EF1 for tornadoes at the higher ends of both scales, while Categories 3 and 4 are similar to EF2 and EF3 tornadoes. This is because the wind speed thresholds for each category on both scales are designed to correspond to the ability of buildings of a certain code to be able to withstand the force of the winds. However, as will be shown in Chap. 40, the lack of consideration of propagation speed of the storm, the precipitation associated with it, ocean basin bathymetry, coastline configuration, and other factors, leaves the accuracy of the relationship between category and damage potential suspect.

Table 28.1  The Saffir-Simpson scale Tropical cyclone category

Sustained wind speed km hr−1 mi hr−1 37–63 23–38

Likely storm surge m ft