W E AT H E R : A C O N C I S E I N T R O D U C T I O N From a world-renowned team at the Department of Atmospheric Sciences at the University of Washington, Seattle, Weather: A Concise Introduction is an accessible and beautifully illustrated text covering the foundations of meteorology in a concise, clear, and engaging manner. Designed to provide students with a strong foundation in the physical, dynamical, and chemical processes taking place in the atmosphere, this introductory textbook will appeal to students with a wide range of mathematical and scientific backgrounds. This textbook provides a practical approach to the study of meteorology. It features: a single case study of a midlatitude cyclone which is referred to throughout the whole book to illustrate the basic principles driving atmospheric dynamics and phenomena; boxes on more advanced topics; appendices for additional coverage; chapter summaries listing the “takehome” points discussed; and color figures and charts clearly illustrating the fundamental concepts. Key terms are evident throughout, and a glossary explains the terms that students will need to understand and become familiar with. Gregory J. Hakim has undergraduate degrees in Mathematics and Atmospheric Science and a PhD in Atmospheric Science from the University at Albany, State University of New York. He joined the Department of Atmospheric Sciences at the University of Washington in 1999, where he served as Department Chair from 2012 to 2017 and is currently a Professor. He is also a leading scientist in the areas of weather analysis, predictability, and dynamics, and his research interests include weather and climate prediction, hurricanes, past climates, and polar circulation patterns. He has served on the advisory panel for the Directorate of Geosciences at the National Science Foundation, as Chair of the advisory panel for the Mesoscale and Microscale Meteorology Laboratory at the National Center for Atmospheric Research (NCAR), as a member of the NCAR Advisory Panel, as a member of the NCAR Strategic Planning Council, and as Chair of the University Corporation for Atmospheric Research’s President’s Advisory Committee on University Relations. Jérôme Patoux earned a Master in Environmental Engineering from the University of Texas at Austin and a PhD in Atmospheric Science from the University of Washington. He has been funded by the National Science Foundation (NSF), the National Aeronautics and Space Administration (NASA), the Office of Naval Research (ONR), and the National Oceanic and Atmospheric Administration (NOAA). He has taught undergraduate introductory meteorology for many years, and has been funded by the NSF to develop weather and climate curriculum. He is a former faculty member from the Department of Atmospheric Sciences at the University of Washington, and currently teaches meteorology at the University of Nantes in France. WEATHER A Concise Introduction GREGORY HAKIM University of Washington JÉRÔME PATOUX University of Washington University Printing House, Cambridge CB2 8BS, United Kingdom One Liberty Plaza, 20th Floor, New York, NY 10006, USA 477 Williamstown Road, Port Melbourne, VIC 3207, Australia 314–321, 3rd Floor, Plot 3, Splendor Forum, Jasola District Centre, New Delhi – 110025, India 79 Anson Road, #06–04/06, Singapore 079906 Cambridge University Press is part of the University of Cambridge. It furthers the University’s mission by disseminating knowledge in the pursuit of education, learning, and research at the highest international levels of excellence. www.cambridge.org Information on this title: www.cambridge.org/9781108417167 DOI: 10.1017/9781108264983 © Gregory Hakim and Jérôme Patoux 2018 This publication is in copyright. Subject to statutory exception and to the provisions of relevant collective licensing agreements, no reproduction of any part may take place without the written permission of Cambridge University Press. First published 2018 Printed in the United States of America by Sheridan Books, Inc. A catalogue record for this publication is available from the British Library. ISBN 978-1-108-41716-7 Hardback ISBN 978-1-108-40465-5 Paperback Cambridge University Press has no responsibility for the persistence or accuracy of URLs for external or third-party internet websites referred to in this publication and does not guarantee that any content on such websites is, or will remain, accurate or appropriate. Contents Preface Introduction CHAPTER 1 Weather Variables 1.1 1.2 1.3 1.4 Temperature 1.1.1 Heat and Temperature 1.1.2 Thermometers 1.1.3 Temperature Measurements 1.1.4 Temperature Scales 1.1.5 Radiosonde Profiles Pressure 1.2.1 Force and Pressure 1.2.2 Atmospheric Pressure 1.2.3 Vertical Distribution of Pressure 1.2.4 Barometers 1.2.5 Pressure Units 1.2.6 Some Useful Numbers Wind 1.3.1 Measuring Wind 1.3.2 Reporting Wind 1.3.3 Additional Sources of Wind Information Precipitation 1.5 Weather Stations Summary CHAPTER 2 Spatial Representations of Weather Data 2.1 The Station Model 2.2 Surface Maps 2.2.1 Isotherms and Temperature Maps 2.2.2 Temperature Fronts 2.2.3 Isobars and Pressure Maps 2.2.4 Highs, Lows, Ridges, and Troughs 2.3 Upper-Level Maps 2.4 Radar 2.5 Satellites 2.5.1 Visible Satellite Images 2.5.2 Infrared Satellite Images 2.5.3 Water Vapor Images 2.5.4 Geostationary Satellites 2.5.5 Polar-Orbiting Satellites Summary Appendix 2.1 Important Satellite Cloud Signatures Appendix 2.2 Contiguous USA Reference Map CHAPTER 3 Our Atmosphere: Origin, Composition, and Structure 3.1 Aspect 3.2 Composition 3.3 Origin and Evolution 3.4 Future Evolution 3.5 Vertical Structure Summary Appendix 3.1 Dynamic Equilibrium CHAPTER 4 Heat and Energy Transfer 4.1 Conduction 4.2 Convection 4.3 Radiation 4.4 4.3.1 The Nature of Electromagnetic Radiation 4.3.2 Temperature and Radiation Radiative Interactions 4.4.1 Absorption 4.4.2 Reflection 4.4.3 Scattering 4.4.4 Radiative Equilibrium 4.4.5 Selective Absorbers 4.4.6 A Window to the Sky 4.4.7 The Greenhouse Effect 4.5 Radiation and Weather 4.5.1 Heat Imbalance 4.5.2 Seasonal Variations 4.5.3 Diurnal Variations 4.5.4 The Influence of Clouds 4.5.5 Land–Ocean Contrasts Summary CHAPTER 5 Water 5.1 The Water Cycle 5.2 Saturation 5.3 Humidity 5.4 Relative Humidity 5.5 Humidity and Temperature 5.5.1 Relative vs. Absolute Humidity 5.5.2 Condensation 5.6 Dew Point Temperature 5.7 Applications of the Dew point Temperature 5.7.1 Surface Weather Maps 5.7.2 Meteograms 5.7.3 Radiosonde Profiles 5.7.4 Back to Relative Humidity 5.7.5 How to Saturate Summary CHAPTER 6 Cloud Formation 6.1 Adiabatic Processes 6.2 Adiabatic Processes in the Atmosphere 6.3 Dry Adiabatic Lapse Rate 6.4 Relative Humidity 6.5 Moist Adiabatic Lapse Rate 6.6 Orographic Lifting 6.7 Lifting by Convergence 6.8 Frontal Lifting 6.9 Convection 6.9.1 Stable Air 6.9.2 Unstable Air and Thermals 6.9.3 Stable vs. Unstable 6.9.4 Fair-Weather Cumulus Clouds 6.9.5 Conditional Instability and Cumulonimbus Summary Appendix 6.1 A Cloud Family Album CHAPTER 7 Precipitation 7.1 Warm vs. Cold Clouds 7.2 Collision and Coalescence 7.3 Ice-Crystal Growth 7.4 Precipitation Types Summary Appendix 7.1 Some Optical Phenomena CHAPTER 8 Wind 8.1 Force and Acceleration 8.2 Pressure Gradient Force 8.3 Sea Breeze and Land Breeze 8.4 Coriolis Force 8.5 Geostrophic Wind 8.6 Gradient Wind 8.7 Surface Winds 8.8 Friction 8.9 Topography 8.9.1 Mountain Breeze and Valley Breeze 8.9.2 Katabatic Winds Summary CHAPTER 9 9.1 Global Wind Systems The Averaged Atmosphere 9.1.1 Surface Temperature 9.1.2 Upper-Level Heights 9.1.3 Surface Pressure 9.1.4 Precipitation 9.2 The Single-Cell Model 9.3 The Three-Cell Model 9.4 Some Large-Scale Circulations 9.4.1 West Coast vs. East Coast 9.4.2 Antarctica 9.4.3 The Sahel 9.4.4 The Indian Monsoon 9.4.5 El Niño Summary CHAPTER 10 Air Masses, Fronts, and Midlatitude Cyclones 10.1 Air Masses 10.2 Fronts 10.2.1 Stationary Fronts 10.2.2 Cold Fronts 10.2.3 Warm Fronts 10.2.4 Occluded Fronts 10.2.5 Large-Scale Influences on Cyclone Structure, and the Tbone Model 10.3 Midlatitude Cyclone Development 10.3.1 The Life Cycle of a Midlatitude Cyclone 10.3.2 Vertical Structure of Cyclones 10.3.3 The February 2014 Cyclone 10.3.4 Where do Cyclones Form? Summary Appendix 10.1 Southern Hemisphere Midlatitude Cyclones Appendix 10.2 The Bergen School of Meteorology CHAPTER 11 Thunderstorms and Tornadoes 11.1 Ordinary Thunderstorm 11.2 Severe Thunderstorm 11.3 Lightning and Thunder 11.4 Supercells 11.5 Tornadoes 11.5.1 Description 11.5.2 Tornado Development 11.5.3 Tornado Alley Summary Tropical Cyclones CHAPTER 12 12.1 Facts and Figures 12.2 Tropical Cyclone Structure 12.3 Tropical Cyclone Development 12.4 12.3.1 Tropical Easterly Wave 12.3.2 Tropical Depression 12.3.3 Tropical Storm 12.3.4 Tropical Cyclone (Hurricane) 12.3.5 Tropical Cyclone Decay Conditions for Tropical Cyclone Development Summary CHAPTER 13 Weather Forecasting 13.1 Weather Forecasts and Uncertainty 13.2 Prognostic Equations 13.3 Ensemble Forecasting 13.4 Chaos and Weather Prediction 13.5 From Forecast Grids to Reliable Forecast Values 13.6 Making a Forecast 13.6.1 Medium to Long-Range Forecasting 13.6.2 Seasonal Outlook Summary Air Pollution CHAPTER 14 14.1 Pollutants 14.1.1 Gases and Compounds 14.1.2 Particulates 14.1.3 Photochemical Smog 14.2 Wind and Stability 14.3 Large-Scale Patterns 14.4 Topography Summary CHAPTER 15 Climate Change and Weather 15.1 Past and Future 15.2 Changing Composition 15.3 A Warmer World 15.4 An Altered Water Cycle 15.5 Changing Global Wind Systems 15.6 Midlatitude and Tropical Cyclones in a Warmer World 15.7 Beyond Weather 15.8 The Forecast Summary Glossary References Credits Index Preface Having taught introductory classes on weather many times, we came to see the need for a textbook on the subject that covers the foundations of meteorology in a concise, clear, and engaging manner. We set out to create an informative, cost-effective text that meets the needs of students who may not have any background in mathematics and science. The result – Weather: A Concise Introduction – is an introductory meteorology textbook designed from scratch to provide students with a strong foundation in the physical, dynamical, and chemical processes taking place in the atmosphere. This textbook is unique in that it: ► provides a concise and practical approach to understanding the atmosphere; ► introduces the basic physical laws early on and then ties them together with a single case study spanning the book; ► presents weather analysis tools early in the book to allow instructors to engage in discussions of current weather in tandem with the basic concepts, thus attracting and retaining student interest; and ► facilitates students’ learning and understanding of the fundamental aspects of weather analysis and forecasting, as well as practical skills, through a careful description of the forecasting process. Modern methods, such as ensemble forecasting, are central to the approach. Features Case Study: February 2014 Cyclone The main concepts of the book are illustrated in Chapters 2–13 by a single case study: a midlatitude cyclone that swept through the eastern half of the USA between February 19 and 22, 2014. This rich case study serves as a common thread throughout the book, allowing students to study it from multiple perspectives. Viewing the storm in the context of different topics provides a familiar setting for mastering new subjects and for developing an holistic understanding of midlatitude cyclones. Boxes on More Advanced Topics Instructors have the option of including more advanced coverage through use of boxes that provide insights on various topics. For example, in Chapter 1, Weather Variables, boxes include an in-depth description of the four laws of physics that are central to the study of the atmosphere. The book contains 25 boxes, affording instructors the opportunity to tailor the level of the material that they present to students in their course. Appendixes for Additional Coverage Appendixes at the ends of Chapters 2, 3, 6, 7, and 10 include additional material on important cloud signatures found in satellite imagery, the concept of dynamic equilibrium, the cloud classification, some optical phenomena, southern hemisphere midlatitude cyclones, and the Bergen School of meteorology. Summary A summary of key points has been included at the end of each chapter so that students can, at a glance, confirm that they have understood the significant take-away facts and ideas. Figures, Charts, and Maps Figures have been designed to convey the key concepts in a simple and selfexplanatory way, keeping in mind that clean representations of information are more helpful to students than complex drawings. Graphs and maps have been created with real data as much as possible, obtained from NOAA, NASA, ECMWF, and similar research-quality sources referenced in the text. Key Terms and Glossary The main text contains terms (in bold) that students need to understand and become familiar with. Many of these terms are listed in the Glossary at the back of the book. The Glossary allows the reader to look up terms easily whenever needed and can also be used to review important topics and key facts. SI Units We have consistently used SI units throughout the book, while providing alternative units whenever possible or relevant. Organization The first two chapters provide a general overview of key variables and weather maps used by meteorologists, which facilitates daily weather map discussions early in the course. We have found that motivating lecture topics with real-time examples using weather map discussions is a very effective way to engage students in the lecture material, and it allows instructors to introduce aspects of weather forecasting at their discretion well in advance of discussing the material more completely in Chapter 13. As a result, students are more invested in adding to their knowledge, which builds systematically toward understanding and predicting weather systems. Chapters 3–8 provide foundational material on the composition and structure of the atmosphere, along with the application of the laws of classical physics to emphasize and explain the role of energy, water, and wind in weather systems. Chapters 9–12 apply the foundational material to understanding the general circulation of the atmosphere (Chapter 9), midlatitude cyclones and fronts (Chapter 10), thunderstorms (Chapter 11), and tropical cyclones (Chapter 12). Chapters 13–15 build further on the first twelve chapters by applying the concepts developed to explain processes that affect how weather forecasts are made (Chapter 13), air pollution (Chapter 14), and climate change (Chapter 15). Instructor Resources A companion website at www.cambridge.org/weather contains PowerPoint slides of the figures in the text as well as a testbank of questions. Acknowledgments We thank: NOAA, NASA, and ECMWF for providing access to data and images; Reto Knutti, Jan Sedlacek, and Urs Beyerle for providing access to IPCC data; Rick Kohrs from the University of Wisconsin-Madison for providing global composite satellite imagery; and Paul Sirvatka from the College of DuPage for providing radar imagery. We also thank Ángel Adames, Becky Alexander, Ileana Blade, Peter Blossey, Michael Diamond, Ralph Foster, Dargan Frierson, Qiang Fu, Dennis Hartmann, Lynn McMurdie, Paul Markowski, Cliff Mass, Max Menchaca, Yumin Moon, Scott Powell, Virginia Rux, David Schultz, Justin Sharp, Brian Smoliak, Mike Warner, Steve Warren, Rachel White, Darren Wilton, Matt Wyant, and Qi Zhong, as well as 13 anonymous reviewers, for their help in the preparation of this book. This project would not have come to life without the support, help, influence, and constructive criticism from many fellow professors, teaching assistants, and students. We cannot acknowledge them all here by name, but we thank them nevertheless for the important role they have played in shaping the development of this book. Introduction Why should we study our atmosphere? Why should we learn about the causes and mechanisms of our weather? Weather affects our daily life: the clothes we wear (rain coat, shorts, hat, should we take an umbrella or sunglasses...?), the means of transportation we choose (walk, take a bus, ride our bike...?), our activities (ski, sail, water our plants, read a book in a coffee shop...?), and probably more. But beyond our daily concerns, weather affects society at large. Schools close when snow impedes traffic. Visitors to ski resorts might be more impatient for snow, while the ski instructors will be keeping an eye on the possibility of avalanches. Rangers are concerned with fog, thunderstorms, and flash floods. Fire patrols look for weather patterns that are conducive to forest fires (dryness, wind). Electricity providers are concerned by wind storms that can damage the infrastructure of the electrical grid and, on larger timescales, also need to plan how weather will affect upcoming energy needs (minimum temperatures impact heating, while maximum temperatures impact air-conditioning). Weather averages, such as prevailing winds, the typical temperature range, and mean precipitation determine how we build our homes and what locations are sensitive to extreme events, such as droughts, floods, hurricanes, tornadoes, etc. On longer timescales, we can ask how humans are changing the atmosphere, and what those changes imply for the weather and climate of the future. To start answering those questions, we need to understand how the atmosphere works. We need to identify the basic processes that drive the atmosphere, and the laws that govern atmospheric processes. By doing so, we will be able to explain the weather phenomena we experience around the year and throughout the world. Furthermore, we will also be able to apply these laws to the current state of the atmosphere, and predict how it will evolve in the future. © Caroline Planque There is a lot of value in becoming a knowledgeable observer of the atmosphere. After reading this book, you will look at the sky differently, you will gain an understanding of weather and climate that will make you more attentive to the world around you. You will have a basic understanding of weather phenomena, of cyclones, thunderstorms, and hurricanes, and you will understand the basic aspects of weather forecasting. You will see beyond the weather forecast you get on your phone, radio, TV, or the internet, and you will be able to make your own forecast in many situations. Weather and Climate Before we continue, let us clarify an important distinction between weather and climate. Weather is the condition of the atmosphere at a particular time and location. Weather varies on timescales of minutes to days. Climate, by contrast, is an average of the weather. It varies on timescales of decades to centuries and beyond. In this textbook, we will be mostly concerned with weather – even though many of the concepts have direct application to climate. Getting Started Our exploration of weather will start with a quick overview of important weather elements that we can observe or measure, and analyze. The choice of variables to observe is influenced by the laws of physics that govern the atmosphere. As we will see shortly, the atmosphere is made of matter (air and water etc.), it contains energy (heat), and it is in motion (wind, convection). Our understanding of weather is based on the fundamental notion that matter, energy, and motion obey conservation laws. To apply these conservation laws to the atmosphere requires observations of temperature (conservation of energy), pressure (conservation of mass), wind (conservation of momentum), along with humidity, precipitation, and clouds. One step at a time, and one building block over another, we will then investigate the physical processes that underlie the atmosphere at work. Finally, we will articulate these processes together to build a picture of weather systems such as midlatitude cyclones, thunderstorms, and hurricanes. In doing so, we will follow the precepts of René Descartes, who advocated, as early as 1637, that every difficult problem should be divided into small parts, and that one should always proceed from the more simple to the more complex. This cornerstone of the scientific method, still in favor today, will be an important aspect of our exploration as we elaborate a thorough understanding of the atmosphere from its most fundamental constituents at the molecular scale to its most complex inner workings as a system for moving heat at the global scale. CHAPTER 1 Weather Variables CONTENTS 1.1 Temperature 1.2 Pressure 1.3 Wind 1.4 Precipitation 1.5 Weather Stations Summary Where should we start with our study of the atmosphere? How should we first approach the weather? Like many scientists, meteorologists first make observations. Then they raise questions, and try to answer them. In this first chapter, we will quickly describe four of the elements, also called variables, of weather that meteorologists regularly observe, measure, and chart on weather maps, before we return to each of them for a more thorough exploration in subsequent chapters. Three of these elements are fairly intuitive: when concerned with the weather, we like to know how warm or cold it will be (temperature), whether it will be windy or not (wind), and whether it will rain or not (precipitation). The fourth variable, atmospheric pressure, is less intuitive, but it may be the most important to a meteorologist, as we will soon discover. Weather results from atmospheric changes. These changes obey certain rules, dictated by the laws of physics. In meteorology, three laws are of particular importance: the law of conservation of energy, the law of conservation of mass, and the law of conservation of momentum. Each describes a particular aspect of the atmosphere, and each requires that we measure certain variables of the atmosphere. The object of this first chapter will be to provide an overview of these variables, a starting point for our exploration of atmospheric changes. We will then return to each variable in subsequent chapters for a more thorough description and analysis. 1.1 Temperature Of primary interest to us is the law of conservation of energy (see Box 1.1). It states that energy is never created or destroyed, but only transferred between locations or transformed between different types of energy. In the atmosphere one form of energy is heat, and weather is largely the result of heat contrasts and heat transfers. Therefore we need to design ways of describing the amount and fluxes of heat throughout the atmosphere, which is accomplished by measuring temperature. Box 1.1. The Law of Conservation of Energy The law of conservation of energy states that the total energy of a system remains constant, if we account for the gains and losses of energy from and to the outside. Energy can be transferred between different parts of the system, or transformed into different types of energy inside the system (e.g., from potential energy to kinetic energy, or chemical energy, or radiative energy), but energy cannot be created (out of nothing) or destroyed. If we think of the atmosphere as a system, we can apply the law of conservation of energy to describe how energy, and in particular heat, is transferred or transformed in the atmosphere. And since heat transfers are related to temperature differences, we need to measure temperature and map these differences. That is one reason why temperature measurements are an integral part of weather observations. We will return to the conservation of energy in Chapter 6, in the form of the first law of thermodynamics. 1.1.1 Heat and Temperature We all have an intuitive feel for temperature, for whether things are hot or cold. But that feeling is very subjective – warm water feels cold after stepping out of a sauna. Science requires objectivity, which is sought by measuring and quantifying variables and processes. But how do we quantify that feeling of warm and cold? What is temperature, really? First, we need to return to the fundamental definition of heat as it is transferred to our body and the environment by contact and interactions at the molecular level. We can think of the atmosphere as a mixture of gases made up of molecules in motion (Figure 1.1). And we can think of heat as the energy associated with this molecular motion. In warm air, molecules are moving more rapidly than in cold air, and therefore have more energy of motion. We call this energy of motion, kinetic energy. Temperature is an indirect measurement of the average kinetic energy of the molecules constituting the air. Here it is useful to think of the air around us as made of individual volumes of air of uniform characteristics, called air parcels. [Although the size of such volumes of air is somewhat arbitrary, the concept of an air parcel will be useful later to explain various processes at work in the atmosphere.] While individual molecules in the parcel might all have different speeds, their average state is indicative of a higher or lower energy level, hence a higher or lower temperature of the air parcel. Figure 1.1. Random molecular motion in a volume of air. When fast-moving air molecules (at a higher temperature) are in contact with slow-moving air molecules (at a lower temperature), they impart some of their energy to the slow molecules, by contact, through collisions. In doing so, they lose energy, and therefore cool down, while the slower molecules gain energy, and therefore warm up. Kinetic energy, and therefore heat, has been transferred from the faster air molecules to the slower molecules, from the warmer to the colder air. We will see in Chapter 4 that this is the basic mechanism for transferring energy at the molecular level through conduction, but it is not a primary heat transfer mechanism for most of the atmosphere. 1.1.2 Thermometers How do we measure temperature? We make indirect measurements of average kinetic energy using thermometers. Liquid-in-glass thermometers, for example, work on the basis that substances expand or contract when temperature increases or decreases. A glass tube is filled with a liquid such as alcohol or mercury. When placed in warmer air, the energy of motion of the air molecules is transferred to the liquid in the reservoir by conduction through the glass. The molecules constituting the liquid now have more energy than before. Being more active, they push each other apart, which makes the liquid expand. In particular, it makes the liquid rise in the tube. Exposed to the same temperature, the liquid will always rise to the same level. Thus, if we attach a scale and units to the glass tube, and calibrate the instrument against known temperatures (such as the freezing point and the boiling point of water), we obtain an instrument that can measure any temperature (Figure 1.2). Figure 1.2. Thermometer in a standard weather shelter. Many other types of thermometer exist, and all make use of a property of matter to determine temperature indirectly. A bimetallic strip, for example, is made of two thin pieces of different types of metal attached to each other. Because the metals expand and contract at different rates, the combined strip bends when the temperature changes. If the tip of the metal strip is made to bend toward a temperature scale, the device can be calibrated and turned into an instrument. Some electronic thermometers use a material whose electrical resistance depends on temperature (a thermistor). Others, called radiometers, measure the radiation emitted by bodies. Because the emitted radiation is a function of the temperature of these bodies, we can once again indirectly deduce their temperature, as we will describe in Chapter 4. 1.1.3 Temperature Measurements As we will discuss at the end of this chapter, temperature measurements, like other weather measurements, are standardized, so that they can be compared and mapped. Weather stations always measure temperature in the shade, to avoid contamination by sunlight – recall that we are interested in measuring the kinetic energy of the air molecules surrounding the thermometer, and not the radiative energy contained in sunlight that might be absorbed by the thermometer. For this reason, temperatures are measured in a shelter, which is elevated at 2 meters (6½ feet) height above a vegetated surface, to avoid contamination by surface effects. Indeed, the air temperature can change rapidly near the ground, even within 2 meters (see Chapter 4 and Figure 4.14 in particular), and it is important that we measure the temperature of the air, not that of the surface. The shelter is painted white, and ventilated, to limit the absorption of sunlight and the concentration of heat inside the box, which could otherwise produce an artificially high temperature. 1.1.4 Temperature Scales You are probably familiar with the Fahrenheit temperature scale, named after Gabriel Fahrenheit, a German scientist who constructed the first mercury thermometer in the eighteenth century and calibrated it against three fixed points: 0, as defined by a mixture of ice, water, and sea salt; 32 in water and ice; and 96 “in the mouth or armpit of a healthy man.” Degrees Fahrenheit are commonly used in the United States. Scientists, and most other countries of the world, prefer to use the Celsius (or Centigrade) temperature scale, named after Anders Celsius, a Swedish astronomer who also lived in the eighteenth century and proposed the temperature of melting ice and the temperature of boiling water as fixed points (0 and 100, respectively). To convert from one scale to another, use the following formulas: ∘C=5/9×(∘F−32°)∘F=9/5×(∘C)+32° (Note that, in the first formula, 32 is subtracted before multiplying by 5/9, while, in the second formula, 32 is added after multiplying by 9/5 – a common source of mistake.) Scientists also use the Kelvin temperature scale, named after Lord Kelvin, a nineteenth-century scientist whose original name was William Thomson. The Kelvin scale is merely an offset version of the Celsius scale, translated in such a way that temperature measurements are always positive numbers. K= ∘C+273.15 °∘C=K−273.15 The smallest possible temperature, zero kelvin (0 K), is called “absolute zero,” and corresponds to the theoretical state in which all molecular motion stops, in which case molecules have zero kinetic energy, and therefore zero temperature. (Note that the symbol for kelvin (K) does not have the degree symbol – the little circle – in contrast with °F and °C.) Here, in keeping with the International System of Units, we use degrees Celsius on weather maps and temperature profiles. If you are more accustomed to degrees Fahrenheit, however, it will be useful to be able to convert from one scale to another, using the formulas above, or Table 1.1 for quick reference. If you hear or read temperatures in degrees Celsius and want to convert them quickly to degrees Fahrenheit, you can take advantage of the fact that 9/5 is approximately equal to 2, and 32 is approximately equal to 30. Thus, you can use the following approximation: ∘F≈2× ∘C+30 This is a very quick calculation that you can easily do in your head: multiply by 2 and add 30. For example, 20 °C multiplied by 2 is equal to 40. Add 30 to obtain 70 °F, which is close enough to the exact conversion, 68 °F. Table 1.1. Common temperature values in degrees Celsius and Fahrenheit °C °F °C °F −40 −40 0 32 −35 −31 5 41 −30 −22 10 50 −25 −13 15 59 −20 −4 20 68 −15 5 25 77 −10 14 30 86 −5 23 35 95 0 32 40 104 1.1.5 Radiosonde Profiles As we will see later in the book, a large part of our weather is dictated by what takes place aloft, as weather systems often extend up into the atmosphere and surface phenomena are often driven by upper air currents. Therefore it is useful to obtain information about the vertical structure of the atmosphere (Figure 1.3). We do so by launching helium-filled balloons carrying instruments that record weather variables up to 35 km altitude as the balloons ascend – helium is used because it is a very light gas (Figure 1.4). These balloons are called radiosondes, since the measurements are radioed back to a receiving station at the surface. Figure 1.3. Temperature profile obtained by radiosonde at Amarillo, Texas, on February 18, 2014, at 00:00 UTC. The red curve indicates temperature (T ) and the green curve indicates the dew point temperature (Td ). Pressure levels (given in hectopascals, hPa) are shown in blue and will be described in the text. Figure 1.4. Radiosonde being launched in Hawaii. Figure 1.3 shows the lower section of a temperature profile, in red, obtained by launching a radiosonde from Amarillo, Texas, on February 18, 2014, at 00:00 UTC. [Except where specified, we use Universal Time Coordinates (UTC) in the rest of the book, i.e., time referenced at the meridian of Greenwich, England.] The green curve indicates the dew point temperature, an important measure of the amount of water vapor in the air, which we will discuss in Chapter 5. Figure 1.3 also gives us an opportunity to describe a particular type of graph that we will encounter many times in this book. In such a graph, altitude is displayed on the y-axis (i.e., the vertical axis), while the variable of interest (temperature, pressure, wind speed) is displayed on the x-axis. Thus, the red curve in Figure 1.3 tells us that the temperature decreases from about 18 °C at 1 km altitude to −60 °C at 12 km altitude, which is in fact fairly typical. For reference, pressure levels are indicated by dashed blue lines, labeled in hectopascals (see Section 1.2.5). We can see that the temperature decreases up to about 200 hPa. We will soon learn that this level of the atmosphere (which changes with time and location) is called the tropopause and defines the upper limit of the troposphere – the layer of the atmosphere where essentially all weather takes place. 1.2 Pressure Atmospheric changes, and therefore weather, involve the movement of air, and the redistribution of air mass in the atmosphere. Such redistribution obeys the law of conservation of mass, which, in meteorology, is expressed in terms of pressure (see Box 1.2). It is therefore important that we measure pressure to analyze and predict the weather. As we did for temperature, however, it is useful that we first understand the nature of pressure in the atmosphere before we describe instruments to measure it. Box 1.2. The Law of Conservation of Mass Much in the same way as energy is conserved, matter, and therefore mass, is conserved in a system. It cannot be created out of nothing and cannot be destroyed: it can only enter or leave the system, or be transformed inside the system (by chemical reaction, for example, or by phase changes, as we will discuss in Chapter 5). This is described by the law of conservation of mass. In meteorology, the distribution of mass in the atmosphere is described by the distribution of density, which is defined as mass per volume. However, because density is difficult to measure, we instead rely on measurements of pressure and temperature to infer density (see Box 1.3). As a result, we can consider pressure measurements as informing us about the distribution of mass in the atmosphere. Notions of mass conservation and pressure will be important for understanding wind and weather systems in Chapter 8 and 10. Box 1.3. The Ideal Gas Law To very good approximation, the atmosphere behaves like an ideal gas, in which molecular collisions result only in a transfer of kinetic energy. As a result, pressure, temperature, and density do not vary independently of each other, but are related by the ideal gas law, pV=nR∗T where p stands for pressure, V denotes volume, n is the number of molecules in volume V, and T is temperature. R* is a universal constant that applies to any gas. For the atmosphere, the ideal gas law is more conveniently written: p=ρRT where ρ is the density (mass per volume) and R is a gas constant specific to the atmosphere. For example, if we heat a fixed volume and mass of air (i.e., the density remains constant), the kinetic energy of the molecules increases (i.e., the temperature increases), and the molecules bump into each other with more impetus (i.e., the pressure also increases). 1.2.1 Force and Pressure As we discussed for temperature, air is a gas made up of molecules in motion, and heat can be equated with the kinetic energy of the molecules. However, we also recognized that these molecules constantly bump into each other. Loosely speaking, the amount of bumping is what we call pressure. (We will come to a more exact definition shortly.) If the molecules bump into each other more frequently, or if the collisions themselves are stronger, there is more pressure. Thus, in a closed container, we could increase the pressure by either decreasing the size of the container at fixed temperature, which would bring the molecules closer to each other and would increase the number of the collisions, or increasing the temperature by heating the air, which would provide more kinetic energy to the gas molecules, and would result in more intense collisions. We can see here that pressure, temperature, and the volume of the container are indeed related, as expressed by the ideal gas law (see box 1.3). However, this is not to say that pressure and temperature are the same thing. Temperature, as a measure of heat, relates to the average kinetic energy of the molecules, independently of the collisions. Conversely, pressure relates to the collisions between molecules, which is not uniquely determined by the speed of the molecules, i.e., it is also a function of the number of molecules in the volume, and therefore the density of the air. In the atmosphere, things are slightly more complicated, as the air is not enclosed in a container, but we will address that issue shortly. In the meantime, we can think of pressure as the amount of bumping between molecules. And since the molecules are moving in all directions, we can see that pressure is applied in all directions as well. In practice, we often need to know what happens when air pushes on particular surfaces, either real, like the ground, or imaginary. For example, we sometimes think of the atmosphere as being made of air columns standing next to each other, and we are interested in knowing how much the air in two adjacent columns is pushing against the “wall” in between the columns. At other times, as described earlier, we will think of the atmosphere as being made of air parcels, i.e., volumes of air delimited by an imaginary envelope. Then we will be interested in knowing how much the air inside and outside the parcel pushes on the envelope. In all cases, pressure is applied to a surface, which results in a force pushing on that surface. As the area on which pressure is applied increases (i.e., more air molecules bump into that surface), the force increases proportionately. In symbolic notation: F=p ×A where p stands for pressure applied to a surface of area A, and F is the resulting force. Equivalently, pressure can be thought of as the force applied to a surface divided by the area of the surface: p=F/A which is, in reality, how pressure is defined in physics. In the above formula, if the force is distributed over a large area, the resulting pressure is small. But if the force is concentrated on a small area, the pressure is high. A simple application comes to mind. If you are wearing regular shoes, your weight, which provides the force, is distributed over the entire sole of your shoes, which results in a small pressure being applied to the floor. With high heels, however, weight is concentrated on a very small area, which results in a much greater pressure, and could dent a soft floor surface. 1.2.2 Atmospheric Pressure By analogy, we now consider the weight of the entire atmosphere applied onto Earth’s surface at sea level. Air might seem weightless, as philosophers of antiquity, including Aristotle, used to think, but gravity pulls air downward, keeping the atmosphere, like everything else, around Earth. In a column of air extending from Earth’s surface all the way up to the top of the atmosphere, the accumulated weight of all the air pushes down on the surface. Recall that weight is a force, so when we divide the weight of the air column by the area of the base of the atmospheric column, we obtain the pressure of the atmosphere, or atmospheric pressure. Note that, in this explanation, we are thinking about atmospheric pressure as pushing essentially downward, but recall that pressure at a point acts equally in all directions. 1.2.3 Vertical Distribution of Pressure We do not need to be at sea level for the concept of atmospheric pressure to make sense. We can repeat the same exercise at some altitude above sea level, say, 3000 m. Air pressure at that altitude is determined by how much the air is compressed due to the accumulated weight of the overlying layers of air. (Note that the amount of air below 3000 m is irrelevant in calculating pressure at 3000 m.) Since there is necessarily less air above 3000 m than there is above sea level, air pressure will be less than at sea level (Figure 1.5). Figure 1.5. Atmospheric pressure is tantamount to the weight of the atmosphere above us. To make a familiar connection with this decrease in pressure with altitude, notice the feeling you experience when you swim to the bottom of a pool – you experience an increase in pressure in your ears and nose. The atmosphere is like a swimming pool, and we live at the bottom of the pool – a pool of air. In the same way that a diver will experience higher pressure at depth in the ocean, atmospheric pressure is highest at ground level, and decreases upward (Figure 1.6 – pressure units are defined in Section 1.2.5). Figure 1.6. Pressure decreases with height. In summary, it is useful to think of atmospheric pressure as the weight of air above a point, or above a unit area. If we are at sea level, we are experiencing the weight of the total atmospheric column, and therefore higher pressure. If we are higher up in the atmosphere, the column above us contains less air than at sea level, and as a result the pressure is lower. 1.2.4 Barometers As the weather changes, air moves around, resulting in atmospheric pressure changes at a given location. We monitor these changes with an instrument that measures atmospheric pressure, which we call a barometer. In the seventeenth century, Evangelista Torricelli, an Italian scientist who was greatly influenced by the writings of Galileo, invented the principle of the mercury barometer. He filled a glass tube with mercury, inverted it in a cistern, and observed that the level of the liquid in the tube would not completely drop, but would stabilize at a certain height above the cistern (about 760 mm). He concluded that the air pressure applied downward onto the mercury in the cistern was forcing the liquid up into the tube to that level. He correctly speculated that the weight of the column of mercury in the tube was balanced by the weight of the atmosphere pressing down on the mercury outside the tube; i.e., the weight of the mercury column is balanced by atmospheric pressure (Figure 1.7). Figure 1.7. Concept of a mercury barometer. Torricelli’s design remained the basic principle behind most barometers for more than 300 years. If atmospheric pressure increases, the mercury level rises in the tube. Conversely, when a storm approaches and atmospheric pressure drops, a shorter column of mercury is required to balance the weight of the atmosphere, and the mercury level drops. That is why atmospheric pressure is often reported as the height of the mercury column, namely about 76 cm, 760 mm, or 30 inches Hg (“Hg” being the chemical symbol for mercury). Note that, in reality, Torricelli first experimented with water. However, about 10 meters of water are required to balance the weight of the atmosphere, i.e., about 33 feet, which made the experiment quite cumbersome! Since mercury is much denser, and thus heavier than water, a much shorter column of mercury is required to achieve the same weight. If your barometer at home is relatively small and round, you must be wondering how it can possibly contain 30 inches of mercury in a tube. And it does not, of course. You probably own an aneroid barometer, which is made of a small empty chamber (evacuated of some air to create a partial vacuum) that can expand and contract. When atmospheric pressure increases, it squeezes the chamber and forces it to contract. When atmospheric pressure decreases, the chamber expands. By connecting the chamber to a moving needle, the expanding chamber can be turned into a precise instrument (Figure 1.8). Figure 1.8. Aneroid barometer. 1.2.5 Pressure units In the International System of Units, pressure is expressed in pascal (Pa), where 1 Pa = 1 N/m2 (one newton per square meter), or alternatively in kilopascal (1 kPa = 1000 Pa). Note that a newton is a unit of force and a square meter is a unit of surface area, which is consistent with our earlier definition of pressure as force divided by area. In meteorology, however, since we think of atmospheric pressure as the weight of the atmosphere exerted on a surface, it can be more intuitive to express it in pounds per square inch (psi), or kilograms per square centimeter (kg/cm2) in the metric system. In that context, we can picture a column of air extending from Earth’s surface up to the top of the atmosphere. If the base of the column has a surface area of one square inch, then the weight of that air column is on average 14.6 pounds (at sea level); thus, atmospheric pressure is, on average, 14.6 psi. Meteorologists, however, go one step further. Since we are interested in the pressure exerted by the atmosphere, we may express pressure... in atmospheres! The average atmospheric pressure would then be close to “one atmosphere.” For historical reasons, one atmosphere is called “one bar,” after the Greek “baros,” meaning weight. Since daily variations in pressure are on the order of a thousandth of an atmosphere (i.e., a thousandth of a bar), meteorologists have historically used millibars (mb), where 1 bar = 1000 mb. The exact value of a bar has been chosen as 100 kPa for practical purposes, which is equal to 1000 hectopascal (hPa, where hecto- is the prefix for 100), so millibar and hectopascal can be used interchangeably. Putting it all together, the average pressure at sea level is approximately: 1013hPa−1013mb−29.92°Hg−14.6psi To be consistent with the International System of Units, we will use hectopascals (hPa) in the rest of this book. 1.2.6 Some Useful Numbers If you observe the variations in atmospheric pressure at a given location over time using a barometer, you will find that the typical range of pressure at sea level is about 980 to 1030 hPa. The pressure will only drop below 980 hPa during intense storms. Similarly, it will only rise above 1030 hPa during exceptionally high pressure events. If you carry your barometer upward, you will find that pressure drops very quickly with height, as we saw in Figure 1.6. As a rough rule of thumb, pressure near the surface drops about 8 hPa for every 60 m of altitude – this is easily noticeable, and measurable, in elevator trips in tall buildings. Recall the pool analogy: pressure is greatest at the bottom of the pool, or ocean of air, and decreases as we rise in the pool. Since we will be interested in what happens higher up in the atmosphere, it is useful to have some standard levels of pressure in mind and know their approximate altitude. Figure 1.9 shows some useful numbers to which we will refer often. We saw that the average pressure at sea level is 1013 hPa. We will often look at maps of weather variables at 850 hPa, which is around 1.5 km altitude (about 5000 ft), and near the top of what we call the “boundary layer” (i.e., near the top of the layer where the effects of the daily variations in temperature near Earth’s surface are felt). At 500 hPa, about half of the atmosphere is above us and the other half below (in terms of mass and weight), and this happens at about 5.5 km, or 18,000 ft. Aircraft cruising altitude is at about 250 hPa (roughly 11 km, or 34,000 ft), which is also the top of the troposphere. Thus, the troposphere, i.e., the first layer of the atmosphere, where most of our weather takes place, contains about 75% of the mass of the atmosphere. Note that 250 hPa also corresponds to the altitude where we find the jet stream, as will be discussed in Chapter 9. Above the troposphere, we enter the stratosphere, where the ozone layer is found, as we will describe in more detail in Chapter 3. Figure 1.9. Altitude and pressure in the atmosphere: some useful numbers to remember. Since water is 1000 times more dense than air, the increase in pressure with depth happens much faster in the ocean than in the atmosphere. Recall that the weight of the atmosphere is balanced by a 10 m column of water in a water barometer. In other words, a 10 m column of water has the same weight as a column of air extending from sea level to the top of the atmosphere. Therefore, we will add the equivalent of another atmosphere of pressure (i.e., one additional bar) if we dive 10 meters down into the ocean – another good number to remember. And for each additional 10 m layer of ocean, we will add another bar of pressure. That is why it is important for divers to pause regularly, especially on their way up to the surface, to let their blood and other internal fluids adjust to the change in external pressure. 1.3 Wind As mentioned before, weather results from atmospheric motions, and air motion is also constrained by the laws of physics. As for energy and mass, air motion is constrained by a conservation law: the law of conservation of momentum (see Box 1.4). Momentum is the product of mass and velocity. The density of air is estimated indirectly from measurements of pressure and temperature, using the ideal gas law (see Box 1.3). The wind velocity, however, needs to be measured. It has two components, direction and magnitude, which need to be measured separately. Box 1.4. The Law of Conservation of Momentum Weather is a result of the atmosphere in motion to redistribute heat. In meteorology, it is useful to describe air motion in the form of momentum, a product of the mass, or density, of the air and its speed, or velocity. Momentum tells us both how much air is in motion, and how fast it moves. As it turns out, much like energy and mass obey conservation laws, momentum is governed by the law of conservation of momentum, which is described by Newton’s laws of motion. In particular, Newton’s second law, which relates forces to changes in momentum, describes how the wind evolves. That is why we measure wind as part of our routine weather observations. Wind and Newton’s second law of motion will be the subject of Chapter 8. 1.3.1 Measuring wind Wind direction is measured using a wind vane. The vane points toward the direction that the wind is coming from (Figure 1.10), which is why meteorologists report wind direction as the direction that the wind is coming from. An easterly wind, for example, is a wind blowing from the east. A northerly wind blows from the north, etc. Figure 1.11(a) shows wind in terms of compass direction. Figure 1.10. A wind vane points toward the direction that the wind is coming from. Figure 1.11. Wind direction convention in meteorology. (a) Compass direction. (b) Degrees from true north. Meteorologists also use degrees from true north to report wind direction (Figure 1.11(b)). Degrees from true north are more convenient for computation and graphing, for example, since they are numerical quantities. Wind speed is measured with an anemometer. The most common type is the cup anemometer (Figure 1.12(a)). As the cups catch the wind, the spindle rotates, and the rate of rotation of the spindle is proportional to the wind speed. It is similar to propeller anemometers, for which the wind turns the blades in proportion to the wind speed (Figure 1.12(b)). Figure 1.12. Two different instruments to measure wind speed. (a) Cup anemometer. (b) Propeller anemometer. The units for wind speed are miles per hour (mph) in the English system, and kilometers per hour or meters per second in the metric system. (Remember that speed is distance divided by time.) Unfortunately, the most common unit used by meteorologists is knots, for historical reasons: mariners were affected by weather before airplanes were. Sailors used to measure the speed of their ship by letting a knotted rope run into the sea, and counting the number of knots per time as a measure of speed. One knot is one nautical mile per hour, and since one nautical mile is the distance of one minute of latitude on Earth (1.85 km, or 1.15 mi): 1 knot−1.85 km/h−1.15 mph−0.5 m/s. Mariners sometimes also use the Beaufort wind scale, which was first designed to relate wind speed to the sea state in the absence of anemometers; the rougher the sea surface, the higher the wind speed. It is now also occasionally used on land and describes typical damage observed during strong wind events (Table 1.2). Table 1.2. Beaufort wind scale Force Description Wind speed (knots) Wind speed (km/h) 0 Calm 0 0 1 Light air 1–3 1–5 2 Light breeze 4–6 6–11 3 Gentle breeze 7–10 12–19 4 Moderate breeze 11–16 20–28 5 Fresh breeze 17–21 29–38 6 Strong breeze 22–27 39–49 7 High wind 28–33 50–61 8 Gale 34–40 62–74 9 Strong gale 41–47 75–88 10 Storm 48–55 89– 102 11 Violent storm 56–63 103– 117 12 Hurricane force >63 >117 1.3.2 Reporting Wind As you know from experience, the wind does not blow constantly and regularly. It stops, resumes, increases, and decreases in strength over short time intervals, producing wind gusts. Figure 1.13, for example, shows a lot of “gusty” variability over the course of 2 minutes. Wind gusts are often due to the presence of swirls of air called “turbulent eddies” that mix the air up and down on a variety of scales (i.e., eddies of different sizes). Since wind speed typically increases with altitude and decreases as it approaches the surface, where there is more friction, turbulent eddies tend to bring higher wind speeds down to the surface intermittently (Figure 1.14). The resulting wind gusts can be twice as strong as the average wind and can damage trees, for example, or affect sailing. In the example shown in Figure 1.13, wind gusts reach speeds up to 20 knots. Figure 1.13. Wind variability, wind gusts, and sustained wind (one-minute average wind). Figure 1.14. Turbulent eddies bring higher wind speeds down to the surface. To reduce the effect of turbulent eddies on wind measurements, we measure the wind away from the surface, at a standard height of 10 m (30 ft). Even at this distance above the surface, the wind speed fluctuates significantly, and an instantaneous wind measurement is not very representative. Therefore, we average the wind speed over a period of time and report the wind as both sustained wind and gusts, where the gusts denote the higher values of wind speed that only last for short periods of time. The standard time average for sustained surface winds is one minute. In the example shown in Figure 1.13, the sustained wind is about 12 knots, and the wind would be reported as a 12 knot wind with 20 knot gusts. Note that the effect of friction on the wind varies with the underlying surface. For example, water surfaces are relatively smooth compared to land surfaces, where trees and buildings act as a drag on the air and slow it down. As a result, under similar weather conditions, sustained winds at the surface are typically 30–50% higher over water than over land (Figure 1.15). Figure 1.15. The ocean is a smooth surface that offers less resistance to the wind than land surfaces, resulting in higher wind speeds over the ocean. 1.3.3 Additional Sources of Wind Information Our understanding of atmospheric motions requires that we measure not only surface winds, but also upper-level winds. Meteorologists use radiosonde measurements to learn about the wind at various altitudes. Wind speed and direction are inferred from the balloon drift, which give us a vertical wind profile (see Figure 5.9 for an example). Measurements are also made routinely on commercial aircraft, which inform us about the weather along the main carrier routes. Finally, wind measurements can also be obtained from satellites, either by tracking cloud movement (speed and direction) or by using radar to measure the small ripples caused by the wind blowing on the surface of the ocean, which is an effect you can also observe on a lake: stronger winds make the surface “rougher.” 1.4 Precipitation The last element of weather that is of particular interest to us is precipitation, a generic term that encompasses rain, snow, hail, sleet, etc. Generally, we can define precipitation as solid or liquid water falling out of the atmosphere. We typically measure precipitation with a rain gauge, of which there are two primary types. In a tipping bucket, water is collected by a funnel and falls onto a tipping plate that tips when enough water has accumulated in the bucket (Figure 1.16). One tip usually corresponds to 0.25 mm (0.01 inches) of rain. The number of tips indicates the amount of rain that has fallen over time. Figure 1.16. Concept of the “tipping bucket” rain gauge. In a simpler type of rain gauge, a funnel also collects raindrops and magnifies the rain accumulation into a thinner cylinder so that the total depth may be accurately read on the scale (Figure 1.17). However, only the total depth of accumulated rain is read on the scale. In other words, we have no indication of the changes in precipitation rate between two readings. Such a rain gauge usually measures up to 25 mm, and the overflow is caught in the outside cylinder, for later measurements. Figure 1.17. Concept of the “funnel” rain gauge. Precipitation of less than 1 mm is reported as a trace. Frozen precipitation like snow is melted and its liquid depth is reported. (The frozen depth is also reported.) On weather maps, precipitation is indicated using specific symbols (see Figure 2.1 in Chapter 2). The most useful symbols are those for rain and snow. However, we will discuss very important weather conditions associated with sleet rain and fog (frozen raindrops in the form of ice pellets), freezing (liquid raindrops that freeze when they hit objects at the ground), (which is not precipitation). 1.5 Weather Stations All the instruments described above are set up in a standardized fashion at thousands of weather stations around the world. Most are now automated, as is the case for the Automated Weather Observing Systems (AWOS) of the Federal Aviation Administration and the Automated Surface Observing Systems (ASOS) of the US National Weather Service (Figure 1.18). At standard sites, thermometers are enclosed in a shelter to prevent direct exposure to the sun, wind, and precipitation. Nevertheless, the shelter needs to be well ventilated so that the measurements are representative of the surrounding air. That is why weather instrument shelters are painted white (to reflect sunlight) and have open slats to promote air circulation (Figure 1.19). Temperature is measured at a standard height of 2 m. Figure 1.18. Example of an automated weather station. Figure 1.19. Standard weather station shelter. Anemometers are set at the top of “masts” where they can measure wind speed and direction at the standard height of 10 m. Over the ocean, weather instruments are attached to masts on buoys, which are themselves anchored to the sea floor (Figure 1.20). Due to the difficulty and expense of installing and maintaining instruments at sea, there are far fewer buoys than there are land stations, as shown in Figure 1.21. When comparing the number of weather buoys at sea to the number of weather stations on land, it is immediately apparent that weather is much better directly monitored over land, and especially in urban areas, than over the ocean. (Note that some regions have chosen to set up denser networks of weather stations, one of which is apparent in Alberta, Canada.) This disparity has important implications, as weather observations are crucial for constraining forecasting models (see Chapter 13). However, the advent of satellite observations in the last few decades has very much reduced the impact of this disparity, as we will see in Chapter 2. Figure 1.20. Example of a weather buoy. Figure 1.21. Map showing the distribution of land-based stations and weather buoys in and around North America. Summary Atmospheric changes obey the laws of physics for the conservation of energy, mass, and momentum. We observe the atmosphere by measuring a number of variables constrained by these laws and relevant to explaining and predicting weather patterns. ► We measure temperature using thermometers and use three different temperature scales: Celsius, Fahrenheit, and Kelvin. ► We measure atmospheric pressure with barometers, and express it using hectopascals. ► We measure wind direction with wind vanes, and wind speed with anemometers. We report wind speed as sustained winds and wind gusts, using knots, miles per hour, meters per second, and kilometers per hour. ► We measure precipitation depth with “funnel” rain gauges with units of millimeters or inches, and precipitation rate with “tipping bucket” rain gauges. Finally, we assemble weather instruments at dedicated weather stations on land using shelters and masts, and buoys at sea. Weather observations are more dense over land than over water. We can use these measurements to create time series of the variables (i.e., variations in time) and weather maps (i.e., variations in space at a given time, as we will see in Chapter 2). Instruments are also attached to a weather balloon, a radiosonde, which records the corresponding variables as the balloon ascends, building a vertical profile of the atmosphere above a given location (i.e., variations with height). CHAPTER 2 Spatial Representations of Weather Data CONTENTS 2.1 The Station Model 2.2 Surface Maps 2.3 Upper-Level Maps 2.4 Radar 2.5 Satellites Summary Local weather is largely the result of large weather systems in motion. Thus, meteorologists gain insight into the manifestations of weather by studying maps and images of the atmosphere on a number of scales, from global, to regional, down to local scales. In this chapter, we will learn how weather information is represented on weather maps and images to reveal the twodimensional dynamics of weather systems. In Chapter 1 we introduced some foundational concepts to initiate a study of the atmosphere guided by scientific principles and the laws of physics and, in particular, conservation laws that constrain the behavior of particular variables: temperature, pressure, wind, and water. Understanding weather, and therefore the atmosphere in motion, now requires that we display these variables on maps, to reveal their spatial distribution. Furthermore, because the analysis of raw observations is often quite challenging, we need to design tools that will enable us to analyze these variables in a way that reveals salient features of the weather. The spatial representation and analysis of data will be the object of this chapter, before we return to each element of weather individually for a more thorough exploration in later chapters. 2.1 The Station Model The variables described in Chapter 1 and measured at weather stations or buoys are typically displayed on weather maps using the station model partly described in Box 2.1. Pressure, temperature, and dew point temperature are displayed as numerical values, while we use symbols for precipitation. The central circle indicates cloud cover. Wind is displayed with wind barbs using the meteorological convention. Wind direction is represented with a line extending from the weather station toward the direction the wind is coming from, while wind speed is indicated with barbs and pennants. Box 2.1. A (Simplified) Weather Station Model Weather observations are reported on surface maps using the weather station model shown in Figure 2.1.1.* Figure 2.1.1. Simplified weather station model. Cloud cover is indicated by shading the circle, as shown in Figure 2.1.2. Figure 2.1.2. Some cloud cover symbols. Significant weather by type and intensity is indicated on weather maps using symbols such as those shown in Figure 2.1.3. Figure 2.1.3. Some weather symbols. Wind direction is represented by a line extending from the weather station toward the direction the wind is coming from, while wind speed is indicated with barbs and pennants, as shown in Figure 2.1.4. Figure 2.1.4. Wind convention. Strictly speaking, the barbs and pennants indicate a range of wind speeds around the central value. For example, a half barb indicates 3 to 7 knots. Hence the values shown in Figure 2.1.5. Figure 2.1.5. Reporting wind speed with barbs and pennants. * See http://www.wpc.ncep.noaa.gov/dailywxmap/wxsymbols.html for a complete description. An example of a surface weather map showing measurements at land and buoy stations on February 20, 2014, is shown in Figure 2.1 (only temperature, dew point temperature, pressure, and wind are shown, for clarity). This is our first look at a weather event that will follow us through the rest of this book: an extratropical cyclone generated east of the Rocky Mountains and sweeping through the eastern half of the United States between February 19 and 22, 2014. Extratropical cyclones are also called midlatitude storms, or frontal cyclones, as cyclones form fronts, but fronts can also form cyclones. Figure 2.1. Example of surface station map on February 20, 2014, at 18:00 UTC. To facilitate reading, temperature is indicated in red, dew point temperature in blue, and sea-level pressure in black, in addition to following the weather station model layout described in Box 2.1. At this stage of its development, the storm (let’s call it “our February 2014 cyclone”) appears in our weather map as a counterclockwise circulation occurring over the eastern half of the United States. By following the direction of the wind barbs, we can get a sense of its size and locate its center somewhere along the border between Iowa and Missouri (see Appendix 2.2 for geographical reference). By contrast, the western United States is relatively quiet with low wind speeds and no organized circulation. 2.2 Surface Maps As evident in Figure 2.1, it is quite difficult to appreciate the spatial distribution of the different variables at a glance. That is why we analyze and complement the surface maps with additional elements. 2.2.1 Isotherms and Temperature Maps We first color the temperature field using an intuitive color scale, typically going from blue (low temperatures) to red (high temperatures), as shown in Figure 2.2. We further contour the temperature observations with contour lines at regular intervals. These contours are called isotherms, from “iso,” meaning “equal,” and “therm”, meaning “heat.” Isotherms are lines of constant temperature. Indeed, if we were to walk along an isotherm with a thermometer, the temperature would remain constant. If we decided to depart from the isotherm, the temperature would either increase or decrease. Figure 2.2. Example of a surface temperature map on February 20, 2014 at 18:00 UTC. Coloring and contouring the temperature map allows us to quickly identify the air masses, where the temperature is relatively uniform. In Figure 2.2, for example, a warm air mass occupies the southeastern sector of the United States, while a cold air mass occupies the northwestern half. In between these air masses are areas of strong temperature change, where the temperature increases from cold to warm over relatively short distances. These regions of sharp temperature change are called temperature fronts – see Section 2.2.2. The rate of change of a variable over a certain distance is called a gradient – a notion we will use often in the rest of this book. In Figure 2.2, for example, we can see that the temperature gradient is relatively small in the cold and warm air masses, i.e., if we drove across an air mass with a thermometer, for example from Austin, Texas, to Atlanta, Georgia, a 1500 km drive, the temperature would not vary much. Frontal regions, however, are, by definition, regions of strong temperature gradient, i.e., if we drove across a temperature front, our thermometer would record a strong temperature increase or decrease over a short distance. Going from Kansas to Arkansas, for example, the temperature increases by 30 °F over only 200 miles. temperature increases by 2.2.2 Temperature Fronts Weather in the midlatitudes is largely the result of air masses in motion. In fact, by matching the air masses visible in Figure 2.2 with the corresponding wind directions in Figure 2.1, you can see that the warm air mass over the southeastern United States is advancing northward, toward colder air, while the cold air mass over the Great Plains is advancing to the southeast, toward the warm air mass. The temperature fronts visible in between are largely the result of these air masses in motion. As we will see in Chapter 10, clouds and precipitation tend to be concentrated along the temperature fronts, i.e., along the boundaries between the air masses. Therefore, we indicate the leading edge of air masses on weather maps, using a blue line and blue triangles for cold fronts, and a red line and red semi-circles for warm fronts, as shown in Figure 2.3. Note that, in each case, the symbols are pointing in the direction of motion. In Figure 2.3(a), the cold air mass is replacing the warm air to the southeast, and therefore the blue triangles point to the southeast. Similarly in Figure 2.3(b), the warm air replaces the cold air to the northeast, and therefore the red semi-circles point to the northeast. [For future reference, know that fronts are drawn on the warm side of the largest magnitude in the temperature gradient.] Figure 2.3. Conventional representation of (a) a cold front and (b) a warm front on weather maps. We will learn more about air masses and fronts in Chapter 10. In the meantime, we will soon add cold and warm fronts to our surface map, but first we need to introduce pressure contours, or isobars, which will help identify the fronts. 2.2.3 Isobars and Pressure Maps As for temperature, when studying isolated point measurements of pressure on a surface weather map, it is quite challenging to capture visually at once the entire structure of the pressure field, with its minimum and maximum values, its areas of rapid horizontal change as opposed to slow change, etc. That is why we also contour the pressure observations, to obtain a twodimensional picture of the pressure field (Figure 2.4). Note that pressure measurements are adjusted to mean sea level before contouring (Box 2.2). Contours of pressure are called isobars, and each isobar represents a particular value of pressure – typically at a 4 hPa interval. We can talk, for example, about the 1000 hPa isobar. By definition, the pressure will be 1000 hPa everywhere along that contour. Figure 2.4. Example of surface pressure map on February 20, 2014, at 18:00 UTC. Box 2.2. Adjustment of Pressure to Mean Sea Level Pressure maps are an important tool of the weather analyst, as pressure is intimately related to wind and weather systems. We will learn in Chapter 8 that wind is caused by horizontal pressure variations. Pressure, however, decreases far more rapidly upward than it changes horizontally. A good rule of thumb is that pressure decreases by about 8 hPa with each 60 m gain in altitude (at low elevation). Thus, the pressure at the top of the Empire State Building (381 m) is about 50 hPa lower than at street level. The pressure at the top of the Rocky Mountains could be as low as 600 hPa. However, the horizontal pressure variations causing wind are much more subtle than this. Recall that the typical range of pressure at sea level is about 980 to 1030 hPa. Therefore, if we reported actual pressure measurements on a surface weather map, as they are measured at weather stations on the ground, we would observe very low pressure values wherever there are mountains. We would suggest the presence of artificially low pressure areas where the mountains stand, with a drop in pressure of 300 or 400 hPa in places. To avoid such artefacts, we calculate what the pressure would be if there were no mountains and we were at sea level. We replace the mountains with air, so to speak, and add the corresponding pressure to the value measured at the weather station. In other words, we adjust the values of pressure to mean sea level. Figure 2.2.1(b) shows the unadjusted pressure field, and a comparison with the topography (Figure 2.2.1(c)) confirms that the low pressure areas correspond to the Rocky Mountains and the Appalachians, at the expense of our February 2014 extratropical cyclone over the Midwest. After adjusting the pressure measurements to mean sea level (Figure 2.2.1(a)), the low pressure center and the troughs corresponding to the cold and warm fronts appear clearly. [Note that we have used a green color scale for the purpose of this exercise, although pressure fields are usually displayed with isobars alone.] Figure 2.2.1. Pressure measurements adjusted to mean sea level (a), with the unadjusted pressure field (b) and topography (c) for comparison. A good analogy for isobars is the lines of constant height on a topographic map. You know that if you are hiking along such a line, you stay at the same height. If you depart from that contour, however, you either climb up toward higher elevations or down toward lower elevations. Similarly on a pressure map, pressure necessarily increases on one side of the isobar and decreases on the other side. Furthermore, if topographic contours are tightly packed together, you know that you are facing a steep slope. Similarly, tight isobars indicate a steep pressure change, i.e., a sharp pressure gradient, and we will learn in Chapter 8 that strong pressure gradients induce strong winds. 2.2.4 Highs, Lows, Ridges, and Troughs The surface pressure field often contains regions of minimum and maximum pressure. Figure 2.5(a) shows an example of a high pressure region where the pressure increases toward a maximum value, which we indicate with a big H (for “high”). We will see that high pressure regions, or “highs,” correspond to anticyclones and are often associated with fair weather (i.e., weak wind and clear skies). In Figure 2.4 the western half of the United States is under a broad area of high pressure. Figure 2.5. Example of (a) high pressure region, or “high,” and (b) low pressure center, or “low,” as would appear on a surface pressure map. Figure 2.5(b) shows an example of a low pressure region in which the pressure decreases to a central minimum, indicated with a big L (for “low”). Note that there is no 988 hPa contour. Therefore, we know that the pressure is above 988 hPa everywhere inside the interior contour. The big L, however, tells us that there is a minimum value less than 992 hPa. The exact value of the minimum pressure is often indicated on standard weather maps. We will see that such low pressure centers correspond to cyclonic weather systems, such as midlatitude storms and hurricanes. They are often associated with active weather, such as clouds, precipitation, and strong wind. We will refer to such minima in pressure as “low pressure areas”, “low pressure centers,” or simply as “lows.” Our February 2014 cyclone, for example, as depicted in Figure 2.4, corresponds to a low pressure center located over northern Missouri. To continue with the topography analogy, it is sometimes useful to think of a high pressure region as a mountain of high pressure, and to think of a low as a basin. For practice, try to convince yourself that you could hike up to the top of the mountain shown in Figure 2.5(a) and down into the bowl shown in Figure 2.5(b). Similarly, if the contours are not closed, we can still find useful features to identify in the pressure field. For example, if higher pressure values extend in one direction, this is like a mountain ridge, as indicated in Figure 2.6(a). Even though there is no isolated high pressure center, the crest of the ridge is still a line of relative maximum pressure, compared to the lower values of pressure down the slopes on each side of the crest. We will refer to this feature as a “ridge of high pressure,” and to the line of relative maximum pressure as the “axis” of the ridge. Figure 2.6. Example of (a) high pressure ridge, or “ridge,” and (b) low pressure trough, or “trough,” as would appear on a surface pressure map. We will see that ridges are especially apparent on upper-level maps, and are associated with fewer clouds and lighter winds. Conversely, when the pressure field is relatively lower along a line, we have the equivalent of a valley. Although we do not call it a “valley of pressure,” we use a similar analogy and call it a “trough of low pressure.” Even though there is not a closed low pressure center, the axis of the trough is still a line of relative minimum pressure (Figure 2.6(b)). We will see that troughs are also frequently observed on upper-level maps and are often associated with regions of clouds and precipitation. Note that we rarely color the pressure field as we did for temperature, because it is more informative to overlay the isobars on top of the (colored) temperature field, as shown in Figure 2.7. We will see that this is a powerful tool in meteorology, as it allows us to analyze temperature and pressure jointly. After adding the highs (H), the lows (L), and the temperature fronts, we obtain a compelling picture of the state of the atmosphere at any given time. To convince yourself fully, return to Figure 2.1 and compare it to Figure 2.7 to appreciate how much more readable our surface map has become. At a glance, we have what we call a synoptic view of the atmosphere, or a view of the “synoptic situation” over the United States (from the Greek “syn” meaning “together” and “opsis” meaning “view”). Figure 2.7. Example of surface weather map showing temperature (colors and isotherms), pressure (isobars), highs (H) and lows (L), as well as temperature fronts on February 20, 2014, at 18:00 UTC. Here we can see that the cold and warm air masses, the cold and warm fronts, and the low pressure center all coincide to form our February 2014 extratropical cyclone. We will see in Chapter 10 that the air masses and fronts circle around the low, and the warm sector, i.e., the sector of warm air delimited by the two fronts, decreases in size as the cyclone evolves, while clouds and rain form along the fronts and at the center of the cyclone. By looking carefully at the pressure field, you will also notice that the fronts happen to be located where the isobars “kink” and change directions. If you think of these two depressed areas as pressure troughs, then the fronts follow the axis of the troughs. This is indeed a useful tip for placing temperature fronts on a surface map, the reason for which we will discover when we know a bit more about wind. 2.3 Upper-Level Maps We tend to think of weather as happening close to Earth’s surface, where we live and where we can observe it. But most of the action actually takes place aloft. Surface weather in the midlatitudes is largely determined by the position and structure of the jet stream, a fast-flowing air current blowing at the top of the troposphere. Disturbances in the jet stream create certain conditions that are more or less conducive to the formation, or demise, of cyclones. That is why meteorologists spend a lot of time studying upper-level maps. We could create such maps by contouring pressure observations at specified heights, as we did for our surface weather map. For example, we could collect pressure measurements at an altitude of 5000 m and build a 5000 meter pressure map. We would then find highs and lows, troughs and ridges, as we did at the surface. In practice, however, it is more convenient to do the opposite: we choose a particular pressure, for example, 500 hPa, and we indicate how high in the atmosphere this 500 hPa value is located. If we now contour these height measurements, we have a 500 hPa “height” map. We can think of it as a two-dimensional surface where the pressure is everywhere 500 hPa, roughly at 5.5 km altitude, but with dips and bumps that make it sag lower or rise higher in places. This seems complicated, but in practice it yields about the same result, and it is more convenient from a mathematical point of view, because working at constant pressure allows us to simplify our equations. [Specifically, the air density disappears from the equations. Recall that density is not measured, but estimated indirectly from other measurements. Therefore, disposing of density is very convenient.] In the end, we can treat this “height” map exactly as we would a pressure map, as illustrated in Figure 2.8. If you picture a series of pressure surfaces at different altitudes, they look like the blue lines shown in Figure 2.8(a) on a cross-section. Recall that pressure decreases with height, and therefore the 850 hPa surface is necessarily above the 900 hPa surface and below the 800 hPa surface. If you follow the 850 hPa surface from west to east, you will see that it is higher to the west (blue H) and lower to the east (blue L). If you now follow the 1.5 km height line (dashed) from west to east and read the corresponding pressure (on the blue lines), you will see that the pressure is higher to the west (about 900 hPa, indicated by a black H) and lower to the east (about 800 hPa, indicated by a black L). In other words, the height surface is high (H) when the pressure is high (H), and it is low (L) when the pressure is low (L). On the actual 850 hPa map, the high and the low might look like Figure 2.8(b), which tells us that the 850 hPa surface is sloping from higher to lower altitudes from H to L. The high (at the constant pressure of 850 hPa) also corresponds to higher pressures (at the constant height of 1.5 km), while the low corresponds to lower pressures. Therefore, for our purposes, we can think about the features on the upper-level height maps as if they were on a pressure map. Figure 2.8. Example of a high and a low as would appear on (a) a crosssection, and (b) a 850 hPa map. The blue lines in panel (a) represent pressure surfaces, while the black lines in panel (b) represent isobars. The blue H represents a region where the 850 hPa surface is high, while the black H represents the corresponding region where the pressure is high at 1.5 km altitude (and similarly for low pressures). Meteorologists frequently use the 850, 500, and 300 hPa pressure surfaces for upper-level analyses. In addition to height contours, we usually plot station data on these maps, as shown in Figure 2.9 for a 500 hPa map, which helps us to visualize the speed of air currents flowing along the contours (see Chapter 8). Finally, recall from the ideal gas law that, for a fixed mass of air, warmer air has a greater volume than colder air. Therefore, for a fixed area, a warmer column of air is “thicker” and associated with higher upper-level heights on pressure surfaces. Thus, meteorologists can also analyze height fields in terms of thickness and heat content of the underlying air columns. Figure 2.9. Example of station data on a 500 hPa upper-level map. Height given in meters. An example of a simplified upper-level map, corresponding to our February 2014 cyclone, is shown in Figure 2.10, where wind speed is also indicated with colors in the background. Closed highs and lows are not so common at 500 hPa, but troughs and ridges are, and often correspond to specific weather features at the surface. Upper-level ridges, for example, often indicate fewer clouds, while upper-level troughs create surface conditions that are conducive to the formation of cyclones (see Chapter 10). Here, the Great Lakes are under a ridge of high pressure, but the Midwest is under a trough of low pressure, with strong winds blowing around the trough, and particularly high winds on the southeastern flank of the trough, over Missouri. The alert reader will notice that this feature of the upper-level jet coincides with the location of the surface low in Figure 2.7. Indeed, we will learn in Chapter 10 that this configuration is conducive to the deepening of extratropical cyclones. Figure 2.10. Example of a 500 hPa map on February 20, 2014, at 18:00 UTC. 2.4 Radar Radars (originally RAdio Detection And Ranging when developed for military purposes) are very useful for measuring and tracking precipitation. They emit radio waves nearly horizontally, since the troposphere is so thin. Some of these waves are transmitted through the atmosphere and are lost if they do not encounter any obstacle. If precipitation is present on their path, however, the raindrops will reflect some of the waves back to the radar, where the returned energy can be measured and analyzed (Figure 2.11). Figure 2.11. Concept of radar as used to measure precipitation. The black lines show the signal emitted by the radar. The red lines show the part of the signal that returns to the radar after being reflected by raindrops. The return signal is an indirect measurement of precipitation. It contains two pieces of information. First, more precipitation causes more reflected signal. Therefore, the radar can be used to estimate the rate of precipitation. Second, by calculating the time it took for the signal to reach the raindrops and come back, we can estimate the distance between the radar and the precipitation, and we can indeed map the structure of the precipitation around the radar. Therefore, a radar provides information about both the location and the intensity of precipitation. By tilting the radar up and down, we can also learn about the vertical structure of the precipitation. By rotating the radar horizontally, we can build a 360º picture of precipitation all around the radar, out to a distance of about 250 km. By aggregating measurements from a network of radars, we can build a composite picture for large areas, as shown in Figure 2.12 for our February 2014 cyclone at a later stage of its development. Recall that precipitation tends to concentrate along the temperature fronts and in the low pressure center. On February 21, the cyclone center had moved over Wisconsin, and the bands of precipitation can be seen to delineate the cold front (from East Texas to Illinois), the warm front (from Michigan to New York), and what we will call the occluded front (arching back over Wisconsin). Figure 2.12. Example of a radar image of precipitation over the United States on February 21, 2014, at 00:00 UTC. A more advanced type of radar, called Doppler radar, can also detect the shift in frequency of the return signal, as compared to the emitted signal, caused by the horizontal motion of the raindrops. This shift in frequency can be translated into a speed, which informs us about the wind – specifically, whether the raindrops are moving toward or away from the radar. Finally, the latest radars have the ability to utilize polarized radar signals, i.e., radio waves that are oriented in the vertical and horizontal plane, to determine the shape of precipitation particles. 2.5 Satellites Since the first weather satellite was placed in space in 1960 – TIROS (Television Infrared Observation Satellite) – satellites have revolutionized our view and understanding of the atmosphere and weather. They provide information about cloud coverage (cloud location and height), as well as information about the structure of weather systems. Importantly, they provide information both over land and over the ocean, where very few direct measurements are available. (Recall that the ocean covers two thirds of the planet.) Weather satellites carry sensors that measure different kinds of radiation. These measurements can then be translated into variables of interest for weather and climate (temperature, cloud coverage, cloud top height, cloud motion, wind, etc.). For our purposes, we will be interested in three important types of satellite sensor: visible, infrared, and water vapor. 2.5.1 Visible Satellite Images A satellite can be equipped with a visible light sensor. Since Earth does not emit visible radiation, the only light that will be measured by such a sensor is visible light from the sun reflected by bright features on Earth, such as clouds, ice, snow, and to some extent bright land areas. Figure 2.13 shows a visible image of our February 2014 cyclone captured by GOES-East, an American geostationary satellite (see Section 2.5.4). Later on February 21, the cyclone was starting to decay, as we will explain in Chapter 10. At this fully developed stage, however, the visible image reveals a beautiful display of clouds spanning thousands of miles. The cold front has now moved further east and extends from the Gulf of Mexico, along the East Coast, to New England. The extensive cloud cover over the low pressure area suggests precipitation over Canada, and it was indeed snowing heavily over the Great Lakes. As for the warm front, it extends over the Atlantic Ocean and connects with another cyclone in its decaying phase over Europe. Figure 2.13. Example of a visible satellite image on February 21, 2014, at 17:45 UTC. For orientation, the United States and Mexico appear in light gray in the western half of the image. Visible images are useful for identifying and tracking cloud features when they are illuminated by the sun. As Earth and the satellite rotate away from the sun to the night side, however, no more sunlight is reflected and no observation is possible (Figure 2.14). Therefore, we resort to a second type of sensor to obtain information about Earth independently of the position of the sun. Figure 2.14. Geostationary satellites can only take visible images during the day. A few hours later, both Earth and the satellite have spun away to the night side. 2.5.2 Infrared Satellite Images As we will see in Chapter 4, all objects emit some level of radiation, and the warmer the object, the more radiation it emits and at a higher frequency. The sun, for example, emits a large amount of radiation, in the visible part of the spectrum. Most objects on Earth are not as hot, however, and therefore emit a lesser amount of radiation, in the infrared (IR) part of the spectrum – radiation of lower frequency than visible light. Now, the infrared spectrum encompasses a broad range of frequencies and slight variations in the temperature of Earth features will lead to slight variations in the amount and frequency of the infrared radiation they emit. An infrared sensor on a satellite can be tuned to detect such variations in radiation, and therefore variations in temperature. Thus, the full picture taken by an infrared sensor, i.e., an infrared satellite image, is really a “temperature image” of Earth. It is relatively straightforward to display a visible satellite image, since it shows in shades of gray what our eyes would see if we were sitting on that satellite looking down at Earth. A visible satellite image is really a black and white photograph, as can be seen in Figure 2.14 – or, lately, a color photograph, with the latest full-color satellite imagery. But how do we display an infrared satellite image? Our eyes have evolved to see visible light, but not infrared radiation. Therefore, there is no such thing as “infrared colors.” So we need to create a “fake” color scale to display the temperature information contained in the image (Figure 2.15). We display warm features with darker shades of gray and colder features with lighter shades of gray (Figure 2.16). Of course, the scale was not chosen randomly: it reflects the fact that cloud tops, being higher in the atmosphere, are colder than the surface. Recall that temperature decreases with height. Since it is more intuitive to represent cloud tops in white, we chose our scale to show cold as white, and warm as dark gray. Figure 2.15. Example of an infrared satellite image on February 21, 2014, at 17:45 UTC, corresponding to Figure 2.13. Figure 2.16. Color scale for displaying satellite infrared images. At first glance, the white features in Figure 2.15 indeed seem to match the clouds of Figure 2.13. After closer inspection, however, we notice that some of the clouds present in the visible image are absent from the infrared image. Indeed, clouds are not all found at the same altitude, and therefore do not all have the same temperature. Clouds with a light gray shading in an infrared satellite image are colder and therefore higher in the atmosphere, while clouds with a dark gray shading are warmer and therefore lower in the atmosphere (Figure 2.17). [Note that, strictly speaking, what we are seeing are the cloud tops. Therefore, a cloud with a light gray shading can be either a high cloud or a tall cloud.] Figure 2.17. The amount of infrared radiation emitted by cloud tops tells us their temperature, which in turn tells us their altitude. (Note that “light gray” and “dark gray” here refer to the appearance of cloud tops on an infrared image, not to the color of the clouds themselves.) Figure 2.15 tells us that, at this later stage of its development, our February 2014 extratropical cyclone contains different types of clouds, in contrast to what seems to be fairly homogeneous cloud cover in the visible image. This is very instructive to the weather analyst, as it reveals different aspects of the circulation pattern. The cloud band along the cold front, for example, is bright white and reaches high in the atmosphere. In contrast, the cloud spiral that extends over Canada and arches back toward the Great Lakes, producing heavy snow, is darker in color, and therefore lower in the atmosphere. In the end, we should never forget that an infrared image shows us temperature, rather than the brightness of the clouds. This explains, for example, why some clouds over the Atlantic Ocean appear in the visible image shown in Figure 2.13, but not in the infrared image shown in Figure 2.15. They are clouds at very low altitude. They are made of liquid cloud droplets and highly reflective; therefore, they clearly appear on a visible image. However, being low in the atmosphere, and therefore at temperatures close to the temperature of the ocean underneath, they are hardly distinguishable from the ocean itself. If we were not aware of the true nature of an infrared image, we could simply overlook the existence of these clouds. 2.5.3 Water Vapor Images Our satellite sensor can also be tuned to detect radiation emitted by specific molecules in the atmosphere. As we will see in Chapter 4, gas molecules absorb and emit radiation at specific wavelengths due to their molecular structure. A gas of particular interest to us is water vapor, as it transports energy through phase changes and plays a major role in the development of weather systems (see Chapter 5). Therefore, we also equip our satellite with a water vapor sensor to produce images such as shown in Figure 2.18. Figure 2.18. Example of a water vapor satellite image on February 21, 2014, at 17:45 UTC, corresponding to Figure 2.13. In this image, darker colors indicate a drier troposphere, whereas lighter colors indicate a higher water vapor content. As for infrared imagery, the altitude of the atmospheric water vapor is a factor in the amount of radiation recorded by the satellite. In particular, dark and light colors are more indicative of the presence or absence of water vapor in the mid to upper troposphere. In other words, a dark region on a water vapor satellite image might still contain a fair amount of water vapor in the lower troposphere, and even low clouds. Note, for example, the presence of numerous low cumulus clouds over the tropical Atlantic in Figure 2.13, even though it is depicted as “dry” in Figure 2.18. Water vapor satellite imagery is particularly useful to weather analysts when it is animated, as it shows the fluxes of water vapor in and out of weather systems, as well as the intrusions of dry air. Note, for example, the tongue of dry upper-tropospheric air wrapping around and into the cyclone in Figure 2.18 (west of the cold front). This highlights another important attribute of the water vapor image: it shows us structure in the atmosphere even where there are no clouds. 2.5.4 Geostationary Satellites Meteorologists rely primarily on two types of satellite: geostationary and polar-orbiting. As their name indicates, geostationary satellites are stationary with respect to Earth (“geo-”). More precisely, they are stationary with respect to a specific location on Earth, so that they constantly observe that same location over time. Since Earth rotates, the satellite needs to move at the exact same speed, otherwise it will start leading or lagging behind the location it is supposed to observe. More exactly, it needs to orbit Earth at the same angular speed, i.e., it needs to describe the same angle over time. At that speed, the satellite will experience a centrifugal force that will tend to pull it out of its orbit. This is similar to the outward pull you experience when you turn at high speed with your car. For the satellite to stay in orbit, the gravitational pull of Earth needs to balance the centrifugal force exactly. This perfect combination happens in the equatorial plane at about 36 000 km from Earth’s center (about 22 000 miles), which defines the geostationary orbit (see Figure 2.19). For reference, Earth’s radius is about 6400 km. So, a geostationary satellite is at about five to six times that distance from Earth, above the equator. At this location, it can hover above a point on Earth, without falling toward Earth. Figure 2.19. Geostationary satellites orbit Earth at the same angular speed as Earth. Since a geostationary satellite always remains above the same location, it is useful for tracking weather changes over a specific country or region. In the United States, for example, one satellite called GOES-West (Geostationary Operational Environmental Satellite-West, currently GOES15...
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