Tropical Circulation Systems and Monsoons
Kshudiram Saha
Tropical Circulation Systems and Monsoons
123
Dr. Kshudiram Saha 4008 Beechwood Road University Park MD 20782 USA
[email protected]
ISBN 978-3-642-03372-8 e-ISBN 978-3-642-03373-5 DOI 10.1007/978-3-642-03373-5 Springer Heidelberg Dordrecht London New York Library of Congress Control Number: 2009935567 © Springer-Verlag Berlin Heidelberg 2010 This work is subject to copyright. All rights are reserved, whether the whole or part of the material is concerned, specifically the rights of translation, reprinting, reuse of illustrations, recitation, broadcasting, reproduction on microfilm or in any other way, and storage in data banks. Duplication of this publication or parts thereof is permitted only under the provisions of the German Copyright Law of September 9, 1965, in its current version, and permission for use must always be obtained from Springer. Violations are liable to prosecution under the German Copyright Law. The use of general descriptive names, registered names, trademarks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use. Cover design: deblik, Berlin Printed on acid-free paper Springer is part of Springer Science+Business Media (www.springer.com)
Dedicated with gratitude to my beloved wife, Usha
Preface
The present book is a sequel to the author’s first book entitled ‘The Earth’s Atmosphere – Its Physics and Dynamics’, published in 2008, and addresses the practical side of the subject with stress on proper analysis and interpretation of tropical data. It deals specifically with circulation systems that are excited by heat sources and sinks in the Earth-atmosphere system over the tropics and extratropics. Together, the two volumes are intended to form a fairly comprehensive course of study of tropical meteorology, in theory and practice. A few new ideas, especially on monsoons, though empirically derived from data and analysis, would appear to be intuitionally born out of nowhere and are put forward by the author in the hope that they will be verified by practicing meteorologists in the years to come. The book is divided into three parts: the first deals with some general aspects of tropical general circulations and their disturbances; the second with monsoons over seven regions of the tropics, and the third with monsoons over Central America and the extratropical belt of North America. Part I, in dealing with the General circulation of the atmosphere in the tropics and its large-scale perturbations, introduces the Hadley, Walker and Monsoon circulations in a unified system. It also introduces the synoptic-scale tropical disturbances, such as Easterly waves, monsoon troughs, depressions and cyclones and gives a detailed account of their formation, structure, development and movement over different parts of the globe. Part II is devoted wholly to Monsoons in different parts of the tropics. The regions covered are Southern Asia, Eastern Asia, Maritime Continent, Australia, North and South Africa, and South America. Part III deals with monsoon over Central America and adjoining Southwestern North America as well as monsoons over North America. Monsoon is one area where the book advances several new ideas and concepts. It offers an entirely new definition of monsoon which allows monsoon to be detected and identified over several hitherto-unsuspected regions of the globe. It unveils the wave structure of the monsoon and identifies the areas to be associated with organized clouding and rainfall. Among the other important issues discussed in the book are the origin of African wave disturbances which later in their life cycle develop into Atlantic hurricanes, Low-Level Jets over several parts of the globe, and monsoon lows and depressions over South America, and formation of tropical cyclones over the South Atlantic Ocean. vii
viii
Preface
Written in an easy-to-understand style, with minimal mathematics, the book is intended for use by students pursuing courses in atmospheric science at undergraduate and graduate levels in colleges and universities and by scientists of other disciplines interested in tropical meteorology. It is likely to appeal to a large circle of readers interested in studies of tropical monsoons. Practicing meteorologists may find it particularly useful and challenging while comparing the results of their analysis and diagnosis of tropical circulation systems with the empirical and conceptual models suggested in the book. Inspite of the utmost care taken in the preparation of the book, it is likely that there have been errors and omissions. The author will be thankful if such lapses are brought to his notice. University Park, Maryland June 15, 2009
Kshudiram Saha
Acknowledgments
The author is grateful to his family, especially his daughters Manjushri and Suranjana, for helping him throughout the preparation of this book. Manjushri typed the long list of references. Suranjana who is co-author of many of the papers referred to in the book helped not only with insertion of nearly 150 diagrams in the book, but also along with her husband Dr. Huug van den Dool provided all the logistic support and technical facilities. Without their wholehearted support, it would have been impossible to complete this work. Huug’s review of the first draft of most of the chapters of the book was very helpful in improving the manuscript. The author’s special thanks are due to the National Centers for Environmental Prediction (NCEP) of the National Weather Service, USA, for several of their data and analysis products incorporated in the book. He expresses his deep gratitude to the numerous authors, publishers and learned Societies, who permitted him to reproduce diagrams and excerpts from their published work. Finally, the author would like to thank all the publishing staff at Springer-Verlag, involved in the book project, for their unfailing courtesy, active cooperation, and helpful attitude.
ix
Contents
Part I
Tropical Circulation Systems – A Survey
1 Large-Scale Tropical Circulations – Some General Aspects . . . . 1.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.2 Tropical Circulation as Part of the Global General Circulation – The Tradewinds . . . . . . . . . . . . . . . . . 1.3 Poleward Boundary of the Tropical Circulation . . . . . . . . 1.4 Heat Sources and Sinks . . . . . . . . . . . . . . . . . . . . . 1.4.1 Definition of Heat Sources/Sinks . . . . . . . . . . . 1.4.2 Diabatic/Adiabatic Heat Sources/Sinks . . . . . . . . 1.5 Some Physical and Dynamical Constraints and Conservation Laws . . . . . . . . . . . . . . . . . . . . . 1.5.1 Direct and Indirect Circulations . . . . . . . . . . . . 1.5.2 Energy Transformations . . . . . . . . . . . . . . . . 1.5.3 Energy Transfer Process – Carnot’s Cycle . . . . . . 1.5.4 Conditional Instability and Convection . . . . . . . . 1.5.5 Cellular Structure – Shallow and Deep Convection . 1.5.6 Coriolis Control-Variation with Latitude . . . . . . . 1.5.7 Conservation Laws . . . . . . . . . . . . . . . . . . 1.6 Equatorial Circulations . . . . . . . . . . . . . . . . . . . . . 1.6.1 Circulation with Heat Sources and Sinks Placed Alternately Along the Equator – Walker Circulations 1.7 Meridional Circulation with Heat Source at the Equator and Heat Sinks in Higher Latitudes – The Hadley Circulations 1.8 Seasonal Migration of the Equatorial Heat Source . . . . . . . 1.8.1 Origin of Monsoon . . . . . . . . . . . . . . . . . . 1.8.2 The Wave Structure . . . . . . . . . . . . . . . . . . 1.8.3 Forcing for the Seasonal Movement of the Equatorial Heat Source . . . . . . . . . . . . . . . . 1.8.4 Intraseasonal Oscillation of Monsoon . . . . . . . . 1.9 Co-existence of Monsoon, Hadley and Walker Circulations – Inclined Troughs . . . . . . . . . . . . . . . . 1.10 Definition of Monsoon . . . . . . . . . . . . . . . . . . . . .
. .
3 3
. . . . .
4 6 7 7 8
. . . . . . . . .
9 9 10 11 11 12 12 13 14
.
14
. . . .
15 17 17 18
. .
19 21
. .
22 24 xi
xii
Contents
1.11
1.12
Global and Regional Distribution of Monsoons . . 1.11.1 Tropical Monsoons . . . . . . . . . . . . 1.11.2 Extratropical Monsoons . . . . . . . . . . 1.11.3 Zonal and Meridional Anomalies . . . . . Co-existence of Monsoon with Desert Circulation .
. . . . .
. . . . .
. . . . .
. . . . .
. . . . .
. . . . .
. . . . .
2 Tropical Disturbances (Quasi-stationary Waves, Easterly/Westerly Waves, Lows and Depressions, Cyclonic Storms, and Meso-Scale Disturbances) . . . . . . . . . . . . . . . . 2.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2 Quasi-stationary Waves . . . . . . . . . . . . . . . . . . . . . . 2.2.1 Quasi-stationary Wave in Wind Field . . . . . . . . . . 2.2.2 Quasi-stationary Wave in Temperature Fields . . . . . 2.2.3 Structure of the Quasi-stationary Wave in Circulations . 2.3 Traveling Easterly (E’ly) Waves . . . . . . . . . . . . . . . . . 2.3.1 Easterly Waves in Tropical North Atlantic . . . . . . . 2.3.2 Easterly Waves in Tropical North Pacific . . . . . . . . 2.3.3 Easterly Waves in the Indian Ocean Region . . . . . . 2.4 Development of Waves . . . . . . . . . . . . . . . . . . . . . . 2.4.1 Meaning of Development . . . . . . . . . . . . . . . . 2.4.2 Development of a Quasi-stationary Monsoon Trough into a Depression . . . . . . . . . . . . . . . . 2.4.3 Development of a Depression into a Cyclonic Storm/Tropical Cyclone/Hurricane/Typhoon . . . . . . 2.5 Meso-Scale Disturbances and Severe Local Storms in the Tropics 2.5.1 General Considerations – Source of Energy of the Storm . . . . . . . . . . . . . . . . . . . . . . . . 2.5.2 Thunderstorms . . . . . . . . . . . . . . . . . . . . . 2.5.3 Hailstorms . . . . . . . . . . . . . . . . . . . . . . . . 2.5.4 Tornadoes . . . . . . . . . . . . . . . . . . . . . . . . 3 Tropical Cyclones/Hurricanes/Typhoons – Their Structure and Properties . . . . . . . . . . . . . . . . . . . . . . . . . 3.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . 3.2 Observed Structure of a Tropical Cyclone . . . . . . . . 3.2.1 Wind Structure . . . . . . . . . . . . . . . . . 3.2.2 Radial and Tangential Components of the Wind 3.2.3 Vertical Motion in a Mean Typhoon . . . . . . 3.2.4 Pressure Distribution . . . . . . . . . . . . . . 3.2.5 Temperature Distribution . . . . . . . . . . . . 3.3 The Eye and the Eye-Wall . . . . . . . . . . . . . . . . 3.3.1 General Considerations – Formation of the Hurricane Eye . . . . . . . . . . . . . . . . . . 3.3.2 Circulation Inside the Hurricane Eye – Evidence of Meso-Scale Vortices . . . . 3.3.3 Concentric Multiple Eye-Walls . . . . . . . . .
. . . . . . . . .
. . . . . . . . .
. . . . . . . . .
25 25 26 27 29
33 33 34 35 36 37 39 39 40 41 43 43 45 48 50 50 51 55 57
. . . . . . . . .
61 61 61 63 64 67 68 69 71
. . . .
71
. . . . . . . .
73 75
Contents
3.4
3.5
3.6
Part II
xiii
Spiral Bands Around the Eye-Wall . . . . . . . 3.4.1 Structure . . . . . . . . . . . . . . . . 3.4.2 Origin and Direction of Propagation . Storm Surge . . . . . . . . . . . . . . . . . . 3.5.1 Introduction . . . . . . . . . . . . . . 3.5.2 Some General Aspects of Storm Surge 3.5.3 Mathematical Models of Storm Surge Prediction of Cyclone Track and Intensity . . . 3.6.1 Early Models – The Steering Concept 3.6.2 Current Models . . . . . . . . . . . .
. . . . . . . . . .
. . . . . . . . . .
. . . . . . . . . .
. . . . . . . . . .
. . . . . . . . . .
. . . . . . . . . .
. . . . . . . . . .
. . . . . . . . . .
. . . . . . . . . .
77 77 78 78 78 79 80 82 83 83
Tropical Monsoons over Continents and Oceans
4 Monsoon over Southern Asia (Comprising Pakistan, India, Bangladesh, Myanmar and Countries of Southeastern Asia) and Adjoining Indian Ocean (Region – I) . . . . . . . . . . . 4.1 Introduction – Physical Features and Climate . . . . . . . . . . 4.2 The Winter Season (December–February) . . . . . . . . . . . . 4.2.1 General Climatic Conditions . . . . . . . . . . . . . . 4.2.2 Disturbances of the Winter Season . . . . . . . . . . . 4.3 The Transition Season (March–May) . . . . . . . . . . . . . . 4.3.1 Western Disturbances . . . . . . . . . . . . . . . . . . 4.3.2 ‘Heat Lows’ over Land and ‘Cold Highs’ over Ocean . 4.3.3 Severe Local Storms . . . . . . . . . . . . . . . . . . 4.3.4 Developments over the Equatorial Indian Ocean . . . . 4.4 Advance of Summer Monsoon to the Indian Subcontinent – General Remarks . . . . . . . . . . . . . . . . . 4.4.1 Advance over the Indian Ocean (April–June) – Stage 1 4.4.2 Onset over the Indian Subcontinent (June–July) – Stage 2 . . . . . . . . . . . . . . . . . . 4.4.3 Advance to Western Himalayas (July–August) – Stage 3 4.4.4 Source of Moisture for Monsoon Rainfall . . . . . . . 4.5 Disturbances of the Summer Monsoon during the Onset Phase . 4.5.1 Onset Vortex over the Arabian Sea and the Bay of Bengal . . . . . . . . . . . . . . . . . . . . . . . . 4.5.2 Monsoon Depressions and Cyclonic Storms . . . . . . 4.5.3 Interaction of Monsoon with W’ly Waves . . . . . . . 4.6 Rainfall over the Indian Subcontinent during the Onset Phase . 4.7 Summer Monsoon – Withdrawal Phase (September–November) 4.7.1 Dates of Withdrawal of Monsoon . . . . . . . . . . . . 4.7.2 Retreating Monsoon Rain over Tamil Nadu . . . . . . 4.7.3 Disturbances of the Withdrawal Phase . . . . . . . . . 4.8 Variability of the Indian Summer Monsoon Rainfall . . . . . . . 4.8.1 Interannual Variability . . . . . . . . . . . . . . . . . 4.8.2 Factors Likely Responsible for Interannual Variability . 4.8.3 Intraseasonal Variability . . . . . . . . . . . . . . . . .
89 89 91 91 92 93 94 94 94 95 98 102 104 105 106 107 107 107 108 110 111 111 111 113 115 115 117 121
xiv
Contents
5 Monsoon over Eastern Asia (Including China, Japan, and Korea) and Adjoining Western Pacific Ocean . . . . . . . . . . . 5.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2 Physical Features and Climate . . . . . . . . . . . . . . . . . 5.3 The Winter Season over Eastern Asia (November–March) . . 5.3.1 Temperature, Pressure, and Wind . . . . . . . . . . . 5.3.2 Quasi-stationary Wave in Westerlies – Its Interaction with Traveling Waves – Cold Surges . . . 5.3.3 Winter Rainfall over Eastern Asia . . . . . . . . . . 5.4 Airmass Transformations and Cyclogenesis over the Oceans . 5.4.1 Cyclonic Disturbances over Eastern Asia and Neighboring Ocean . . . . . . . . . . . . . . . . 5.5 Transition Period (April) . . . . . . . . . . . . . . . . . . . . 5.5.1 Development of ‘Heat Low’ over Eastern Asia . . . . 5.6 Origin of Monsoon over Eastern Asia . . . . . . . . . . . . . 5.7 Seasonal March of the Summer Monsoon . . . . . . . . . . . 5.8 Stationary States and Jumps . . . . . . . . . . . . . . . . . . 5.9 Meteorological Developments Associated with the Jump to Central China . . . . . . . . . . . . . . . . . . . . . . . . 5.9.1 Tibetan Plateau Monsoon . . . . . . . . . . . . . . . 5.9.2 The Meiyu (Plum Rain) Front over China . . . . . . 5.10 Jump of East Asian Monsoon to Extratropical Latitudes . . . 5.10.1 Evidence of Jump in Climatological Fields . . . . . . 5.10.2 Zonal Anomaly in Seasonal Variations . . . . . . . . 5.10.3 Climatological Rainfall over Eastern Asia During July . . . . . . . . . . . . . . . . . . . . . . 5.11 Monsoon over Japan . . . . . . . . . . . . . . . . . . . . . . 5.11.1 Geographical Location and Climate . . . . . . . . . 5.11.2 The Baiu Front – Its Seasonal Movement and Activity 5.12 Monsoon over Korea . . . . . . . . . . . . . . . . . . . . . . 5.12.1 Historical Background . . . . . . . . . . . . . . . . 5.12.2 Physical Features and Climate . . . . . . . . . . . . 5.12.3 Winter Monsoon over Korea . . . . . . . . . . . . . 5.12.4 Summer Monsoon over Korea – Changma Season . . 5.12.5 Korea’s Climatic Zones (After McCune, 1941) . . . . 6 Meteorology of the Maritime Continent (Region – III) (Comprising Philippines, Indonesia and Equatorial Western Pacific Ocean) . . . . . . . . . . . . . . . . . . 6.1 Introduction . . . . . . . . . . . . . . . . . . . . . 6.2 Climate of the Maritime Continent . . . . . . . . . 6.2.1 Pressure . . . . . . . . . . . . . . . . . . 6.2.2 Temperature . . . . . . . . . . . . . . . . 6.2.3 Relative Humidity and Cloudiness . . . . 6.2.4 Rainfall . . . . . . . . . . . . . . . . . .
. . . . . . .
. . . . . . .
. . . . . . .
. . . . . . .
. . . . . . .
. . . . . . .
. . . . .
123 123 124 125 125
. . .
128 130 130
. . . . . .
130 133 134 134 136 137
. . . . . .
137 138 139 142 143 145
. . . . . . . . . .
146 147 147 147 150 150 150 150 151 152
. . . . . . .
155 155 156 156 157 158 158
Contents
6.3
xv
. . . . . .
. . . . . .
. . . . . .
. . . . . .
159 159 160 160 165 168
7 Monsoon over Australia (Region – IV) . . . . . . . . . . . . . 7.1 Introduction – Location and Physical Features . . . . . . . 7.2 Early Studies . . . . . . . . . . . . . . . . . . . . . . . . 7.3 Climate of Australia and Surrounding Oceans . . . . . . . 7.3.1 Ocean Surface Temperature (SST, C) . . . . . . . 7.3.2 Air Temperatures . . . . . . . . . . . . . . . . . 7.3.3 Atmospheric Pressure (Isobaric Height) . . . . . 7.3.4 Wind and Circulation . . . . . . . . . . . . . . . 7.4 Monsoon over Australia . . . . . . . . . . . . . . . . . . 7.4.1 Onset of Monsoon . . . . . . . . . . . . . . . . . 7.4.2 Co-existence of Monsoon and Hadley Circulations – Interhemispheric Movement . . . . 7.4.3 Summer Monsoon Rainfall over Australia . . . . 7.5 Annual Rainfall of Australia and Its Seasonal Variability . 7.5.1 Annual Rainfall . . . . . . . . . . . . . . . . . . 7.5.2 Seasonal Variability . . . . . . . . . . . . . . . . 7.6 Variability of Australian Rainfall with ENSO . . . . . . . 7.7 Tropical Disturbances in the Australian Region – Depressions and Cyclones . . . . . . . . . . . . . . . . . 7.8 Tropical-Midlatitude Interaction in the Australian Region . 7.8.1 Northerly and Southerly Bursters . . . . . . . . .
. . . . . . . . . .
. . . . . . . . . .
. . . . . . . . . .
171 171 172 173 173 174 176 177 182 182
. . . . . .
. . . . . .
. . . . . .
183 185 187 187 188 189
. . . . . . . . .
189 191 192
6.4 6.5
Factors Affecting the Climate of the Maritime Continent 6.3.1 Geographical Location and Topography . . . . 6.3.2 Ocean Currents . . . . . . . . . . . . . . . . . 6.3.3 Equatorial Trough, the ITCZ and Monsoons . . The Maritime Continent – A Heat Source . . . . . . . . The Maritime Continent and the ENSO . . . . . . . . .
8 Monsoon over Africa (Region – V) . . . . . . . . . . . . . . . . . 8.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.2 Physical Features and Environment . . . . . . . . . . . . . . 8.3 Climates of Africa and Surrounding Oceans . . . . . . . . . . 8.3.1 Sea Surface Temperature and Wind . . . . . . . . . . 8.3.2 Air Temperature . . . . . . . . . . . . . . . . . . . . 8.3.3 Isobaric Height (gpm) . . . . . . . . . . . . . . . . . 8.3.4 Wind and Circulation . . . . . . . . . . . . . . . . . 8.3.5 Rainfall over Africa . . . . . . . . . . . . . . . . . . 8.4 Equatorial Westerlies over Africa . . . . . . . . . . . . . . . 8.5 The Equatorial Trough over North Africa – Its Zonal Anomaly 8.6 Structure of the Circulation Associated with the Equatorial Trough . . . . . . . . . . . . . . . . . . . . . . . 8.6.1 Zonal Circulation . . . . . . . . . . . . . . . . . . . 8.6.2 Meridional Circulation . . . . . . . . . . . . . . . . 8.7 Origin of African Wave Disturbances . . . . . . . . . . . . . 8.7.1 Early Studies . . . . . . . . . . . . . . . . . . . . .
. . . . . . . . . . .
195 195 196 198 198 198 199 201 205 207 209
. . . . .
210 211 211 212 212
xvi
Contents
. . . .
213 217 219 219
. . . .
220
9 Monsoon over South America (Region – VI) . . . . . . . . . . . . . 9.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.2 Physical Features and Environment . . . . . . . . . . . . . . . 9.2.1 Physical Dimension of the Continent . . . . . . . . . . 9.2.2 Topography . . . . . . . . . . . . . . . . . . . . . . . 9.2.3 Oceanic Environment and Its Influence on Climate . . 9.3 Climatological Features . . . . . . . . . . . . . . . . . . . . . 9.3.1 Air Temperature and Pressure . . . . . . . . . . . . . . 9.3.2 Atmospheric Circulation – Monsoon . . . . . . . . . . 9.3.3 Co-existence of Monsoon and Hadley Circulations . . 9.3.4 Rainfall over South America . . . . . . . . . . . . . . 9.4 Quasi-stationary Waves and Their Associated Weather . . . . . 9.4.1 Weather Phenomena Related to the Northern Boundary 9.4.2 Weather Phenomena Associated with the Southern Boundary . . . . . . . . . . . . . . . . . . . 9.5 Tropical Disturbances over South America . . . . . . . . . . . 9.5.1 Types of Disturbances . . . . . . . . . . . . . . . . . . 9.5.2 Monsoon Lows and Depressions . . . . . . . . . . . . 9.5.3 Upper-Tropospheric Cyclonic Vortices . . . . . . . . . 9.6 A Tropical Cyclone over the South Atlantic Ocean . . . . . . . 9.6.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . 9.6.2 Formation of the Initial Vortex – Interaction with W’ly Waves . . . . . . . . . . . . . . . . . . . . 9.6.3 Structure, Movement and Development of the Vortex .
223 223 224 224 225 225 227 227 230 233 235 236 238
8.8 8.9
8.7.2 Influence of Midlatitude Forcing . . . . . . . . 8.7.3 Sudan – The Breeding Ground . . . . . . . . . 8.7.4 Role of Topography . . . . . . . . . . . . . . . Structure, Development and Movement of the Waves . . Interaction of South African Monsoon with Midlatitude Waves of the Southern Hemisphere . . . . . . . . . . .
. . . .
. . . .
. . . .
240 241 241 242 243 248 248 250 251
Part III Extratropical Monsoons 10
Monsoon over Central America and Adjoining Southwestern North America (Region – VII) . . . . . . . . . . . . . 10.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.2 Heat Sources and Sinks and Their Seasonal Movement . . . . . 10.3 The Climate of Central America and Adjoining North America . 10.3.1 Surface Temperatures and Winds . . . . . . . . . . . . 10.3.2 Upper Air Temperatures . . . . . . . . . . . . . . . . 10.3.3 Upper Air Height (gpm) . . . . . . . . . . . . . . . . . 10.3.4 Upper Air Wind Field and Circulation . . . . . . . . . 10.4 Rainfall over Central America and Adjoining Areas . . . . . . . 10.4.1 Annual Rainfall over Mexico . . . . . . . . . . . . . .
255 255 256 257 257 259 260 262 264 267
Contents
xvii
10.4.2 10.5
11
Source of Moisture for Rainfall over the Arizona-Sonoran Desert . . . . . . . . . . . . . . . . . Some Characteristic Features of Weather over Central America . 10.5.1 Weather Associated with W’ly Waves . . . . . . . . . 10.5.2 Weather Associated with ‘Northers’ . . . . . . . . . . 10.5.3 Land and Sea Breezes on the Pacific Coast of Mexico . 10.5.4 Temporales of the Caribbean Sea and the Gulf of Mexico . . . . . . . . . . . . . . . . . . . . . . . . 10.5.5 Hurricanes and Tropical Storms . . . . . . . . . . . .
Extratropical Monsoon over North America . . . . . . . . . . . . . 11.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11.2 Climatological Background of North American Monsoon . . . . 11.2.1 Physical Features of the Land . . . . . . . . . . . . . . 11.2.2 Semi-permanent High and Low Pressure Systems over Oceans . . . . . . . . . . . . . . . . . . 11.3 The Seasonal Movement of Heat Sources and Sinks . . . . . . . 11.4 Seasonal Circulations – Monsoons . . . . . . . . . . . . . . . . 11.4.1 The Winter Monsoon (December–February) . . . . . . 11.4.2 The Spring Transition Season (March–May) . . . . . . 11.4.3 The Summer Monsoon (June–August) . . . . . . . . . 11.4.4 The Autumn Transition Season (September–November) 11.5 Interaction of Monsoons with W’ly Wave Disturbances . . . . . 11.6 Some Characteristic Features of East Coast Monsoon . . . . . . 11.6.1 Seasonal Variations and Reversals . . . . . . . . . . . 11.6.2 Monsoonal Characteristics of the East Coast Region . . 11.7 Role of the Appalachian Mountain Range – Leeside Cyclogenesis – Northeast Storms . . . . . . . . . . . . . . . . 11.8 Interaction of Monsoon with Storms and Hurricanes . . . . . .
268 270 270 270 271 272 272 275 275 275 276 277 278 279 279 281 281 286 287 287 287 288 288 289
Appendix: Meanings of Uncommon Words/Terms Used in the Book . .
291
References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
299
Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
313
Guide to Numbering Figures and Equations
Figures are numbered serially chapterwise. For example, Fig. 5.12 is diagram No. 12 of Chapter 5. Equations are numbered serially, sectionwise and chapterwise. For example, Eq. (5.12.4) is Equation No. 4 in Section 12 of Chapter 5.
xix
Part I
Tropical Circulation Systems – A Survey
Chapter 1
Large-Scale Tropical Circulations – Some General Aspects
1.1 Introduction Meteorology in general and tropical meteorology in particular have made tremendous strides during the last half a century or so after the introduction of computer technology and space satellites. Undeniably, we now know much more about our atmosphere and its behaviour than ever before. It is now well established that tropical circulation forms an integral part of the global general circulation and that there is a continual exchange of heat, momentum and moisture between the tropics and the rest of the atmosphere. Yet, tropical circulation has some distinctive characteristics of its own which need to be identified and studied independently in a more comprehensive manner. At present, uncertainty prevails in several areas of interest and there are many dark or twilight areas. In a seminal paper, Charney and Shukla (1981) had observed that quite unlike the midlatitude weather systems the predictability of which was limited by short period baroclinic wave activity, tropical weather systems which are determined by circulations between long-period heat sources and sinks are more predictable. Surface and boundary layer characteristics, such as ocean surface temperature, seasonal ground heating and cooling, have much longer lifespans and as such amenable to prediction over longer periods of time. It is proposed to show in the present text that tropical circulation systems are basically forced by boundary layer heat sources and sinks, on different time and space scales, though wave activity arising from flow instability plays a significant role in weather-forming processes. However, several questions regarding tropical circulations remain unresolved. For example, with a heat source centered at the equator and flanked by heat sinks in both zonal and meridional directions, what sort of circulations and weather patterns would evolve over the equatorial region? In this area, we haven’t gone far beyond what we learnt from the excellent theoretical work of Matsuno (1966), Gill (1980, 1982) and others. An uncertainty shrouds the existence or non-existence of double equatorial troughs in some oceans and continents. For that matter, what is the origin of the SW Pacific Convergence Zone (SPCZ), or the SW Atlantic Convergence Zone (SACZ)? Are they to be regarded as equatorial, tropical, subtropical or extratropical convergence zones? Monsoon is probably one area where our knowledge K. Saha, Tropical Circulation Systems and Monsoons, C Springer-Verlag Berlin Heidelberg 2010 DOI 10.1007/978-3-642-03373-5_1,
3
4
1 Large-Scale Tropical Circulations – Some General Aspects
and understanding lack precision. Firstly, its definition and then the manner of its advance and retreat over different parts of the globe! In the past, monsoon has been defined almost exclusively in terms of either rainfall or seasonal reversal of the prevailing surface wind. Little is known of the structure of monsoon and the manner of its advance and retreat in different areas. Notwithstanding commendable progress made in studies of monsoon (for a recent review of literature, see Webster et al., 1998) over the globe in general, we yet do not have a clear idea of its structure and how it advances or retreats over the Arabian Sea or the Bay of Bengal, for example, during the northern summer. What brings up the monsoon current over the Arabian Sea which remains relatively cold compared to its surrounding land areas during summer? Then, again, what brings about the well-known intraseasonal oscillation in monsoon rainfall and other meteorological parameters? In some parts of the tropics, the equatorial trough of low pressure which changes its orientation with season shows strong correlation with the genesis of tropical cyclones/hurricanes/typhoons. What is the connection between the two? Another area where we lack precise information is how large mountain systems, such as the Andes in South America, the Rockies in North America, the Mountains of East Africa and the Great Himalaya complex with Tibetan Plateau in Asia affect circulation systems over the respective continents. Our knowledge of the effect of ocean currents and ocean surface temperatures on equatorial circulation is also inadequate. Here we have not gone far beyond what we learnt from the pioneering studies of Walker (1924), Bjerknes (1966, 1969) and others. There are many such areas where we need further and better information from observations and their analysis and diagnosis, along with theoretical and dynamical studies, than heretofore. In a lecture delivered during a seminar in 1979, Charney had observed that ‘Data by themselves are not sufficient, nor is mind. We need a combination of both’. A summary of this lecture was prepared by J. Shukla (1979) from notes taken by him. To-day, equipped with much better network of surface and upper-air observations and improved methods of data analysis, we have much better opportunity to have a fresh look into some of the above-mentioned outstanding problems of tropical meteorology and apply our mind to them than ever before. Frankly, that is the main objective of this book.
1.2 Tropical Circulation as Part of the Global General Circulation – The Tradewinds From time immemorial, mariners sailing over oceans in low latitudes for trading purposes encountered a set of highly-steady seasonally-reversing winds which they called the tradewinds and which they used to great advantage in moving from place to place. In late 17th century, Edmund Halley (1686) was the first to make a detailed study of the tradewinds with the data then available and hypothesized that the observed
1.2
Tropical Circulation as Part of the Global General Circulation – The Tradewinds
5
winds at the surface were part of a direct thermally-driven circulation between a heat source and a heat sink, which reversed its direction between winter and summer But it was Hadley (1735) who offered an explanation for the cause of the tradewinds as well as their observed reversal of direction on the basis of differential heating between the equator and the poles and the rotation of the earth. He argued that a general equatorward drift of the tradewinds at low levels required a compensating poleward drift at high levels in order to prevent an undue accumulation of mass near the equator. Further, a general westward drag by the tradewinds near the earth’s surface at low latitudes due to the rotation of the earth required a compensating eastward drag by the westerlies in high latitudes so as to prevent a general slowing down of the earth. It was found later that the general westward or eastward drift of the wind was consistent with the principle of conservation of absolute angular momentum of the earth. A parcel of air moving equatorward from high latitudes in order to conserve the angular momentum of its original latitude would acquire an increasingly westward drift, while a poleward-moving parcel would acquire an increasingly eastward drift. This was due to the fact that the earth’s surface moved faster at the equator than at higher latitudes. Hadley’s idealized single-cell circulation model held ground and went unchallenged for nearly a century and it was once thought that Hadley’s model was representative of mean meridional circulation over all parts of the globe at all times of the year. However, later observations called for a modification of Hadley’s idealized single-cell model. The new observations revealed the presence of a well-marked high pressure belt over the subtropics and a low pressure belt further poleward near 60◦ latitude, which suggested a meridional pressure gradient and a poleward drift of air, instead of an equatorward drift near surface, and a compensating equatorward drift at some height, over the midlatitudes. Further, the westerly wind over the midlatitudes were found to be baroclinically unstable and characterized by large-scale eddy motion. Amongst the early attempts to modify Hadley’s original scheme were those of Thomson (1857) and Ferrel (1859) who introduced a shallow indirect cell, characterized by a poleward flow near the surface and equatorward flow at some height, over the midlatitudes, within the framework of the idealized single-cell Hadley circulation model. Further modifications to the meridional circulation model were made in the light of later observations. That tropical circulation forms a distinct integral part of a three-cell global general circulation model was first spelt out by Rossby (1947) and others almost two centuries after Hadley (1735) had proposed his single-cell poleto-equator zonally-averaged annual-mean general circulation model. The model proposed by Rossby (1947) and others, is in Fig. 1.1. Rossby’s three-cell meridional circulation model shows a direct circulation cell over the tropical belt with rising motion over the equatorial region and sinking motion over the subtropical belt, an indirect circulation cell over the midlatitudes and a direct circulation cell over the polar latitudes, with a polar front located at a latitude of about 60◦ . The Rossby model has, by and large, stood the test of time
6
1 Large-Scale Tropical Circulations – Some General Aspects
Fig. 1.1 Schematic of a three-cell meridional circulation model proposed by Rossby (1947) and others
and found general acceptance by the scientific community to this day. In this model, the classical Hadley circulation has been shown to have its rightful place over the tropical belt only, as against the original model with a pole-to-equator circulation. In the Rossby model, the tradewinds blowing towards the equator pick up increasing amounts of heat and moisture from the warm ocean surface and deliver them in rising currents over the equatorial belt to form cloud and precipitation and then turn around in the upper troposphere to blow as anti-trades towards the subtropical belt where they subside to low levels to blow as low-level trades or tradewinds again. This constitutes the direct kinetic energy-producing tradewind circulation cell of low latitudes shown in the Rossby model.
1.3 Poleward Boundary of the Tropical Circulation In the past, there appears to have been some controversy and divergence of opinion regarding the definition of tropics, particularly its poleward boundary. Climatologists divide climates on the earth’s surface into three distinct latitudinal belts, calling that which lies within the first 30◦ parallels of latitude, the tropics or the hot climate zone, as distinct from the other 30-degree belts which lie poleward with temperate and frigid climates. Defined in this way, the tropics is the hottest part of the earth’s surface with the sun always shining overhead somewhere over the region. Some people even put the poleward boundary of the tropics at the solstices, that is at 231/2◦ of latitude north and south of the equator. However, meteorologists, by and large, are aware that the poleward boundary of the tropical circulation is not determined by any fixed latitude, since it meanders about its annual mean location considerably with longitude and season. To them, the ridge of the global subtropical high pressure belt at surface, which divides the easterly tradewinds from the midlatitude westerly winds, would appear to be a more reasonable and acceptable choice. That this last is the correct view is also suggested by a three-dimensional examination of the structure of the tropical atmosphere vis-à-vis that of the higher
1.4
Heat Sources and Sinks
7
latitudes. Besides surface temperatures, the vertical distribution of temperature over the two belts is such as to create a fairly sharp thermal and circulation divide between the tropics and the midlatitudes, which creates a strong horizontal temperature gradient across a narrow boundary zone to drive a westerly wind maximum known as the subtropical jet (STJ) along the poleward boundary of the tropics. One of the earliest to find the existence of this jet was Palmen (1951) who studied the wintertime mean meridional circulation on the earth’s surface. Palmen’s finding has been substantiated by several other studies (e.g., Krishnamurti, 1961; Wallace and Hobbs, 1977; Galvin, 2007). According to observations, the boundary moves equatorward of its annual-mean latitude during the winter and poleward during the summer in a hemisphere. So, the tropics of our concern in the present text will be the latitudinal belt between the subtropical belts of the two hemispheres, bounded in each hemisphere by the ridge of the subtropical high pressure at surface and the subtropical jet in the upper air. The tropical circulation, however, is not isolated. In some parts of the globe, a tropical circulation system may extend to high extratropical latitudes, or an extratropical circulation system extend to very low tropical latitudes, and there is continual interaction and exchange of heat, mass, and momentum between the two circulations.
1.4 Heat Sources and Sinks Before proceeding further, we need to define what we mean by heat sources and sinks and state how we can identify them in the atmosphere.
1.4.1 Definition of Heat Sources/Sinks At any given time of the year, when the atmosphere over a place is continually heated (cooled) relative to its surrounding, the differential heating creates a heat source (sink) over the place. We define a heat source or a heat sink by the following criteria: dH/dt = 0,
∇ 2 dH/dt < 0,
Heat source
(1.4.1)
dH/dt = 0,
∇ 2 dH/dt > 0,
Heat sink
(1.4.2)
Where H denotes the steady-state heat content of the air at a given place at time t, and ∇ is a Del operator. Here, H = ρcp T, where ρ is density, cp is specific heat at constant pressure p, and T is temperature determined by local heat balance. Thus, to maintain a steady-state, a heat source must give up its excess heat to the environment, and a heat sink must make up its deficit heat by acquiring heat from the environment.
8
1 Large-Scale Tropical Circulations – Some General Aspects
1.4.2 Diabatic/Adiabatic Heat Sources/Sinks 1.4.2.1 Diabatic Heating/Cooling A diabatic heating or cooling process is one in which a working sample of the atmosphere is free to exchange heat with its environment. There are three important processes of diabatic heating or cooling in the atmosphere in the tropics. These are: (i) Absorption or emission of short- and long-wave radiation; (ii) Latent heat released by condensation of water vapour, or lost through evaporation of water, and mixing; and (iii) Sensible and latent heat gained from, or lost to, a boundary surface through turbulence and convection.
1.4.2.2 Adiabatic Heating/Cooling In an adiabatic process, the air sample has to work in a closed system and is barred from sharing its heat with the environment. This means that any change that occurs in its temperature is due to its own expansion or compression. In this process, air with upward velocity cools by expansion to a lower pressure and that with downward velocity warms up by compression to a higher pressure. 1.4.2.3 Temperature Change in the Atmosphere Due to Condensation Heating Both diabatic and adiabatic processes are important in producing temperature change at a point in the atmosphere. The most important diabatic heat source in a tropical disturbance is the release of latent heat of condensation of water vapour in the atmosphere which amounts to nearly 2.5 × 106 J kg–1 and this value increases with temperature. The temperature change produced by this process may be computed by using the approximate relation (Anthes, 1982) (∂T/∂t)cond = (L/cp )C,
(1.4.3)
where C is the local condensation rate (mass of water vapour condensed per unit mass of air per unit time) and L is the latent heat of condensation of water vapour. The local condensation rate in a saturated updraft, for example in a mature thunderstorm, may be computed by using the approximate relation C˜ ∼ −ω∂qs /∂p
(1.4.4)
where ∂qs /∂p represents the quantity of water released between two pressure surfaces assuming a saturated adiabatic lapse rate of temperature. Anthes (loc. cit.) finds that for a saturated adiabat through, say, 24◦ C, this term would yield about
1.5
Some Physical and Dynamical Constraints and Conservation Laws
9
5 g kg–1 of water in the 200 mb-layer between 800 and 600 mb. For an updraft velocity of 1000 mb/h, the value of C is 25 g kg–1 h–1 . Substitution of this value of C in (1.4.3) yields a condensation heating rate of about 1500 K (day)–1 . Of course, such an enormous heating rate is never observed in the atmosphere. This is because the diabatic heating is almost totally compensated by adiabatic expansion and cooling, so that actual change of temperature at any level is only a small imbalance between diabatic heating and adiabatic cooling. It is only near the lower boundary where vertical motions are constrained to be small that diabatic heating produces large changes of temperature. 1.4.2.4 Identification of Heat Sources and Sinks Heat sources and sinks are created in the atmosphere by both diabatic and adiabatic processes and may be identified generally by their effects on the distribution of temperature and pressure at surface and aloft. At the lower boundary of the atmosphere, a diabatic heat source (sink) may be identified with a ‘heat low’ (‘cold high’). In the atmosphere above, a diabatic heat source is created by condensation of water vapour and a diabatic heat sink by the reverse process of evaporation of drops, or by radiative cooling. In the atmosphere, in the absence of precipitation, convection generally leads to adiabatic cooling and subsidence to adiabatic warming. The processes lead to a ‘cold low’ and ‘warm high’ respectively. This means that a ‘cold low’ is produced when warm air rises and expands adiabatically by rising to a lower pressure, and a ‘warm high’ when a sample of cold air is compressed and heated adiabatically by subsidence to a higher pressure, i.e., to lower elevations.
1.5 Some Physical and Dynamical Constraints and Conservation Laws Circulations forced by heat sources and sinks are, however, required to comply with certain physical and dynamical constraints and conservation laws. Some of these are stated below.
1.5.1 Direct and Indirect Circulations When a heat source over a place gets continually heated up, it tends to maintain itself as a source by transferring the excess heat to a heat sink in the neighborhood. Likewise, when a heat sink continually gets colder, it tends to maintain itself as a heat sink by transferring its excess cold to a neighboring heat source. The heat exchanges are assumed to take place via a direct kinetic energy-producing vertical circulation in which warm air rises and cold air sinks. But due to adiabatic processes in the atmosphere, the rising air cools and the sinking air warms up. So, the transfer
10
1 Large-Scale Tropical Circulations – Some General Aspects
process requires an independent source of energy to provide for the kinetic energy of the direct circulation. This source must be an indirect circulation which would generate sufficient available potential energy to provide for the kinetic energy of the direct circulation.
1.5.2 Energy Transformations Since the total energy of the circulations in a closed frictionless domain must remain constant, it follows that the required available potential energy (both zonal and eddy, as defined by Eqs. (1.5.1) and (1.5.2) below), must be generated in one part of a system to provide for the kinetic energy of the other part, in accordance with the following relationships (after Krishnamurti and Surgi, 1987): Generation of zonal available potential energy (Pz ) γ [ < T > ][ < Q > ]dm (1.5.1) G(Pz ) = m
Generation of eddy available potential energy (Pe ) γ [ < T Q > + < T >∗ < Q >∗ ]dm G(Pe ) =
(1.5.2)
m
Conversion from zonal available potential energy (Pz ) to zonal kinetic energy (Kz ) < Pz .Kz >= − [ < ω > ] [ < α > ] dm (1.5.3) m
Conversion from eddy available potential energy (Pe ) to eddy kinetic energy (Ke )
[ < ω α > + < ω >∗ < α >∗ ]dm
< Pe .Ke >= − m
where γ is a stability parameter, and the symbols used with meanings are: ω vertical p-velocity, where p is pressure, T temperature, Q total diabatic heating per unit mass (m) of air, α specific volume, Deviation from zonal mean, Deviation from horizontal area average, m integration over mass, m, <> zonal average, [] time average, and ∗ deviation from time average.
(1.5.4)
1.5
Some Physical and Dynamical Constraints and Conservation Laws
11
1.5.3 Energy Transfer Process – Carnot’s Cycle In physics, Carnot’s heat engine works in a closed system and the process of heat transfer between a source and a sink is in a cycle which is strictly reversible. In the open atmosphere, the principle of reversibility clearly does not hold, since the working substance which is a part of the environment is exposed to the environment with which it can exchange heat. Yet, the general principle of the Carnot’s heat engine has been found to be useful in interpreting the atmospheric heat transfer processes. Let us see how it works. In the first stage, the working substance which is in contact with a heat source which in the atmospheric case is latent heat released by condensation draws a certain quantity of heat from the source by rising and expanding isothermally to a somewhat lower pressure; in the second stage, it expands adiabatically by rising to a much lower pressure and temperature; in the third stage, it descends isothermally to a somewhat higher pressure and delivers a certain quantity of heat to the heat sink with which the working substance is now placed in contact, from where an adiabatic compression to a higher pressure and temperature during the fourth stage will restore the working substance to its original pressure and temperature. Thus, the working of the Carnot’s cycle in the atmosphere represents a heat transfer process in which cold air is raised adiabatically to become colder and warm air lowered adiabatically to become warmer, a process which generates available potential energy by an indirect circulation.
1.5.4 Conditional Instability and Convection The tropical atmosphere most often remains in a state of conditional instability in the sense that it is unstable (∂θ e /∂z < 0) in the lower troposphere below about 700 hPa, but stable (∂θ e /∂z > 0) above, where θ e is the equivalent potential temperature of the air which may be defined as ‘the temperature attained by a parcel of air which is raised adiabatically from its existing level to a level where all its moisture is condensed out and the latent heat of condensation added to it, and then brought down dry-adiabatically to a standard level, usually 1000 mb.’ However, it does not follow from this stability condition that it will automatically lead to convective overturning. In the tropics, moisture is usually confined to a shallow boundary layer in contact with the earth’s surface. So, convective overturning is only possible when this low level moisture is lifted to higher levels by synoptic-scale convergence. It is not surprising, therefore, that most of the tropical disturbances such as depressions and cyclones develop over oceans where low level convergence is able to lift the boundary layer moisture to higher levels for large-scale release of the latent heat of condensation which will potentially lead to explosive growth of the disturbance. Such developments may also occur over land areas where there
12
1 Large-Scale Tropical Circulations – Some General Aspects
is a copious supply of moisture from neighboring oceans and strong lapse rate of temperature may develop during afternoon-evening hours. In all cases, cyclogenesis must lift the low-level moisture to higher levels for the latent instability energy of the atmosphere to be released.
1.5.5 Cellular Structure – Shallow and Deep Convection It is well-known from laboratory experiments as well as theoretical and observational studies that when the atmosphere is thermodynamically and/or hydrodynamically unstable, it breaks down into convective cells of different dimensions depending upon the depth of the layer involved and the amount of moisture present. These convective cells may appear in the sky in the form of clouds of different horizontal and vertical extents. For example, when instability is present in a shallow layer above the earth’s surface and there is a limited supply of moisture, we may see only small-scale fair weather cumulus type of clouds. On the other hand, when instability involves a deeper layer of the atmosphere and there is deep moisture convergence at low levels with divergence above, clouds may appear in different layers and some of the cloud cells may grow to great heights. The cellular structure of the atmosphere then becomes quite clearly visible. However, when more than one layer is involved, the cells must arrange themselves vertically so as to secure sufficient vertical compensation to ensure a small pressure change that is actually observed at the earth’s surface. In a developing system, however, there is continual rearrangement of the cells which may result in a large pressure change. In the tropics, the cellular structure of atmospheric circulation is ubiquitous. However, the actual structure of the cells in any case depends upon the configuration and dimensions of heat sources and sinks and their orientation. It must also satisfy the requirements of energy conservation and transformations and physical and dynamical constraints discussed earlier in this section.
1.5.6 Coriolis Control-Variation with Latitude The relationship between pressure and wind is largely controlled by the Coriolis force, apart from friction and gravity. In higher latitudes where Coriolis control is strong, pressure rapidly adjusts to the wind field, but in the tropics where the Coriolis control is weak, the wind tends to adjust to the pressure field. So, in the absence of friction or any other force, when there is a pressure gradient across the equator where the Coriolis control vanishes, there can be a direct cross-equatorial airflow down the pressure gradient. However, away from the equator, the wind becomes increasingly quasi-geostrophic.
1.5
Some Physical and Dynamical Constraints and Conservation Laws
13
1.5.7 Conservation Laws In steady state, tropical circulation is required to satisfy certain conservation laws. These include: (a) The Law of conservation of heat energy, and (b) The Law of conservation of potential vorticity. (a) Conservation of Heat Energy: As mentioned earlier, condensation heating is one of the powerful diabatic heat sources in the atmosphere but its effect is almost totally offset by adiabatic processes. The balance between the two processes is governed by the thermodynamic energy equation in the form ∂T/∂t = −(u∂T/∂x + v∂T/∂y) + σ ω + (1/cp )δQ/dt
(1.5.5)
Where T is temperature, u, v are the components of the wind along the coordinate axes x, y respectively along the pressure surface p, ω is the vertical p-velocity, cp is the specific heat at constant pressure, δQ/dt is the rate of diabatic heating, and σ is the static stability parameter given by the relation, σ (= κT/p–∂T/∂p), where κ = R/cp , and R is Gas constant. The balance is reached when the left-hand side vanishes. In the tropics, the third term on the right hand side of (1.5.5) balances largely with the second term, while in the midlatitudes, the first term becomes important as well. (b) Conservation of Potential Vorticity: The principle of conservation of potential vorticity stated in the form (1.5.6) (ζ +f )/(δθ/δp) = Constant
(1.5.6)
(Where ζ is relative vorticity, f is Coriolis parameter, θ is potential temperature, and p is pressure) is a powerful constraint on tropical circulation though the equation was derived on the assumptions that the atmosphere was frictionless, adiabatic, and barotropic. Obviously, the assumptions are not quite realistic, since frictional forces are always at work in the earth’s boundary layer, and adiabatic processes may be more or less compensated by diabatic processes, and baroclinicity cannot be ruled out from the tropical atmosphere. Yet, the principle of conservation of potential vorticity is useful in interpreting the distribution of vorticity in airflows negotiating large mountains. For example, during northern winter, the cold NE-ly winds with negative relative vorticity over the Tibetan Plateau, on descending to the Plains of Northern India, develops a narrow zone of positive relative vorticity before resuming the normal negative relative vorticity over Central India. This change appears to be in conformity with the principle of conservation of potential vorticity. The principle appears to explain similar changes in relative vorticity in several other mountainous regions of the globe.
14
1 Large-Scale Tropical Circulations – Some General Aspects
1.6 Equatorial Circulations In the annual mean, the equatorial region of the earth receives the maximum solar radiation and may act as a perennial heat source. However, owing to distribution of continents and oceans and differential heating between them, heat sources and sinks are created along the equator. Heat sources and sinks are also created over equatorial oceans by powerful warm and cold ocean currents, especially between their western and eastern parts.
1.6.1 Circulation with Heat Sources and Sinks Placed Alternately Along the Equator – Walker Circulations A general problem of this type was first addressed by Matsuno (1966) theoretically by using a one layer homogeneous divergent barotropic model and later by Gill (1982) and others. In the first part of his treatment, Matsuno studied the different types of wave motions like Rossby and inertio-gravity oscillations that are excited in the equatorial atmosphere when he placed mass sources and sinks alternately along the equator, but in the second part he addresses the problem of forced stationary circulation and obtained the following interesting results: (i) In latitudes somewhat away from the equator, the surface tends to be raised where mass is added, and lowered where mass is extracted; (ii) In the immediate vicinity of the equator, however, the deviations of the surface elevation is less than that in the higher latitudes in magnitude. Consequently, high and low pressure cells tend to be divided into two parts, one on each side of the equator; (iii) Strong zonal flow is created along the equator when mass flows from source to sink. The zonal flow along the equator is intensified by convergence of flow from higher latitude circulations and weakened by divergence of equatorial flow to higher latitudes; (iv) In the higher latitude region, the velocity fields are in geostrophic balance. When he applied the same boundary conditions to the two-level model of the atmosphere in which heat sources and sinks are placed alternately along the equator, the differential heating produces low pressure and high pressure respectively. The wind blows geostrophically in the high latitude region. The convergence or divergence of such winds along the equator induces vertical motions. These vertical motions counteract to the imposed heat sources and sinks, and their effects are strongest along the equator. Consequently, the warm air associated with the heat source is split into two parts, by the cold air belt located at the equator. Similarly, the cold air associated with the heat sink is split into two parts by the warm air belt located at the equator. Matsuno’s results in the case of the forced stationary circulation are depicted in Fig. 1.2(a, b, c). In Fig. 1.2, mass sources and sinks are shown in terms of high (H) and low (L) pressure systems, along with the induced vertical circulations in the lower panel
1.7
Meridional Circulation with Heat Source at the Equator and Heat Sinks
15
Fig. 1.2 Forced stationary horizontal (b) and vertical (c) circulations caused by imposition of mass sources (+) and mass sinks (–) alternately along the equator (a). In (b), pressure replaces mass; H for source, L for sink, CV denotes convergence, DV divergence (adapted from Matsuno, 1966)
(c) which has come to be known as the Walker Circulation. The arrows show the directions of air motion. It is easy to see from Fig. 1.2(c) that the regions of low-level convergence and upper-level divergence are those of dense clouding and heavy precipitation (as shown by hatching below the bottom line in the lower panel) and where they diverge at low levels and converge aloft are relatively cloud-free areas with little or no precipitation. According to observations, heavy clouding and precipitation occur along the equator in some preferred regions, such as equatorial Eastern Indian Ocean, Western Pacific Ocean, the Amazon basin of South America, and some parts of equatorial Africa.
1.7 Meridional Circulation with Heat Source at the Equator and Heat Sinks in Higher Latitudes – The Hadley Circulations The classical Hadley circulation cell visualizes a zonally-symmetric annual-mean meridional circulation between a heat source with its cyclonic circulation centered at the equator and a heat sink characterized by anticyclonic circulation centered at
16
1 Large-Scale Tropical Circulations – Some General Aspects
the ridge of the subtropical high pressure with rising motion at the equator and subsidence over the subtropical belt. Tradewinds diverging from the high pressure belt were assumed to travel equatorward, converge at the equator and rise in penetrative convection producing cloud and rain and thereby releasing latent heat of condensation of water vapour carried by the tradewinds. At the equator, the rising currents are assumed to diverge in the upper troposphere and flow poleward as anti-trades to give up heat and subside over the tropical belt in order to flow back again towards the equator as tradewinds in some kind of a meridional-vertical circulation. It was this mean meridional circulation which was assumed to transfer sensible and latent heat from the equator poleward. However, when applied to the real atmosphere, the idealized Hadley circulation cell model as described in the preceding para, faces several difficulties. Observation shows that the equatorial tropopause with a temperature of about –80◦ C at a height of about 16 km above sea level is the coldest place in the tropical atmosphere and as such a transfer of heat from the equator poleward in the upper troposphere against a temperature gradient is not possible, as envisaged in the classical model. The inadequacy of the classical single-cell model in this regard was first pointed out by Fletcher (1945) who during a flight across the equatorial eastern Indian Ocean during the Second World War found two well-organized zonally-oriented cloud bands, one on each side of the equator, with little or no cloud in between over the equator. To overcome the difficulty of explaining the observation with the classical Hadley circulation model, Fletcher proposed a revised model with a small-scale meridionalvertical circulation cell interposed between the equator and the Hadley cell. The equatorial cell was supposed to have subsidence at the equator and rising motion where the Hadley cell circulation would converge into the equatorial circulation at some distance away from the equator. We call this convergence zone, the Tropical Convergence Zone (TCZ). The Fletcher model was supported by Rossby (1947) and several others. But it soon became apparent that the Fletcher model which visualized a single equatorial vertical circulation cell may also have problems in explaining some observed phenomena in the atmosphere. One of these concerns the formation of different types of clouds and release of condensation heating in different layers of the atmosphere. In small-scale cumulus-type clouds, maximum heating is likely to be small and confined mostly to a lower level. In large cumulonimbus-type of clouds, maximum heat is released by condensation in the upper troposphere. In the vertical distribution of diabatic heating in the mean tropical atmosphere, the level of maximum heating appears at about 400 mb (Holton, 1979). This fact appears to suggest that more than one layer of convective clouds may be involved in releasing latent heat of condensation in the atmosphere and that the upper layer may, in fact, contribute a greater amount of diabatic heat to the warming of the equatorial troposphere than the lower layer. In the present text, we, therefore, suggest a further revision to the Classical Hadley cell model, by dividing the single equatorial cell into two cells, a Lower Equatorial (LE) and an Upper Equatorial (UE), as shown in Fig. 1.3.
1.8
Seasonal Migration of the Equatorial Heat Source
17
Fig. 1.3 Schematic showing the suggested revised model of the Hadley circulation with two equatorial cells (UE, Upper-Equatorial) and (LE, Lower Equatorial) and the location of the TCZ; Symbols DV means Divergence, CV Convergence. Arrow shows the direction of air motion
1.8 Seasonal Migration of the Equatorial Heat Source 1.8.1 Origin of Monsoon A northward movement of the equatorial heat source from the equator is accompanied by change in the structure of the associated Hadley circulations described in the preceding section. The SE’ly tradewinds, which were earlier confined to the southern hemisphere, cross the equator directly and turn into SW/W’ly tradewinds. So the movement of the equatorial heat source brings about a latitudinal expansion of the belt of the W’ly tradewinds to the south of its trough with influx of cool, moist airmass from across the equator. On converging into the circulation around the heat source on its equatorward side at low levels, the convergence produces the well-known Intertropical Convergence Zone (ITCZ) where the converging winds rise in penetrative convection, precipitate, and diverge in the upper troposphere. A branch of the diverging currents returns equatorward as NE-ly antitrades to sink over the region of the heat sink where on sinking it joins the low-level diverging current which converges into the ITCZ in order to complete a vertical circulation which has come to be known as the Monsoon circulation, while the other branch of the diverging upper currents moving poleward sinks over higher latitudes to form the upper branch of the Hadley circulation. However, within the framework of the Hadley circulation of the northern hemisphere, there is a secondary vertical circulation associated with the TCZ where the converging low-level E/NE tradewinds rise in subdued convection, often producing clouding and light rain. The structure and properties of the monsoon circulation appear in clear perspective when we consider a general case of horizontal circulation around the equatorial heat source, interacting with the tradewind circulations at low levels, as shown schematically in Fig. 1.4, for summer in: (a) Northern Hemisphere, and (b) Southern Hemisphere.
18
1 Large-Scale Tropical Circulations – Some General Aspects
Fig. 1.4 Schematic showing the structure of the horizontal circulation around the equatorial heat source (EQ.TR.) and the locations of the ITCZ and the TCZ where the tradewinds (thick bold lines with arrows) of the two hemispheres converge: (a) Northern Hemisphere (N.H.), (b) Southern hemisphere (S.H.)
1.8.2 The Wave Structure The circulations depicted in Fig. 1.4 reveal the following general features: 1. The equatorial heat source as shown has two troughs of low pressure, one zonally-oriented (marked EQ.TR.) and the other meridionally-oriented (unmarked); they separate out four heat sinks, two on each side; 2. The winds diverging from each heat sink appear to converge into the circulation around the heat source, forming two distinct convergence zones with a zone of divergence in between on each side of the EQ.TR: these are the tradewinds; (a) the cold NE trades on the poleward side representing the Hadley circulation, and the cold SW trades on the equatorward side representing the Monsoon circulation in the Northern Hemisphere (N.H.); (b) the corresponding flows in the Southern Hemisphere (S.H.) are the cold NW Monsoon current on the equatorward side, and the cold SE Monsoon current on the poleward side; 3. The tradewind circulation on either side of the EQ.TR appears to have the structure of a wave characterized by two convergence zones separated by a zone of divergence; a wave associated with the TCZ on the poleward side, and a wave associated with the ITCZ on the equatorward side of the EQ.TR;
1.8
Seasonal Migration of the Equatorial Heat Source
19
4. Over most parts of the globe, the average zonal wavelength of the monsoon wave appears to be about 2000–2500 km; 5. In the waves, it is the convergence zones where penetrative convection and precipitation occur, with relatively clear, or less cloudy, conditions in the divergence zone in between; 6. The Monsoon and the Hadley circulations appear to co-exist with the equatorial heat source circulation at all times and in all monsoon climes. They may interact with traveling waves in forming tropical disturbances.
1.8.3 Forcing for the Seasonal Movement of the Equatorial Heat Source It is well-known that in the summer hemisphere, given the same solar radiation, temperature rises and pressure falls over both land and ocean, but the changes occur much more rapidly and with greater amplitude over land than over neighboring ocean due to much lower heat capacity of the land compared to that of the ocean. Conversely, for the same reason, in the winter hemisphere, temperature drops and pressure rises much faster and with greater amplitude over the land than over the neighboring ocean. An example of this difference in pressure tendency between land and ocean is presented in Fig. 1.5, which shows the latitudinal distribution of mean sea level pressure along longitude 70◦ E over the Arabian Sea and adjoining Western India during January (DJF) and July (JJA).
Fig. 1.5 Latitudinal distribution of mean sea level pressure along 70◦ E over the Arabian Sea sector during January (DJF) and July (JJA)
20
1 Large-Scale Tropical Circulations – Some General Aspects
So, it is the gradient of pressure tendency (also called isallobaric gradient) between the land and the ocean that appears to force the equatorial trough of low pressure over the ocean to move towards the heat source over the neighboring land in the summer hemisphere with a velocity given approximately by the well-known kinematic relation (Petterssen, 1956) C = −∇(∂p/∂t)/∇ 2 p
(1.8.1)
where C is the velocity vector, p is pressure, t is time, ∇ is Vector Del operator, (∂p/∂t) is pressure tendency, the denominator is the curvature of the pressure field in the trough, and the isallobaric gradient is in a direction at right angle to the axis of the trough of the equatorial heat source. In a study of the onset, advance and withdrawal of summer monsoon over the Indian Subcontinent, Saha and Saha (1980) tested the validity of this hypothesis qualitatively by applying it to the case of advance and withdrawal of summer monsoon along a narrow longitudinal belt parallel to the West Coast of India (about 73◦ E). Since pressure and height are related hydrostatically, the tendencies of mean monthly height of the 850 mb surface at eight stations along the meridian were calculated using the mean height data available from Ramage and Raman (1972). Results are shown in Fig. 1.6.
Fig. 1.6 Values of 850 mb mean monthly height tendency (gpm/month) in different months at Minicoy (MNC), Bangalore (BNG), Bombay (BOM), Ahmedabad (AHM), Jodhpur (JDP), Delhi (DLH), Lahore (LHR), and Peshawar (PWR). Height data are taken from Ramage and Raman (1972). Arrows show the direction of the height tendency gradient during periods of advance and withdrawal of monsoon
1.8
Seasonal Migration of the Equatorial Heat Source
21
In regions where monsoon is interhemispheric, the forcing for cross-equatorial flow appears to be provided by the isallobaric gradient between the heat source on one side and the heat sink on the other. The crossing appears to open a floodgate for cold, humid airmass of the winter hemisphere to rush into the summer hemisphere to start the process of advance of summer monsoon in that hemisphere. A reversal in the direction of the isallobaric gradient is what causes the retreat of the monsoon. The isallobaric gradient is clearly northward (indicated by a long arrow in Fig. 1.6) during the period April–June, and southward during the period September–November.
1.8.4 Intraseasonal Oscillation of Monsoon The advance and retreat of monsoon as a result of the movement of the equatorial heat source and its associated heat sinks would produce an intraseasonal oscillation in usual meteorological fields of the region. Figure 1.7 illustrates the rationale behind occurrence of intraseasonal oscillation on this account in the field of atmospheric pressure. This intraseasonal oscillation appears to result from the superposition of the monsoon perturbation upon the seasonal oscillation of pressure over the region. An example of intraseasonal oscillation in mean monthly maximum and minimum temperatures at Delhi (India) is shown in Fig. 1.8.
Fig. 1.7 Illustrating the formation of intraseasonal oscillation in atmospheric pressure (P´ denotes perturbation pressure) during advance and retreat of monsoon: Symbols used: H – High, L – Low, C – Cold, W – Warm. Other symbols have their usual meanings
22
1 Large-Scale Tropical Circulations – Some General Aspects
Fig. 1.8 Distribution of mean monthly maximum (continuous line) and minimum (dashed line) temperatures at Delhi
Intraseasonal oscillation is observed in other meteorological parameters as well; for example, in temperature, humidity, rainfall, etc. It is well-known that people living in monsoon regions experience prolonged rainy spells twice during the year, once during advance and a second time during retreat of monsoon. Also, as for temperature, it is not always cool and humid during the monsoon season. During certain periods within the season, temperature rises and weather becomes unbearably hot and humid, at least twice during the season, once during advance and a second time during retreat. For example, at New Delhi, with the arrival of monsoon towards the end of June, the daytime air temperature may rapidly drop from about 46 to 40◦ C or even lower, but it does not remain at that reduced value for long, since it rises again at least twice, the first in August when the monsoon wave moves up to the Western Himalayas, and a second time in October during the retreat of the trough. An intraseasonal oscillation of this kind occurs in other airmass properties as well, such as sunshine hours, air quality, physical comfort level, etc. Similar intraseasonal oscillations are believed to be occurring in monsoons in other regions as well.
1.9 Co-existence of Monsoon, Hadley and Walker Circulations – Inclined Troughs Over most parts of the tropics, the equatorial troughs of low pressure are seldom zonally or meridionally oriented. They are inclined to latitudes or longitudes. There may be several reasons for the observed inclination, but the main reason appears to be the general inclination of the coastline. In summer, when a powerful heat low develops over the land, its sphere of influence extends over the neighboring ocean and tends to draw an oceanic trough of low pressure which may come within its sphere of influence. So, the oceanic trough gets inclined to the heated land. In winter, when a cold high pressure develops over the land, the oceanic trough will move away from the cold land. Thus, in general, the trough will be inclined towards the nearest land during summer and away from it during winter.
1.9
Co-existence of Monsoon, Hadley and Walker Circulations – Inclined Troughs
23
Inclined troughs are found over all the global oceans, on both sides of the equator, during both summer and winter. The areas which are particularly noted for occurrence of inclined troughs are the following: 1. 2. 3. 4.
The western and the Eastern parts of the Pacific and the Atlantic Oceans; Eastern Arabian Sea and the Bay of Bengal; A wide area of the southwestern Indian Ocean close to the coast of Madagascar; A wide area of the southeastern Indian Ocean close to the Indonesian Islands and the coast of northwestern Australia; 5. A long stretch of the southwestern Pacific Ocean, extending southeastward from New Guinea across the Coral Sea.
Fig. 1.9 Streamlines showing the circulations around an NW–SE oriented inclined trough during northern summer: (a) a plan view; (b) vertical circulation in a zonal section through the center of the trough
24
1 Large-Scale Tropical Circulations – Some General Aspects
There is re-organization of the associated circulation cells when a traveling wave interacts with a quasi-stationary inclined equatorial trough. The TCZ is then activated by additional convergence at low levels and divergence at high levels which favor development. Figure 1.9 shows schematically (a) a plan view of a NW–SE oriented inclined trough, and (b) the likely zonal-vertical circulation across the trough in the N.H. An inclined trough is more than of academic interest, for it appears to mark out the regions noted for development of tropical cyclones. But, what is the connection between the trough and the cyclone and how is development favored by an inclined trough? This aspect of the question will be looked into further in the next chapter after we have introduced the easterly and the westerly waves which initiate the development.
1.10 Definition of Monsoon In the past, the word ‘monsoon’ has been defined exclusively in terms of either seasonal rainfall (Rao, 1976), or reversal of the direction of the prevailing surface wind between summer and winter (Ramage, 1971). For people wholly dependent upon rainfall for water resources, food production, etc., especially in developing countries of Southeast Asia, Africa and South America, a year with subnormal rainfall means to them poor monsoon and that with above-normal rainfall good monsoon. To these people, the word ‘monsoon’ is almost synonymous with rainfall. The Arabs first noted the seasonal reversal of the surface wind direction while sailing over the Arabian Sea more than a thousand years ago and coined the word ‘Mawsim’ for the phenomenon. This definition has also continued to this day. Ramage (1971) defines monsoon by the following criteria: (1) The prevailing wind direction shifts by at least 120◦ between January and July; (2) The average frequency of prevailing wind directions in January and July exceeds 40%; (3) The mean resultant winds in at least one of the months exceed 3 m s–1 ; and (4) Less than one cyclone-anticyclone alternation occurs every two years in either month in a 5◦ latitude-longitude rectangle. Definitions on similar lines, seeking to identify monsoon by either rainfall, or the seasonal reversal of the prevailing surface wind, have been advanced by several workers. None of the other characteristic features of the monsoon are ever mentioned in these definitions. It was Edmund Halley (1686) of England who appears to have been the first to conceive monsoon as a seasonally-reversing tradewind circulation when he wrote in his celebrated paper in the Philosophical Transactions of the Royal Society of London, the following:
1.11
Global and Regional Distribution of Monsoons
25
But as the cool and dense air, by reason of its greater gravity, presses on the hot and rarefied, it is demonstrable, that this latter must ascend in a continued stream, as fast as it rarefies, and that being ascended, it must disperse itself to preserve the equilibrium; that is, by a contrary current, the upper air must move from those parts where the heat is greatest: so by a kind of circulation, the north-east tradewind below will be attended with a south-westerly above, and the south-easterly with a north-west wind above. And that this is more than a mere conjecture, the almost instantaneous change of the wind to the opposite point, which is frequently found in passing the limits of the trade winds, seems to assure us; but that which above all confirms this hypothesis, is the phenomenon of monsoons, by this means most easily solved, and without it hardly explicable.
Our analysis of monsoon as a divergent tradewind circulation which converges into the seasonally-migrating equatorial heat source, producing a wave on the equatorward side of its trough prompts us to offer the following holistic, comprehensive, and yet simple, alternative definition of tropical monsoon: “Monsoon is a large-scale perturbation of the tradewind circulation associated with the seasonal movement of the equatorial heat source, which converges into the circulation around the heat source on its equatorward side at low levels, producing a wave with two zones of precipitation and a zone of clearance in between, and a host of other characteristic changes in airmass properties, during its advance and retreat.”
The foregoing definition brings within its fold not only the observed rainfall and reversal of wind direction, but also a host of other characteristic changes in prevailing meteorological conditions, such as changes in air temperature, cloudiness, etc. It also produces an intraseasonal oscillation in meteorological parameters during advance and retreat because of its wave structure. It is believed that the definition given here can be used as a dependable criterion for identifying and tracking monsoon over any part of the globe at any time of the year. Our definition implies co-existence of monsoon with Hadley and, at some locations, Walker-type east-west circulations. However, in offering the above definition of monsoon, we should not lose sight of some other regions of the globe, especially the extratropics, where such thermal contrast may arise between a large continent and a neighboring ocean and force a seasonal movement of the oceanic trough of low pressure between land and ocean. Of course, in such regions, the prevailing winds and eastward-propagating baroclinic wave disturbances may often interfere with the weak monsoon circulation that may develop and make its identification difficult.
1.11 Global and Regional Distribution of Monsoons 1.11.1 Tropical Monsoons Since tropical monsoon constitutes a perturbation of the tradewind circulation that converges into the ITCZ, the seasonal movement of the ITCZ offers a practical means to identify the leading edge of the monsoon over a region at any time of the year (e.g., Riehl, 1954; Saha, 1978; Walisser and Gautier, 1993; Saha et al., 1998).
26
1 Large-Scale Tropical Circulations – Some General Aspects
In other words, the region swept out by the ITCZ in the course of its movement between summer and winter constitutes the domain of the tropical monsoon in that region. A definition of a monsoon region on similar lines based on the seasonal movement of the ITCZ was also advanced by Asnani (1993). In general, tropical monsoons are interhemispheric in character and usually follow the seasonal movement of heat lows over the land from one hemisphere to the other. Three continental sectors stand out in this respect, viz., Asia, Africa, and the Americas. In the Asia sector, the movement is between Australia and Asia; in the African sector, between Northern and Southern Africa; and in the American sector between South America, and Central and North America. The monsoon regions in these sectors are as follows: Asia–Australia sector Region I Indian Sub-continent and adjoining SE Asia Region II Eastern Asia including China, Korea and Japan Region III Maritime Continent including Indonesia and Philippines Region IV Australia Africa sector Region V North and South Africa American sector Region VI South America Region VII Central America and adjoining States of Southwestern North America In general, the domains of oceanic monsoons are much narrower, and confined to latitudes closer to the equator. However, they are much wider over the western and the eastern parts of the Atlantic and the Pacific Oceans where they tend to join up with the continental monsoons along inclined troughs.
1.11.2 Extratropical Monsoons Monsoonal-type atmospheric circulations between continents and oceans are also observed in some extratropical regions of the globe where seasonal contrasts in temperatures between land and ocean and differential heating between them bring about seasonal movements of the circulation systems and associated rainbelts across the coastline and to higher latitudes. They are also observed over some extratropical mountain plateau regions where the circulations and the associated rainbelts are affected by seasonal reversals of temperature and pressure between the Plateau and the neighboring Lowlands or oceans. However, it is often difficult to see the full impact of such monsoons in the face of interference from strong planetary wind systems and frequent movement of baroclinic wave disturbances across the region. The concept of a monsoon in extratropical latitudes is not new, though it has been opposed by some meteorologists (e.g., Ramage, 1971). It was advocated by several
1.11
Global and Regional Distribution of Monsoons
27
workers (e.g., Alisov, 1954; Khromov, 1957 and others), though the steadiness of the prevailing surface wind, which was cited to be one of the main characteristics of a monsoon circulation, was found to be somewhat low on account of the interference from disturbances (Klein, 1957). It is proposed to show in Part 3 of this text that the idea of an extratropical monsoon is not just a fantasy, but a stark reality. It is driven by differential heating between a heat source and a heat sink in much the same way as in a tropical monsoon, but in a different environment. A tropical monsoon results from the seasonal movement of an oceanic heat source called the equatorial heat source from one hemisphere to the other and its movement is limited to the tropical belt. For this monsoon, the subtropical high pressure cells with their anticyclonic circulations and diverging tradewinds serve as heat sinks. On the other hand, for an extratropical monsoon, the same subtropical high pressure cells would act as heat sinks but heat sources would lie over the extratropical belt. The exact mechanism of this process will be elaborated in Chaps. 5 and 11 with a few case studies. Figure 1.10 shows the approximate domains of tropical and extratropical monsoons over different parts of the globe including the oceans.
Fig. 1.10 Global distribution of principal monsoon regions (stippled areas) bounded by the January and July locations of the ITCZ: January (full line); July (dashed line). Regions marked I to VII are areas of tropical continental monsoons; VIII and IX are domains of Extratropical monsoon; SPCZ denotes Southwest Pacific Convergence Zone, and SACZ the Southwest Atlantic Convergence Zone
1.11.3 Zonal and Meridional Anomalies As shown in Fig. 1.10, large zonal and meridional anomalies exist in the distribution of monsoons over the globe. These anomalies arise from the following:
28
1 Large-Scale Tropical Circulations – Some General Aspects
(a) Geographical distribution of continents and oceans; (b) Influence of ocean currents – oceanic monsoons; and (c) Effects of large-scale orographic barriers. (a) Geographical Distribution of Continents and Oceans: Whatever may be the geological origin of continents, their locations in a certain alignment on the earth’s surface, appears to be very important for the seasonal movement of monsoons. This happens in three major continental sectors, viz., Asia-Australia, North and South Africa, and North and South America. A quick glance at the world map shows that in each sector, the landmasses are oriented more or less in a general NW–SE direction, and monsoons in general follow this direction in their interhemispheric movement following the seasonal movement of the sun. (b) Influence of Ocean Currents – Oceanic Monsoons: Over all the world oceans, the subtropical anticyclonic gyres and the coastal ocean currents exercise large control on the distribution of ocean surface temperature. Their seasonal shift gives rise to oceanic monsoons and in this process ocean currents play a very major role. For example, in the Pacific Ocean, the seasonal location of the ITCZ is affected by the following major ocean currents: (1) The poleward-flowing warm Kuroshio Current in the northwest; (2) the equatorward-flowing cold California Current in the northeast; and (3) the equatorward-flowing cold Peruvian or Humbolt Current in the southeast. On account of the cross-equatorial flow of the Peruvian Current, the equatorial eastern Pacific Ocean remains cold almost throughout the year. Similarly, in the Atlantic Ocean, the corresponding ocean currents are: (1) the poleward-flowing warm Gulfstream in the northwest; (2) the equatorward-flowing cold Canaries current in the northeast; and (3) the equatorward-flowing cold Benguela current in the southeast. The equatorward-flowing cold Benguela Current keeps the temperature of the equatorial eastern Atlantic Ocean low almost throughout the year. On account of these cross-equatorial flows of cold ocean currents, the ITCZ seldom appears south of the equator over these oceans. The situation over the Indian Ocean is somewhat different. Here, in the western Indian Ocean, because of the dominant influence of the cold Somali current during northern summer, a zonal anomaly of surface temperature is maintained between the western and the eastern parts of the equatorial ocean almost throughout the year. In the Southern Indian Ocean, the equator-flowing cold West Australian current keeps the surface temperature of the eastern part relatively cold, while in the western part the poleward-flowing warm Agulhas current maintains a warm ocean surface. These movements of the ocean currents are reflected in the equatorial distribution of ocean surface temperature shown in Fig. 1.11. In some of the oceans where equatorial belt remains cold, the ITCZ is displaced from the equator to the relatively warmer parts of the ocean or to the nearest warm continent. For example, the SW Pacific Convergence Zone which runs from equatorial New Guinea eastsoutheastward across the Coral Sea area may be looked upon as the displaced ITCZ of the southern Pacific Ocean. Similarly, in the western part of the equatorial Indian Ocean which remains relatively cold throughout the year, the ITCZ is displaced to about 15◦ S during northern winter and to about 25◦ N during
1.12
Co-existence of Monsoon with Desert Circulation
29
Fig. 1.11 Equatorial distribution of mean ocean surface temperature during February (continuous line), and August (dashed line) (after Defant, 1961)
northern summer. Similar displacements may also be observed over other parts of the tropical oceans and continents. (c) Effects of Mountain Barriers: Large mountain systems, such as the Himalayas of Asia, the Mountains of East Africa, the Rockies of North America, the Andes of South America, and the Great Dividing Range of Australia, all appear to affect the circulations in their respective areas by relocating the ITCZ and other troughs and ridges of pressure through their mechanical and thermal effects on circulation. In accordance with the principle of conservation of potential vorticity, the blocking of a W’ly flow without horizontal shear approaching a north-south oriented mountain range will produce a trough of low pressure in the run-up to the base of the mountain, then a high pressure as it climbs up to the top, and then a low pressure as it descends on the leeside, till it resumes its zonal flow. In the case of an E’ly flow approaching such a mountain range, there is a difference. The flow senses the presence of the mountain barrier from a long distance and accordingly adjusts the pressure in such a way that will enable it to get over the top and emerge as an E’ly current on the leeside. This it does by first producing a low pressure trough and then a high pressure ridge of such intensity as will enable it to cross the mountain and emerge on the other side as an E’ly current. The ITCZ or the TCZ, as the case may be, is relocated whenever the airflow around the equatorial trough has to negotiate these high mountain barriers.
1.12 Co-existence of Monsoon with Desert Circulation It is by no means an accident that the world’s principal monsoons co-exist with subtropical deserts on their poleward sides. In the northern hemisphere, the great central Asian desert, the great Saharan desert in Africa, the Mojave-Sonoan desert
30
1 Large-Scale Tropical Circulations – Some General Aspects
in Central America and adjoining southwestern North America, all lie on the poleward sides of monsoons in the respective regions. Similar co-existence of monsoons with deserts on their poleward sides is also observed in the southern hemisphere in all the continents; for example, the Great Gibson and Western Deserts in Australia, the Namib-Kalahari desert in Southern Africa and the Patagonia desert in South America all lie on the poleward sides of their respective monsoons. So, a coexistence of monsoons and desert circulations appears to be a global phenomenon which has drawn the attention of meteorologists for several decades. Ramage (1966) observed a seesaw type inverse relationship between surface pressure anomaly in the ‘heat low’ over the Thar Desert in Pakistan and monsoon rainfall anomaly over the neighboring Arabian Sea off Bombay. He observed that a fall (rise) of surface pressure in Thar Desert over Pakistan was correlated with increase (decrease) of rainfall over the neighboring Arabian Sea. In a theoretical study of the dynamics of deserts and recurrent drought in the Sahel at the southern periphery of the Saharan desert, Charney (1975) observed that to maintain the radiative equilibrium of the atmosphere over a surface with high albedo, such as a sandy soil surface, against radiative heat loss, air must descend and that it is this descent or adiabatic subsidence and warming of the air that leads to continued dryness and maintenance of the desert. He also recognized the contribution of the descending branch of the Hadley circulation to the desertification process but expressed the view that the impact of the radiative heat loss was greater than subsidence warming.
Fig. 1.12 Mean vertical circulations (resultant streamlines) along the Greenwich meridian at 12 GMT during August, showing the monsoon and desert circulations STF marks the location of the subtropical front between the desert and the Mediterranean Sea circulations (after Saha and Saha, 2001b)
1.12
Co-existence of Monsoon with Desert Circulation
31
Blake et al. (1983) who carried out a detailed observational study of heat balance of the atmosphere over a part of the Saudi Arabian desert (Rub-al-Khali) in May 1979 found that there was strong subsidence in the middle and upper troposphere which often descended to very low levels at night. Strong convection over the desert surface during daytime was limited to the lower troposphere only and there was strong outflux of sensible heat in the lower and middle troposphere from the desert to the monsoon over the adjacent Arabian Sea. Saha and Saha (2001b), using NCEP reanalysis data, computed the mean vertical circulation along the Greenwich meridian over the western part of the Saharan desert at 12 GMT during August and showed how the monsoon and the Hadley circulations co-exist with the desert circulations and are linked to one another. Their results are presented in Fig. 1.12. Figure 1.12 shows that in the meridional circulation, while the monsoon winds converge into the ITCZ along about 12◦ N at low levels producing strong penetrative convection which diverges in the upper troposphere both equatorward as well as poleward, there is strong sinking motion in the Hadley circulation over the desert. In fact, the sinking motion over the desert area is so strong that it prevents the thermally-direct low-level desert convection from rising beyond 700 mb. Along the northern boundary of the desert heat low which extends beyond latitude 30◦ N, cool, moist winds diverging from a high pressure area over the neighboring Mediterranean Sea converge into the heat low circulation, forming a subtropical front (STF), the presence of which was first reported by Soliman (1958). In the present text, we have designated the STF as the TCZ
Chapter 2
Tropical Disturbances (Quasi-stationary Waves, Easterly/Westerly Waves, Lows and Depressions, Cyclonic Storms, and Meso-Scale Disturbances)
2.1 Introduction It is widely believed that weather in the tropics is largely seasonal and monotonously of one type; hot and humid during summer and cold and dry during winter. In reality, it is not always quite so, for, the monotony is often broken by different types of tropical disturbances ranging from synoptic-scale wave disturbances that form in the tradewinds that converge into the circulations around the equatorial trough of low pressure, and by severe local storms originating in meso-scale and small-scale disturbances. The synoptic-scale wave disturbances which have a mean horizontal scale of about 1000–3000 km and a Lagrangian time scale of 3–7 days usually move westward with a speed of about 5–8 m s–1 . Some of these disturbances, under favorable conditions, develop into depressions and cyclones. They are mostly thermally driven and powered by the latent heat of condensation of water vapour carried by the tradewinds. For this reason, the most intense of these disturbances form and develop over the warmer parts of the oceans where there is almost unlimited supply of heat and moisture from the underlying surface for their development. They are called cyclones in the Indian Ocean, hurricanes in the Atlantic and typhoons in the Pacific Oceans, though they are fundamentally the same. Out of the large number of lows and depressions that form in the tropics, a limited number only develop into cyclones or severe cyclones and, out of these, only a small percentage develop an ‘eye’ at the center. The severe cyclones are associated with extremely high winds and torrential rain. Some of them cause storm surges which inundate many coastal belts and cause heavy loss of life and property. Many of them breed deadly tornadoes which knock down many standing structures and cause heavy loss of life and property. Tornadoes have drawn the attention of meteorologists from early times and we have an extensive literature on the subject, dealing with their formation, structure, development and movement. Table 2.1 gives a list of the synoptic-scale disturbances classified on the basis of their scale and tangential wind speed, adopted by most meteorological services. Most of the above-mentioned disturbances originate in quasi-stationary waves when they interact with traveling disturbances. Embedded within the abovementioned large- or synoptic-scale disturbances, there are several types of mesoor small-scale disturbances of shorter duration which are very violent in nature. Just K. Saha, Tropical Circulation Systems and Monsoons, C Springer-Verlag Berlin Heidelberg 2010 DOI 10.1007/978-3-642-03373-5_2,
33
34
2 Tropical Disturbances Table 2.1 Classification of synoptic-scale tropical disturbances Classification
Horizontal scale (km)
Range of wind speed (m s–1 )
Low/trough Depression Deep depression Cyclonic storm/tropical storm Tropical cyclone (severe cyclonic storm/hurricane/typhoon)
2000–2500 1500–2500 1500–2000 1000–1500 <1000
<8.5 8.5–13.5 14.0–16.5 17.0–31.5 >32.0
as in midlatitudes, these include thunderstorms, hailstorms, tornadoes, squalls, etc., but seldom fronts. In the present chapter, we first look at the structure and properties of quasistationary wave disturbances that form in zonal currents due to differential heating between land and sea. We then look at the structure and properties of wave disturbances that form and move in the tradewind easterlies and are called Easterly waves. The problem of development of a quasi-stationary trough of low pressure into a depression or a cyclone is addressed next. The role of condensation heating in development is emphasized. It is hypothesized that with enhanced condensation heating mainly on one side of the trough axis and without any appreciable change of structure, a quasi-stationary trough cannot develop further. A transition to a depression or cyclone requires, as a first step, docking of a traveling wave with a stationary wave; a phenomenon which brings about a change of structure with an additional zone of condensation heating. A developing low pressure system requires condensation heating to extend tangentially around the center, gradually turning the quasi-stationary wave disturbance into an axi-symmetric circulation system. Thus, a docking of a traveling wave with a quasi-stationary wave appears to be a prerequisite or necessary first step in development. The role of tropical easterly waves and subtropical/midlatitude westerly waves in development of tropical wave disturbances is discussed in some detail. Some aspects of subsynoptic or meso-scale disturbences and severe local storms are also discussed.
2.2 Quasi-stationary Waves A quasi-stationary wave forms in meteorological fields over several parts of the tropics wherever land – sea thermal contrast exists and is characterized by higher temperature and lower pressure (‘heat lows’) over the warm land and lower temperature and higher pressure (‘cold highs’) over the neighboring cold ocean during summer. The fields are reversed during winter. At low levels, the circulation around the ‘heat low’ has westerlies on the equatorward side of the trough of the heat
2.2
Quasi-stationary Waves
35
low and easterlies on the poleward side. However, in the presence of an easterly thermal wind, the low-level westerlies give way to upper-level easterlies at some height. Both low-level westerlies and upper-level easterlies get involved in wave development process. The existence of a quasi-stationary wave over a region is reflected in several meteorological fields, as shown by a detailed study by Saha and Saha (1996) for the monsoon region of southern Asia. Here we produce evidence from their study in respect of wind and temperature fields.
2.2.1 Quasi-stationary Wave in Wind Field The wave structure in the wind field is revealed best by the zonal anomaly of the meridional component of the wind (v-component) at 850 and 700 hPa along 20◦ N during July, presented in Fig. 2.1. Between the northerly and the southerly flows at both the pressure surfaces in Fig. 2.1, the location of three monsoon troughs, one each over the eastern part of the Arabian sea, the western part of the Head Bay of Bengal and the western part of South China sea can be identified. The mean zonal wavelength and amplitude of the wave in the v-field would appear to be about 2000–2500 km and 4 m s–1 respectively. The wave structure is revealed not only in the lowlevel westerlies but also in the upper level easterlies (not shown), and other parameters.
Fig. 2.1 Zonal anomaly (deviation from zonal mean) of the meridional component (v) of the wind (m s–1 ) at 850 hPa (dashed) and 700 hPa (dotted) along 20◦ N during July. Positive southerly, negative northerly (from Saha and Saha, 1996)
36
2 Tropical Disturbances
2.2.2 Quasi-stationary Wave in Temperature Fields Figure 2.2 shows the distribution of zonal anomaly (deviation from zonal mean) of a 10 year (1976–1985) mean July temperature over southern Asia along latitude
Fig. 2.2 Quasi-stationary wave in a 10 year (1976–1985) mean July temperature anomaly (deviation from the zonal mean) field along 20◦ N over Southern Asia. Upper panel (a) shows the distribution of the anomaly at MSL, 850 and 700 hPa; Lower panel (b) shows the distribution of the anomaly in a vertical section up to 50 hPa (from Saha and Saha, 1996)
2.2
Quasi-stationary Waves
37
20◦ N, at MSL, 850 and 700 hPa (upper panel) and at different pressure surfaces in the troposphere (lower panel). The upper panel brings out the land-sea thermal contrasts with the warm sectors generally appearing over the land and the cold sectors over the sea. The contrasts appear to stand out more clearly in the lower panel which suggests that the stationary wave is basically a phenomenon of the lower troposphere, its upper boundary being at about 500 hPa. It appears to have a wave-length of about 2500 km and amplitude of about 1◦ C in the mean temperature field. Above 500 hPa, the structure of the wave appears to change considerably, as it comes under the influence of the uppertropospheric warm high over Southern Asia which forms part of the planetary-scale monsoon wave at 200 hPa (Kanamitsu and Krishnamurti, 1978). A characteristic feature of the thermal structure above 500 hPa is a pronounced secondary temperature maximum (warm anomaly) situated directly over the lowlevel temperature maximum over the land and a minimum over the low-level temperature minimum over the ocean. Features similar to those presented in Fig. 2.2 were also found along latitudes 15 and 25◦ N, but not along 10, or 30◦ N. The meridional domain of the stationary wave would, therefore, appear to be between 15 and 25◦ N with its strongest intensity along 20◦ N.
2.2.3 Structure of the Quasi-stationary Wave in Circulations Figure 2.3 shows schematically the (a) horizontal (LL-lower level, UL-upper level), and (b) zonal-vertical circulations through a quasi-stationary wave over India (along 20 N) (after Saha and Saha, 1996). 2.2.3.1 Horizontal Circulation The plan view in Fig. 2.3(a) shows the horizontal circulation in the stationary wave at the lower level (LL) and the upper level (UL). Note that in (a), there is flow convergence at the lower level and divergence at the upper level to the west of the low pressure center and divergence at the lower level and convergence at the upper level to the east. 2.2.3.2 Zonal-Vertical Circulation Figure 2.3(b) shows the vertical circulation along the zonal section p, q, r, s, t in (a). While interpreting the vertical circulation in Fig. 2.3(b), it must be borne in mind that the prevailing zonal wind over the region is westerly at low levels which gives way to easterlies above about 700 hPa. This variation of the wind with height is due to the thermal wind being easterly over the region. For this reason, the vertical tilt of the troughs (Tr) and ridges (R) is steeply eastward in the lower troposphere, and slightly westward in the upper troposphere. Further, it is to be borne in mind
38
2 Tropical Disturbances
Fig. 2.3 Schematic showing (a) a plan view of the horizontal (LL-lower level, UL-upper level), and (b) vertical circulations along a zonal section (p, q, r, s, t) in (a). Other symbols have their usual meanings. DV, divergence, CV, convergence
that due to adiabatic changes in the atmosphere, the thermally indirect circulation with warmer air sinking and colder air rising generates available potential energy, while the direct circulation with warmer air rising and colder air sinking increases kinetic energy. Thus, the direct and indirect circulations are intimately coupled to each other and co-operatively drive the observed stationary wave circulation. Note also that deep convection with heavy rainfall (RR) occurs to the west of the low-level trough of low pressure over the region between p and r where the lowlevel convergence is capped by upper-level divergence. Shallow convection with little or no rain occurs to the east between r and t where upper-level convergence is superimposed on low-level divergence. The difference can, perhaps, be understood
2.3
Traveling Easterly (E’ly) Waves
39
from vertical stability considerations as well. Increased adiabatic warming of the sinking air at low levels and adiabatic cooling of the rising air at upper levels make the atmosphere to the west vertically unstable, while increased adiabatic cooling of the rising air at low levels and adiabatic warming of the sinking air at the upper levels to the east tends to make the atmosphere between r and t stable. It is the differential heating between the lower and the upper troposphere that along with the degree of moisture injection may be responsible for the difference in the level of convection between the two sides. A significant characteristic of the synoptic-scale stationary wave circulations revealed by our analysis is their three-dimensional cellular structure in space, consisting of low and high pressure cells in all directions. The idea of a cellular structure of the atmosphere in the vertical involving more than one layer of divergence and convergence and associated vertical motion is not new. It was suggested by studies of Dines (1912, 1919), Sutcliffe (1947), Arnason (1955), Landers (1955) and several others, to account for observed small pressure tendency at surface. Since surface pressure changes observed in the tropics are in no way different from those in midlatitudes, one would argue that the vertical structure of the tropical atmosphere would also be, by and large, similar. The cellular structure appears to exist in both zonal and meridional planes. Quasistationary waves in equatorial/tropical easterlies often interact with their counterparts in subtropical westerlies. Such interactions result in large amplification of the waves in most cases.
2.3 Traveling Easterly (E’ly) Waves Several studies have addressed the problem of easterly waves over the North Atlantic and North Pacific Oceans that travel westward and under favorable conditions give rise to development of hurricanes and typhoons. We shall refer to some of these studies in the pages that follow. There is converging evidence that the E’ly waves originate in a zone of strong cyclonic wind shear in the belt of the tradewind easterlies that converge into the equatorial trough of low pressure over these oceans.
2.3.1 Easterly Waves in Tropical North Atlantic Following early work of Dunn (1940) who noted that the periodic westward movement of disturbed weather across the Caribbean islands in the latitudes of Puerto Rico was associated with the movement of isallobaric lows and highs at mean sea level in regular succession, Riehl (1945, 1954) made detailed synoptic case studies and found that the observed movement of disturbed weather and isallobaric centers at surface was due to the westward movement of waves in the overlying easterly tradewinds. His detailed analysis threw light on the structure and properties of these
40
2 Tropical Disturbances
waves. On the basis of his studies in the Caribbean, Riehl (1945) proposed a model of an easterly wave, the characteristic features of which were the following: (1) The westward passage of the wave over a station is marked by a sharp clockwise turning of the wind direction from NNE to SSE; (2) A marked jump in the depth of the sub-cloud moist layer; and (3) A change in cloud and weather pattern from fair-weather small-scale cumulus clouds to large-scale towering cumulus and cumulonimbus clouds and rain and thunderstorms. Riehl (loc. cit.) showed that the depth of the moist layer, which is related to development of clouds, is directly controlled by the field of motion in the vertical. (a) Where an upper-tropospheric divergent flow is superimposed on a lowertropospheric convergent flow, vertical motion is strongly upward and it leads to a deepening of the moist layer and development of large clouds. (b) A superposition of upper-tropospheric convergence on lower-tropospheric divergence suppresses upward motion and vertical growth of the moist layer. The vertical variation of the basic current can be easily understood in terms of a change of thermal wind. The profile (a) will prevail in regions where the thermal wind is westerly, and (b) where the thermal wind is easterly. A possible implication of a change of thermal wind across a borderline between the two regions in the development of a tropical disturbance will be discussed in a later part of this chapter. Fujita (1971) used cloud motion vectors at low and high levels from satellite cloud imagery around 15 GCT on 14 July 1969 to construct streamlines to demonstrate the presence of easterly waves and their associated cloud systems over the tropical north Atlantic.
2.3.2 Easterly Waves in Tropical North Pacific As in the Atlantic, easterly waves form in the tradewinds over the north Pacific and travel westward. Palmer (1952), Yanai (1961a,b, 1963, 1968a), and several others studied the structure and properties of these waves using different methods. Yanai et al. (1968), Wallace and Chang (1969), Nitta (1970), Wallace (1971) and others used the technique of power-spectrum and cross-spectrum analysis for the purpose. A compositing technique used by Williams (1970) and Reed and Recker (1971) yielded valuable information regarding the average properties of these waves. Chang (1970) used time composite satellite cloud pictures to identify and track these wave disturbances over long distances in the tropical Pacific. A number of workers (e.g., Krishnamurti and Baumhefner, 1966; Yanai and Nitta, 1967; Holton, 1971)
2.3
Traveling Easterly (E’ly) Waves
41
carried out diagnostic studies, using hydrodynamical equations, to infer different characteristics and obtain dynamically consistent pictures of these waves. Reed and Recker (1971) summarized the average structure and properties of easterly waves in the equatorial western Pacific region as follows: the meridional wind maxima of nearly opposite phase occurred in the lower and upper troposphere. Negative temperature deviations were found in the vicinity of the wave trough at low and high levels; positive deviations were observed at intermediate levels. Highest relative humidities occurred in the trough region. This was also the region of strongest upward motion and maximum cloud amount and rainfall. The maximum upward velocity of 2.5 cm s–1 was found at 300 hPa. Convergence was strongest in the sub-cloud layer; divergence was concentrated near 175 hPa. The maximum anticyclonic vorticity was also observed at that height. Besides the above, there were several other properties of these waves that were revealed by their study. According to theoretical reasoning (Riehl, 1954) and numerical modeling (Holton, 1971), it was apparent that the wave structure would be sensitive to the vertical profile of the basic current in which the wave was embedded in addition to the horizontal profile. This was actually found to be so in the equatorial western Pacific where the vertical structure of these waves changed systematically as they travelled from the easternmost to the westernmost islands from near the dateline to a longitude of about 135E. The wave axis, which tilts eastward with height at the eastern end of the islands, becomes vertical in the central region and acquires an opposite tilt in the west. The change in the tilt is believed to be caused by the variation with longitude of the vertical shear of the basic current, as shown in Fig. 2.4. It may be seen that the vertical wind shear at Koror (7.5N, 135E) which lies at the western end of the network is very different from that at Majuro (7.0N, 171E) which lies near the eastern end and also from that of the mean of three island stations which lie in between. Yanai (1964) suggested a model of the easterly wave based on his studies in the equatorial eastern Pacific, which is shown in Fig. 2.5. The upper panel shows a side view of the wave with isolines of vorticity (thin continuous lines) and the directions of vertical motion (double-shaft arrows) on the left and vertical distribution of the basic current on the right. The lower panel shows a horizontal view of the wave with streamlines (solid arrows) and isotachs (thin dotted lines) on the left, and the latitudinal distribution of the basic current at 850 hPa on the right.
2.3.3 Easterly Waves in the Indian Ocean Region Easterly waves constitute a prominent feature of the atmospheric circulation over the Indian Ocean monsoon region. They are observed in quasi-stationary as well as traveling modes (For quasi-stationary wave mode, see Sect. 2.2).
42
2 Tropical Disturbances
Fig. 2.4 Mean zonal wind speed over western North Pacific for the period July–September, 1967. The profile labelled KEP is the mean for Kwajalein (9.0N, 167E), Eniwetok (12N, 162E) and Ponape (7.0N, 157.5E). The thin vertical line denotes the average wave speed (–9 m s–1 ) observed during the period of study (after Reed and Recker, 1971)
The traveling E’ly waves usually originate over the Pacific Ocean and arrive over the Indian Ocean region after crossing the SE Asia peninsula. However, south of the equator, they originate in the extreme eastern parts of the Indian Ocean adjoining the northwestern coast of Australia. They usually move westward along the south equatorial trough of low pressure between the equator and 15◦ S. Traveling E’ly waves that enter the Northern Indian Ocean region from the east appear to have more or less the same structure as the stationary E’ly wave over the region, shown in Fig. 2.3. However, their vertical structure varies considerably depending upon the prevailing wind structure. An inclined quasi-stationary monsoon trough appears to offer the most favorable orientation for development of the trough into a depression, for, it provides, inter alia, an east–west component of the circulation associated with it to interact with an approaching E’ly wave. Where the trough is under influence of two opposing forces, it gets inclined in the direction of the more powerful heat low. In the Indian Ocean region, inclined monsoon troughs may be found on both sides of the equator. In the Arabian Sea sector, development is favored when the trough is inclined in a NE–SW
2.4
Development of Waves
43
Fig. 2.5 Model of an easterly wave in tropical eastern Pacific (after Yanai, 1964)
direction, while an inclination in a NW–SE direction favors development in the Bay of Bengal sector.
2.4 Development of Waves 2.4.1 Meaning of Development Before we discuss the problem of development of a trough of low pressure into deeper pressure systems, such as a depression or a cyclone, it may be desirable to clarify the meaning of the term ‘development’. The development of a low pressure system means a faster and greater fall of pressure at the center of the low pressure system relative to its surroundings, as measured by the Laplacian of pressure tendency, ∇ 2 (δp/δt). Alternatively, it can also be defined by an increase in the absolute vorticity of the circulation, i.e., by dη/dt, where η is the absolute vorticity and the operator d/dt denotes change following motion. If the system is stationary, d/dt can be replaced by ∂/∂t, the local tendency. It can be easily shown by integrating the hydrostatic equation that the pressure tendency at any given location depends upon the net flow divergence and/or mass advection in the overlying column of air, as shown below:
44
2 Tropical Disturbances
The pressure ps at the earth’s surface (z=0) is given by ∞ ps =
ρgδz, 0
where ρ is density of air, and pressure is assumed to be zero at height infinity. Differentiating ps with respect to t, and using the continuity equation, ∂ρ/∂t = −∇ · ρV we get the pressure tendency equation in the form ∞ ∂ps /∂t = −
∞ ∇H · (ρV)gδz −
0
{∂(ρw)/∂z}gδz 0
where V is the horizontal velocity vector, w is vertical velocity, ∇ H is horizontal Del operator and the first term on the right-hand side gives the contribution of horizontal mass convergence or divergence through the sides of the column, while the second term gives the contribution of vertical mass convergence or divergence through the base of the column. The second term may be simplified to g(ρw)0 , since ρ and w are both zero at infinity. We may, therefore, write the pressure tendency equation in the simplified form ∞ ∂ps /∂t = −
∇H · (ρV)gδz + g(ρw)0 0
Thus, the pressure tendency at a place is negative or positive, according as there is net mass divergence from, or convergence into, the overlying air column. Lines drawn through places of equal pressure tendency are called isallobars. According to Table 2.1, a low pressure system can develop into many different stages, ranging from a depression to a most intense hurricane. It is estimated (Frank, 1971) that only about 10% of all tropical disturbances reach storm strength and about one in three depressions makes the transition to storm or hurricane. Further, the formation of an ‘eye’ in a hurricane marks a most crucial final stage in the transformation of a tropical disturbance. Most tropical disturbances (90%) fail to form the eye and never reach the highest intensity. Further, it is observed worldwide that most disturbances develop from a pre-existing trough of low pressure, such as that found in association with quasi-stationary monsoon troughs discussed in the preceding section and that after formation and development they usually move along the axes of the monsoon troughs. As a quasi-stationary feature of the tropical circulation, the quasi-stationary wave is a dynamically stable system but may get destabilized whenever a traveling Easterly wave interacts with it.
2.4
Development of Waves
45
2.4.2 Development of a Quasi-stationary Monsoon Trough into a Depression 2.4.2.1 Role of Tropical E’ly Waves In the past, mechanisms such as barotropic, baroclinic, or combined CISKbarotropic-baroclinic instability have been suggested to explain the formation of monsoon depressions. However, observations show that inspite of the presence of these types of instabilities in the mean monsoon atmosphere, it is only occasionally (i.e., once or twice a month) that a depression forms and only a small percentage of the depressions develop further to storm strength. So, what actually triggers the formation of a monsoon depression in the first place? Some past studies (e.g., Koteswaram and George, 1958; Saha et al., 1981) suggested the involvement of easterly waves in the process. Koteswaram and George (1958) suggested that when upper-air divergence ahead of an approaching easterly wave trough was superimposed upon the convergent area of the low-level monsoon trough, the trough would develop into a depression. Saha et al. (1981) found that in a 10-year (1969–1978) period, out of the 52 depressions that formed over the Bay of Bengal during July–August, 45 (87%) could be traced to an interaction with westward-propagating easterly waves which moved in from areas to the east. When a traveling easterly wave approaches the quasi-stationary monsoon wave and docks with it, the docking strengthens a pre-existing zone of subsidence to the east of the trough axis by added subsidence of the indirect circulation of a new zone of condensation heating created to the east of the subsidence zone. Enhanced subsidence leads to increased adiabatic warming and a fall of surface pressure immediately to the east of the trough axis. Thus, the initiation of condensation heating on either side of trough axis with a zone of enhanced subsidence warming in between produces a complete depression wave. The depression circulation is centered immediately to the east of the monsoon trough axis. The development of the quasi-stationary trough into a depression is accelerated if the trough is oriented in a NW–SE direction and a cross-equatorial flow from the winter hemisphere converges into the TCZ. The involvement of cross-equatorial flow in development of a monsoon depression was suggested by several studies in the past (e.g., Sikka, 1980b, Sikka and Gray, 1981; Saha and Saha, 2004a,b) in the case of the monsoon trough over the Bay of Bengal, though without mentioning the branch of the cross-equatorial flow likely to be involved. However, it is important to note here that with a monsoon trough oriented in a NW–SE direction, it is only the Bay of Bengal branch of the cross-equatorial flow which can lead to cyclogenesis of the trough; not the Arabian Sea branch which activates the ITCZ. A simultaneous activation of both ITCZ and TCZ will, of course, accelerate the development. We may visualize the development process as illustrated in Fig. 2.6. There is enough observational evidence from different parts of the tropics, which appears to support the conceptual model of the formation of a monsoon depression, suggested above and shown schematically in Fig. 2.6. Rao (1976) cites numerous
46
2 Tropical Disturbances
Fig. 2.6 Schematic illustrating the formation of a monsoon depression in a quasi-stationary monsoon wave trough by interaction with a traveling E’ly wave
examples of typical cloud distributions in the field of a monsoon depression over the Indian monsoon region, which strongly support the structure suggested in Fig. 2.6. Several theoretical and numerical experiments (e.g., Charney and Eliassen, 1964; Krishnamurti et al., 1976; Shukla, 1978) on the formation of a monsoon depression have concluded that the eddy available potential energy required for development of a pre-existing perturbation in the monsoon current is, in all probability, supplied by condensation heating. The conceptual model presented here suggests a mechanism which can bring about the additional zone of condensation heating required for the development. 2.4.2.2 Role of Subtropical W’ly Waves Like an easterly wave, a wave in subtropical/midlatitude westerlies can initiate the development of the quasi-stationary monsoon trough. It is well-known that W’ly waves in the course of their eastward travel often develop large amplitudes when they reach the longitudes of the quasi-stationary monsoon wave and interact with it across the subtropical belt in the upper troposphere (see Fig. 2.7). The interaction produces extended troughs and ridges when the waves get into the same phase. During interaction, an exchange of energy occurs between the waves in which cold subsiding and adiabatically-warming air from higher latitudes move
2.4
Development of Waves
47
Fig. 2.7 Interaction of the quasi-stationary monsoon trough with a W’ly wave trough in the upper troposphere in the northern hemisphere (trough locations are shown by thick dashed lines)
equatorward to intensify the subsidence warming to the west of the monsoon trough axis and warm rising and adiabatically-cooling air from lower latitudes move poleward to intensify the rising currents to the east of the midlatitude W’ly wave trough axis. So, when, after coupling, the waves separate, and the westerly trough moves away eastward, a fracture occurs in the extended trough axis and a cut-off low pressure forms around the easterly wave trough, which may subsequently develop into a monsoon depression. It should, however, be noted that in this interaction, it is the TCZ which gets involved and activated and becomes a zone of penetrative convection, clouding and precipitation; not the ITCZ on the southwestward side of the trough axis. Thus, as a result of this interaction, we have two zones of condensation heating with a zone of enhanced subsidence warming between them, which are favorable for the formation of a monsoon depression. Here also, as in the case of interaction with an easterly wave alone, the depression is likely to be centered immediately to the east of the low-level monsoon trough axis. The arrival of a cross-equatorial flow from the winter hemisphere to the trough zone at low levels at this time may lead to explosive cyclogenesis. There is evidence to suggest that eastward-propagating large-amplitude W’ly waves interact with quasi-stationary monsoon troughs over all continents, leading to formation of monsoon depressions. Over the Indian subcontinent, the W’ly waves which move across the Himalayas and the Tibetan plateau influence the formation, development and movement of monsoon depressions over the region, especially the Bay of Bengal (e.g., Saha and Chang, 1983). Summarizing the above paragraphs, we may state that both E’ly and W’ly waves can contribute to the development of a quasi-stationary monsoon trough into a depression. Explosive development may result if a cross-equatorial current from the winter hemisphere is also drawn into the circulation around the monsoon trough at low levels to activate the ITCZ and the TCZ during the interaction. The low frequency of such a favorable combination may, perhaps, explain why a monsoon
48
2 Tropical Disturbances
depression forms so infrequently, usually one or two a month, though both easterly and westerly waves interact with the quasi-stationary monsoon trough more frequently.
2.4.3 Development of a Depression into a Cyclonic Storm/Tropical Cyclone/Hurricane/Typhoon While interaction with a traveling wave appears to be an essential first step in the formation of a depression, further development of a depression into a tropical storm or tropical cyclone appears to be a long drawn-out and complicated process requiring the following pre-requisites for further growth. ∗ A depression must be in an environment which would favor condensation heating by sustained moisture convergence at low level over a long period of time. The following environmental conditions have been found to be necessary for the formation of a cyclonic storm or tropical cyclone; ∗ Pre-existence of a warm land, or an ocean with SST exceeding 26.5◦ C, and a closed low or trough of low pressure which ensures cyclonic circulation, flow convergence and positive relative vorticity at low levels; ∗ A deep mixed layer of the ocean with continuous availability of water vapour either through surface evaporation or advection of moist air, which ensures strong moisture convergence at low levels; ∗ Existence of conditional instability and weak vertical wind shear in the overlying atmosphere; ∗ Strong upper air divergence above low level convergence, which ensures penetrative convection, rapid build-up of large cloud clusters and precipitation and release of latent heat of condensation; such conditions are usually ensured in monsoon troughs which are inclined to latitudes, as depicted in Fig. 1.5; ∗ Release of enough latent heat of condensation to enforce strong subsidence warming and rapid fall of surface pressure near the center of the depression. ∗ For a tropical cyclone with an eye to form, condensation heating must extend all around the center of the depression. In other words, the cyclone must be axi-, or quasi-symmetric. Furthermore, the development of a depression involves a reduction in the horizontal scale of the disturbance to no more than 30% of the original scale (see Table 2.1). How is this reduction brought about and with what effect on the physical and dynamical characteristics of the depression in the process? These are important questions to which we can provide only tentative answers at present, since the whole process is nonlinear and there are many dissipative processes at work. In fact, the process to full axial symmetry is a long and complicated one and there is little guarantee that a given disturbance will mature to its fullest extent. However, the process
2.4
Development of Waves
49
may be visualized as follows: The introduction of two zones of condensation heating on either side of the center of a low or depression induces increased subsidence warming and fall of pressure near the center. In response to the pressure fall, air from the outer boundary region starts flowing inward. The inflowing air becomes increasingly convergent with strong upward motion and development of towering clouds and precipitation. Another consequence of the inward flow is a rapid strengthening of the tangential wind, because the inflowing air is required to conserve absolute angular momentum. The scale of the disturbance thus gets gradually reduced till it stabilizes at some minimum radial distance from the center, where the pressure gradient force is balanced by the centrifugal, Coriolis and the frictional forces. This radial distance marks the inner boundary of a broad zone of maximum tangential wind, strongest updraft, and torrential precipitation, which constitutes the eye-wall. In the course of the development process, the cloud zone which was originally confined to two narrow zones on either side of the depression center gradually extends along the circumference till the disturbance attains its nearly full axial symmetry with potential to develop an eye at the center. However, the process of formation of an eye in a tropical cyclone appears to be a complicated one. It is discussed further in Chap. 3, which deals specifically with the formation, structure and properties of tropical cyclones. It follows from the foregoing considerations that depressions which have limited sea travel or which move over cold oceans have little chance to develop further. Also, the requirement of a weak vertical shear of the basic wind appears to be a very crucial one, for, it is important that the cool, moist air that converges at low levels in a trough zone must be lifted by upper level divergence as close to the trough zone as possible and that this can only happen in a low-level strongly convergent zone, such as the TCZ, as discussed in Chap. 1. Gray (1968) in a worldwide survey of tropical cyclones for a 20-year period identified the initial genesis points of cyclones (Fig. 2.8). According to this survey, the areas most susceptible to formation of tropical cyclones are: (1) a wide
Fig. 2.8 Initial genesis points of tropical cyclones during a 20-year period (Gray, 1968)
50
2 Tropical Disturbances
belt of ocean extending from the eastern Arabian Sea to the Western Pacific Ocean up to the dateline north of the equator; and (2) a more extensive belt extending from the coast of South Africa to about 160◦ W south of the equator. Two other major cyclone-breeding areas lie in the Western Atlantic Ocean and Eastern Pacific Ocean in the northern hemisphere. It is observed that in all the areas, development occurs around inclined troughs of low pressure. Undoubtedly, one may expect to see many more genesis points than found by Gray’s survey after introduction of earth-orbiting satellites. For example, a tropical cyclone formed over the southwestern part of the Atlantic Ocean off the coast of Brazil in March, 2004, which was visible in a satellite cloud imagery, but not reported by conventional observational systems.
2.5 Meso-Scale Disturbances and Severe Local Storms in the Tropics 2.5.1 General Considerations – Source of Energy of the Storm Besides large and synoptic-scale wave disturbances discussed in the foregoing sections, several types of subsynoptic- or meso-scale disturbances with diameters ranging from a few hundred meters to about 100 km, such as thunderstorms, hailstorms, tornadoes and squall-lines, also form in the tropical atmosphere. Most of them form in a conditionally unstable atmosphere within the synoptic-scale disturbances and develop by mutual co-operation between the large and the small scales. A small-scale disturbance with shallow convection may form in a conditionally unstable environment when it is heated from below or cooled from above, but its further growth in most cases depends upon the support of the synoptic- or large-scale disturbance which provides the necessary mechanism for its development. Thus, the basic requirements for development of a meso-scale tropical storm in any locality appear to be largely the same as enunciated in the case of development of a tropical cyclone in Sect. 2.4.3, viz, (i) The existence of a conditionally unstable atmosphere; (ii) Moisture convergence at low levels; and (iii) An effective mechanism to raise the low-level moisture to higher levels and release the latent heat of condensation for further growth of the cloud through penetrative convection, the net effect of which is an adiabatic warming of the environment by subsidence. Normally, depending upon the extent of mutual co-operation between the cloud and its environment, a small-scale cloud may grow to different stages and display different physical and dynamical characteristics. For example, it may grow to a stage where it gives only thunder and lightning, but no high wind, rainfall or hail at the ground. On the other hand, it may develop to a stage where it displays all these characteristics and, perhaps, additionally also a highly developed tornado vortex.
2.5
Meso-Scale Disturbances and Severe Local Storms in the Tropics
51
When fully developed, most of them make a significant contribution to the total precipitation and the circulation and weather phenomena in the tropics, apart from its impact upon society. In this section, we discuss some aspects of their formation, internal structure, electrical properties, and other physical and dynamical properties to the extent revealed by observations from several field projects, such as the well-known thunderstorm project (Byers and Braham, 1949) in USA. In fact, these experiments have largely confirmed some of the basic characteristics of atmospheric convection, such as strong updrafts inside the cloud cell, entrainment and detrainment of air, and net warming of the environment by subsidence.
2.5.2 Thunderstorms A thunderstorm derives its name from its characteristic property of producing thunder and, inevitably, lightning in the sky besides stormy conditions including high, squally winds and sudden heavy downpour at the ground. Occasionally, it may turn into a hailstorm or even breed tornadoes. Because of its association with thunder and lightning, it is also sometimes called an ‘electrical storm’, though the term is normally used in the context of space weather. Generally, it appears in the sky as a giant cumulonimbus cloud with a large spread-out anvil at the top and occasional lightning flashes. However, when it is far away from a station, one may see lightning flashes only, but hear no thunder. But, as it draws near, lightning flashes become more frequent and one may also hear thunder from it. Since light waves travel a million times faster than sound waves through the air, the difference in their times of arrival gives an indication of the distance of the storm from a station at any time. Severe weather usually follows when the storm is directly over a station. A thunderstorm derives its kinetic energy from the available potential energy stored in a conditionally unstable atmosphere in which the equivalent potential temperature decreases with height up to at least 3 km, and increases aloft. In such an atmosphere, cloud growth is limited to small-scale cumulus clouds only. Their further growth to higher levels is prevented by the strong static stability of the upper troposphere. In the tropics, heat thunderstorms usually develop in late spring and early summer months. During that period, the temperature at the ground and the air layer immediately above it goes through a diurnal cycle, with a temperature inversion near the ground during late night and early morning hours and a steep temperature lapse rate developing during the daytime. So, it is during the late hours of the day that conditions become favorable for penetrative convection and rapid development of the cloud, resulting in the formation of a large cumulus or cumulonimbus cloud. 2.5.2.1 Cellular Structure of a Thundercloud and Vertical Currents When viewed from outside, a thundercloud may appear as a single large cloudmass with a diameter of a few kilometres (usually 10–12 km) and top rising to great heights, sometimes even beyond the tropopause at 16–18 km a.s.l., and lasting for a few hours. This may be compared with an individual fair weather cumulus
52
2 Tropical Disturbances
cloud which has an average diameter of about 1 km, a height of 1–2 km, and a life span of 15–20 min only. In reality, and as found by several aircraft probes during a thunderstorm project (Byers and Braham, 1949), there are several individual cloud cells inside the large cloudmass, all almost merged into one another in some kind of a conglomerate under a single canopy, each having a short lifespan of its own while being replaced by a new cell. Thus, the cloud cells inside a large developing cumulonimbus cloud are in violent agitation, somewhat like rising bubbles in a boiling cauldron. The rise and fall of the seething cloud cell tops can be clearly seen from an aircraft flying above the clouds. It is found that during the developing stage, updrafts dominate over downdrafts inside the cloud. But when precipitation starts falling through the cloud downdrafts increase and ultimately dominate over updrafts, thereby signalling the demise of the thunderstorm. Direct observations or calculations of vertical air velocity inside a growing cumulus or cumulonimbus cloud are scanty. Fortunately, the Thunderstorm Project, referred to in the preceding para, studied the structure and circulation patterns of thunderclouds in various stages of development and found by direct aerial measurements that there is only updraft in the interior of these clouds till the precipitation stage is reached when updraft is replaced by downdraft in the portion of the cloud through which rain or snow falls. The general distribution of updraft velocity with height in the central portion of a growing cumulus cloud and a mature cumulonimbus cloud and the radial distribution of updraft velocity from the central region to the periphery of a cloud of horizontal extent of 8 km at the height of maximum updraft at about 9.0 km are shown in Fig. 2.9(a) and (b) respectively. The vertical distribution Fig. 2.9(a) shows that there is a gradual increase of updraft velocity from a value of about 5 m s–1 at height 1.5 km to 20 m s–1 at 8–10 km and then a gradual decrease to low velocity aloft. The radial distribution (b) shows that the maximum updraft of about 18–20 m s–1 occurs near the center of
Fig. 2.9 (a) Distribution of updraft velocity (m s–1 ) with height in the center of a growing large cumulus cloud (thin line) and a fully mature cumulonimbus cloud (thick line); and (b) the radial distribution of updraft velocity at the level of maximum updraft (8–10 km a.s.l.) (Byers and Braham, 1949; Saha, 1962)
2.5
Meso-Scale Disturbances and Severe Local Storms in the Tropics
53
the storm with a gradual decrease of the velocity to a value of 2–3 m s–1 at about 3.5 km from the center. 2.5.2.2 Hydrometeors Inside the Thundercloud The updrafts in a thundercloud lift moisture to different levels of the atmosphere leading to formation of water substances in different phases of water. In the first stage, as moist air is lifted to levels above the condensation level at about 1.5 km, water vapour condenses on the existing cloud condensation nuclei to form cloud drops. Some of the drops increase in size as they are carried upward by collecting smaller drops and deposition of new water vapour round them but they remain in liquid form till they reach the freezing level at about 5 km. Above the freezing level, drops don’t freeze immediately, since they can remain in a supercooled state to a temperature as low as –40◦ C. However, as they are lifted by strong updrafts to the cooler regions of the upper troposphere, an increasing number of them turn into snow crystals. Beyond a height of about 10 km, most of the snow crystals turn into ice crystals. In fact, the anvil part of a thundercloud near the tropopause consists mainly of ice crystals. 2.5.2.3 Precipitation, Downdraft and Squalls The water substances inside the storm in the form of supercooled water drops, snow and ice crystals, after they grow to large sizes, start falling through the cloudmass under their own weight with a velocity which depends upon the updraft velocity and the aerodynamic properties and resistance of the environmental air through which they fall. As they fall through the freezing level, the ice and snow crystals begin melting by drawing heat from the warmer environment and thereby cool the air through which they fall by evaporation from their surfaces. The result is a downdraft of extremely cold air in the precipitating part of the cloudmass. On reaching the ground along with large drops of rain and sometimes pellets or hailstones, the downdraft spreads in all directions producing squally winds. In meteorology, a squall is defined as a strong wind characterized by a sudden rise of its velocity to 16 knots or more which is sustained for at least two minutes. Squalls of abnormally high winds can do a lot of damage to lightly-built structures, uproot trees, and knock down living creatures out in the open. 2.5.2.4 Lightning and Thunder Lightning and thunder are two most characteristic properties of a thunderstorm. Lightning is an electrical discharge that occurs between two oppositely charged regions inside a cumulonimbus cloud, or between the cloud and the earth’s surface. The basic requirement is a separation of electric charges between two parts of the cloud and build-up of a strong electrical field between them. By induction, the field produces a path of ion-molecules through which the discharge occurs.
54
2 Tropical Disturbances
More than half of all lightning discharges take place inside the cumulonimbus clouds and are known as intracloud discharges. But, lightning also occurs between the cloud and the ground which are known as streak or forked lightnings. Because of its intrinsic interest in the context of harmful effects upon living creatures by way of death and destruction, ignition of forest fires, and disturbances to power and electrical communications, etc., the cloud-to-ground lightning has been studied more extensively than any other form of lightning. Most of the electrical energy in a lightning discharge goes into producing heat. It is estimated that air along the path of a lightning discharge may be raised to a temperature exceeding 10,000◦ C. A part of this heat energy goes into producing radiation and light, while a part goes into producing sound waves through longitudinal compression and rarefaction which we hear as thunder. Lightning activity over the globe is believed to contribute importantly to the maintenance of the earth’s electric field. It is well-known that the ionosphere which is positively charged loses its charge continuously to the ground which is negatively charged. This implies that unless replenished regularly, the ground will soon lose its negative charge and the atmosphere will lose its electric field. It is believed that this requirement is fulfilled by lightning activity over different parts of the globe, which supplies the much-needed negative charge from the cloud to the ground. The cloud also sends positive charge from its upper part to the ionosphere. The harmful effects of lightning strikes can be minimized by keeping indoors during a thunderstorm. It
Fig. 2.10 Schematic showing the structure of a thunderstorm (after Newton, 1966). Arrow shows the direction of air motion. Symbols: A rectangle represents an ice pellet, a star a snowflake, and a circle a raindrop. F.L. stands for freezing level. Double-shaft arrow shows the direction of the wind believed to be steering the storm
2.5
Meso-Scale Disturbances and Severe Local Storms in the Tropics
55
is a common practice in most parts of the globe to erect lightning rods in buildings in order to protect them from possible damage from lightning strikes. 2.5.2.5 Structure of a Thunderstorm Figure 2.10, after Newton (1966), is a sketch showing some of the many abovementioned features of a mature thunderstorm, the top of which shoots up to a level well above the tropopause. It is likely that the structure will differ if the top remains below the tropopause level, but it is believed that the main features shown will remain similar. 2.5.2.6 Squall Line A squall line is a line of active thunderstorms which may be continuous or with breaks over a length of 100 km or more. It is accompanied by all the characteristics usually associated with a thunderstorm, viz., strong updrafts and downdrafts, precipitation, lightning, thunder and squalls. It is a type of meso-scale convective severe local storm system which is usually associated with moving winter cold fronts and summertime troughs of airmass discontinuity. It is distinguished from other types of severe local storms by a smaller width-length ratio.
2.5.3 Hailstorms A hailstorm is a mature thunderstorm from which significant hailstones fall to the ground. The size of hailstones can vary from 5 mm to 15 cm or more in diameter. Bilham and Relf (1937) quote a report of a severe hailstorm at Porter, Western Nebraska in USA, in 1928, in which the largest of the round hailstones was as big as a large grapefruit. It measured 43 cm in circumference indicating a diameter of about 13.7 cm, had a density of about half of that of water and weighed 681 g. A cumulonimbus is the only type of cloud that is known to yield hailstones. Hailstorms are of intrinsic interest because of their destructive power. They are known to inflict damage to aircraft in flight and standing crops on the ground. They can also destroy plant and animal life. In tropical countries, such as India and Nigeria, hailstones can be of even larger sizes. There was once a report of a hailstorm in Rajasthan in India which wiped out a big herd of cattle comprising of giant-size buffaloes grazing in the field. On 27 May, 1959, an Indian Airlines aircraft flying at 5700 m a.s.l. near Delhi in India was caught up in a hailstorm and the hailstones encountered caused holes in the aircraft, the largest of which had a diameter of about 15 cm (Saha, 1962). The clean-cut of the holes which were practically spherical in shape would give an impression that the hailstones that caused these holes were of comparable size, although the impression may not be quite justified in the absence of data based upon relationship between the size of the hailstone and that of the hole in different ranges of size and energy of the hailstone. Wegener (1911) recorded that the largest known hailstone had a weight of 1 kg and the size
56
2 Tropical Disturbances
of a skittle ball (Kegelkugel), indicating a diameter of about 14 cm if a mean density of 0.7 g/c.c. were assumed. Because of the potential dangers attached to hailstones, they have drawn the attention of meteorologists from early times. However, little definite is known as to how hailstones form and grow inside a thundercloud. Several theories and hypotheses have been advanced in this regard. For a summary of some of these studies, the reader is referred to an article by the author (Saha, 1962). A plausible theory of the formation of large hailstones must take into account the following facts of observation, reiterated below: 1. A hailstone consists of concentric shells of opaque and clear ice. 2. It forms in convective clouds the tops of which reach great heights in the atmosphere. In the tropics, the tops of cumulonimbus clouds have been known to extend to 12–5 km a.s.l. or even greater heights. 3. The mean distribution of updrafts in towering cumulus and cumulonimbus clouds is such that there is a gradual increase of updraft velocity from a value of about 2.5 m s–1 at height 1.5 km to about 10 m s–1 at 8–10 km and then a gradual decrease to lower velocity aloft (Byers and Braham, 1949). 4. All thunderstorms do not precipitate hailstones, although it is generally agreed that the evolution of an ice phase is a common feature of all thunderclouds. Observations show that only a very small proportion of even the severest thunderstorms yield large hailstones. 5. There is a pronounced geographical and seasonal variation in the distribution of hailstones. In general, thunderstorms occurring in particular areas and seasons only yield hailstones. 6. In the size-spectrum, hailstones of very large diameter are found to occur in low concentration (1/m3 or less). 7. Most often, hailstones are observed to fall earlier than raindrops. Sometimes they fall simultaneously. 8. Hail starts rather suddenly and usually lasts from a few minutes to less than half an hour. 9. There may be several ups and downs inside and outside the cloud before a hailstone comes to the ground. 10. The presence of a large thickness of cloudmass between the base of the cloud and the level of freezing of large clouddrops prevents the growth of hailstones by prematurely precipitating out the larger drops in the form of heavy rain. There have been several theories regarding the formation of large hailstones in a hailstorm, but one that takes care of most of the above-mentioned facts of observation and rather promising is one that was advanced by Schumann (1938). He gave the following equation for the ultimate diameter of a hailstone: L = (σ/m)(2D − 6) − (4u/m)(3κρσ/πg)1/2 × {(2D)1/2 − 2.45}
(2.5.1)
2.5
Meso-Scale Disturbances and Severe Local Storms in the Tropics
57
where L σ m u κ ρ g D
is the depth of fall of the hailstone above the freezing level, the average hailstone density, the concentration of condensed water, the updraft velocity, the Reynolds number, the air density, the acceleration due to gravity, and the ultimate diameter of the hailstone, the initial diameter being assumed 3.0 cm at height 9.0 km a.s.l. which is the mean height of maximum updraft
Using the above equation and the following values (in c.g.s. system) of the abovementioned parameters, L = 4.5 × 105 cm, σ = 0.7 gm (c.c)–1 , m = 20 × 10–6 gm (c.c)–1 , u = 2000 cm s–1 , κ = 0.1, ρ = 5 × 10–4 gm (cm)–3 , and g = 981 cm s–2 , Schumann obtained a value of 12.1 cm for D. The present position regarding these theories is that several questions relating to formation of large hailstones remain unanswered. In the absence of direct observations on many of the parameters involved, it is rather difficult to discuss the relative merits and demerits of the various theories that have been advanced from time to time. It is, however, generally agreed that strong updraft promotes growth of large hailstones. But it needs to be emphasized that the size of a hailstone that is balanced by updraft is the maximum size with which it leaves the cloud or reaches the ground. The growth of a hailstone beyond the size which can be supported by updraft occurs while it descends through a large mass of supercooled drops and snowflakes in the supercooled region of the cloud. This appears to be an all-important point in any plausible theory of formation of large hailstones. Regarding the observed structure of hailstones consisting of alternate shells of white ice and clear ice, there have been two main schools of thought. One led by Humphreys (1940) holds the view that the formation of alternate shells of opaque and clear ice is due to the alternating movements of a frozen cloud or rain drop between the realms of rain and snow across the freezing level by vertical aircurrents. The other theory, due to Gaviola and Fuertes (1947), seeks to explain the observed structure by assuming that while the surface is wet, a transparent layer forms, and when it dries, an opaque white layer forms.
2.5.4 Tornadoes A tornado is a mesoscale violently-rotating atmospheric vortex protruding downward from the base of a large cumulonimbus cloud in the shape of a funnel which often reaches the ground with disastrous effects on life and property. It occurs mostly during passage of a heavy thunderstorm or squall line over a locality. Because of its
58
2 Tropical Disturbances
funnel shape, it is also sometimes called a funnel cloud. When appearing over water, it is called a water spout. The typical characteristics of a full-grown tornado vortex as revealed by observational and theoretical studies are the following: (a) A diameter of 100–200 m, a variable depth below the cloud base with intermittent touch-downs, a life-span varying from an hour to several hours, and an erratic path varying in length from 1 to 100 km; (b) Revolving winds with high positive relative vorticity and updraft locally inside the cloud, as found in a helical vortex; (c) A maximum tangential velocity, 100–150 m s–1 ; (d) A maximum updraft velocity, 100–150 m s–1 ; (e) Extremely low pressure and high temperature inside the vertical column; (f) Entrainment of environmental air at explosive rate; (g) Frequent lightning flashes and increasingly thunderous and roaring sound at the approach of a tornado; (h) Often heavy precipitation following the passage of a tornado. It should, however, be emphasized that there have been few direct observations made inside a live tornado, because of the obvious dangers to human lives and inability of any meteorological instruments to withstand the impact of the high revolving winds. This means that most of our current information about tornadoes is derived indirectly from detailed post-mortem examination of death and destruction it leaves behind. [The author once investigated a tornado which occurred in the district of Cooch Bihar in eastern India in 1963. The indirectly estimated values of some of the abovementioned characteristics are quoted mostly from the findings of that study.] 2.5.4.1 Tornado Circulation and Intensity Circulations in a tornado, both horizontal and vertical, are sometimes so strong that they can uproot large trees, bend electric poles, knock down buildings and houses, wipe out crop fields, pick up heavy objects from the ground and throw them around like deadly projectiles, derail running trains, sink boats and ships plying in rivers, and occasionally lift boats from river banks and throw them into water, and so on. There was a report of a tornado which while moving over a house in eastern India picked up a young girl of about 10 years from the compound of the house and later deposited her safely a few hundred metres down the path. Further down, it sucked up all the muddy water from a wide canal 1.2 m deep and deposited it all along its route. The direction of the circulation and probable speed of the horizontal and vertical components of the wind are usually estimated from the extent and intensity of these damages and destructions. Casualties to human and animal life are usually caused by collapse of poorly-built houses and by their exposure to heavy metallic projectiles when exposed in the open. In USA, windspeeds are sometimes estimated on the basis of observed damages using a scale known as the Fujita scale (Fujita,
2.5
Meso-Scale Disturbances and Severe Local Storms in the Tropics
Fujita-scale Scale of damage
Speed limits (m s–1 ) (horizontal)
F-0 F-1 F-2 F-3 F-4 F-5
18–32 33–49 50–69 70–92 93–116 117–142
Light damage Moderate damage Considerable damage Severe damage Devastating damage Incredible damage
59
1981). The Fujita scale (also known as F-scale) is a six-point scale which, according to the Meteorological Glossary (second edition) of the American Meteorological Society (2000), corresponds to the following wind-speed estimates: Further clarification of the Fujita-scale is furnished by the following descriptions: F-0 Slight damage to chimneys; branches broken; shallow-rooted trees knocked over. F-1 Surface of roofs peeled off; mobile homes pushed off foundations or overturned; moving autos pushed off roads. F-2 Roofs torn off frame houses; mobile homes demolished; boxcars pushed over; large trees snapped or uprooted. F-3 Roofs and some walls torn off well-constructed houses; trains overturned; most trees in forests uprooted; heavy cars lifted off ground and thrown. F-4 Well-constructed houses leveled; structures with weak foundations blown off; large missiles generated. F-5 Strong frame houses lifted off foundations and carried considerable distances; automobile-sized missiles flying through the air for distances in excess of 100 m; trees debarked. 2.5.4.2 Geographical and Seasonal Distribution of Tornadoes Although tornadoes occur in most parts of the globe, they seem to have a strong geographical and seasonal bias. In USA where their frequency of occurrence per unit area appears to be the highest in the world with more than 1000 per year, majority of the tornadoes form over the Great Plains of the midwest comprising the States of Nebraska, Kansas, Oklahama, and north Texas, and the southeastern States of Mississipi, Alabama, Louisiana and Florida. This continuous zone of the States is often referred to as the tornado alley. They can occur throughout the year at any time of the day but the maximum frequency is during spring and early summer. Multiple tornadoes have been observed in association with certain types of synoptic-scale disturbances. They usually move from west to east with prevailing upper winds. In West Africa where they occur mostly during summer, they move from east to west. In eastern India and Bangladesh, they occur mostly during spring and early summer months and are known as ‘Kal-Baisakhis’ or deadly storms of the month of Baisakh (April–May) and usually move from a NW-ly direction.
60
2 Tropical Disturbances
Observations appear to suggest that tornadoes tend to form in thunderclouds in which a large number of ‘mammatos’ appear at the cloudbase suggesting the presence of strong vertical currents inside them. The exact mechanism of how the vertical currents organize themselves to generate a tornado, however, is not known and open to speculations. One school of thought views a tornado as a meso-scale revolving storm with a horizontal and vertical structure very similar to that of a mature miniature tropical cyclone.
Chapter 3
Tropical Cyclones/Hurricanes/Typhoons – Their Structure and Properties
3.1 Introduction A tropical cyclone represents one of the most mature stages in the development of a tropical disturbance. However, it is well-known that out of the numerous low pressure disturbances which form every year in different parts of the tropics, only a very small percentage develops into a mature stage and that of those which get to that stage, only about a half develop an eye around the center. In this section, we present a brief account of the structure, development and movement of a tropical cyclone as we know them to-day from the results of extensive studies that have been carried out on the subject in the past. Readers interested in getting further information on the subject are referred to an excellent monograph on tropical cyclones by Anthes (1982) published by the American Meteorological Society. The National Weather Service in USA classifies hurricanes in five categories using the well-known Saffir-Simpson scale on the basis of central pressure and/or the strength of the tangential wind (Table 3.1).
3.2 Observed Structure of a Tropical Cyclone In satellite cloud imagery, a fully-matured tropical cyclone appears as a huge dense cloud mass of diameter about 200 km with a small opening of about 10–20 km diameter at the center and surrounded by a number of spiralling cloud bands which, starting from a great distance of say 1000–1500 km, appear to be converging into the cloud mass. The central opening is known as the ‘eye’ of the cyclone. These features appear in almost all mature tropical cyclones. An example is presented in Fig. 3.1 which shows a NOAA satellite view of hurricane ANDREW over the Gulf of Mexico on 23 August 1992. Hurricane ANDREW at one point was a category-5 hurricane and had a welldeveloped eye at the center. However, such pictures provide us with only an external view of the cyclone. For detailed meteorological information regarding the internal structure of the cyclone, we have to look to other sources which can obtain a closer K. Saha, Tropical Circulation Systems and Monsoons, C Springer-Verlag Berlin Heidelberg 2010 DOI 10.1007/978-3-642-03373-5_3,
61
62
3
Tropical Cyclones/Hurricanes/Typhoons
Table 3.1 Categories of hurricanes (USA) (after Simpson, 1974) Category of hurricanes
Central pressure (hPa)
Wind (m s–1 )
1 2 3 4 5
>980 965–979 945–964 920–944 <920
31–40 41–46 47–54 55–65 >65
Fig. 3.1 NOAA satellite view of hurricane ‘ANDREW’ over the Gulf of Mexico, 23 August 1992
view of the system. In fact, our current knowledge of the three-dimensional structure of meteorological conditions inside tropical cyclones has been derived mainly from radiosonde, radar and aircraft data. Since 1936, an expanding network of radiosonde observatories has been operating in different parts of the globe. They routinely record pressure, temperature and humidity at different levels of the atmosphere. Similarly, several countries exposed to the dangers of these storms have set up radar stations to observe and measure various storm parameters. But, perhaps, the most comprehensive and reliable sets of data regarding the three-dimensional structure of tropical cyclones have come from aircraft probes of these storms. The meteorological community owes a great deal of gratitude to the valiant pilots of the hurricane research aircraft of the United States of America
3.2
Observed Structure of a Tropical Cyclone
63
Weather Bureau (now National Weather Service) who, since the forties of the twentieth century, at great risk to their lives, penetrated deep into hundreds of these dangerous storms including the eye and collected data at various flight levels and at various distances from the center for each individual cyclone. Detailed reports of these flights are available with the National Hurricane Research Laboratory and the National Hurricane Center at Miami (Florida) and constitute, perhaps, our best source of data for determining the structure of cyclones. The valuable reports have been used in two ways. A few workers constructed the structure of individual cyclones. For example, Hawkins and Imbembo (1976) reported the wind and temperature structure of hurricane ‘INEZ’ that occurred in September 1966. Sheets (1980) reported the structure of wind and pressure distributions inside hurricane ‘ANITA’ in September 1977. Several workers (e.g., Hughes, 1952; Gray and Shea, 1973; Frank, 1977; Gray, 1979) used a compounding technique in which data from several cyclones were stratified according to the size, intensity, season and quadrant of each cyclone and many other parameters. The compounding technique reduced random errors, and, by averaging, eliminated important asymmetries between individual cyclones. The analysis of such averaged data provides, perhaps, a most complete three-dimensional composite picture of the large-scale structure of a tropical cyclone. A brief description of the radial and vertical distributions of wind, pressure, temperature and several other parameters in some individual as well as composite cyclones is provided in the pages that follow. The information furnished is representative of conditions in respect of hurricanes in North Atlantic Ocean and typhoons in Western North Pacific Ocean.
3.2.1 Wind Structure Hawkins and Imbembo (1976) present the horizontal and vertical distributions of several important wind parameters in respect of hurricane ‘INEZ’. Their streamline and isotach analyses of low-level winds at about 950 hPa in this hurricane are presented in Figs. 3.2 and 3.3 respectively. Figure 3.3 shows the wind speeds (isotachs) at different locations around the eye of the hurricane INEZ. The analyses of the wind field presented in Figs. 3.2 and 3.3 reveal, inter alia, the following: (1) Considerable degree of asymmetry in the direction and speed of the flow outside the eye-wall; (2) The spiralling winds starting from great distances appear to be converging towards the center and meet in concentric circles along the eye-wall; (3) At the center of the eye, there appears to be little or no wind.
64
3
Tropical Cyclones/Hurricanes/Typhoons
Fig. 3.2 Analyses of low-level winds (streamlines) at 950 hPa in hurricane ‘INEZ’ (Hawkins and Imbembo, 1976)
3.2.2 Radial and Tangential Components of the Wind 3.2.2.1 Radial Wind A vertical cross-section of the radial component of the wind (m s–1 ) in western Atlantic composite hurricane (Gray, 1979) is presented in Fig. 3.4. The cross-section shows a layer of inflow of environmental air at low levels and a layer of outflow aloft, the two flows being separated by a surface which slopes outward with height from about 800 mb at the boundary of the eye wall to about 300 mb at the outer boundary of the hurricane at a distance of about 1000 km from the center. The strongest inflow occurs in the layer between about 950 and 850 mb, while the strongest outflow occurs between about 300 and 100 hPa. A layer of inflow appears to occur above the tropopause level.
3.2
Observed Structure of a Tropical Cyclone
65
Fig. 3.3 Analyses of the wind field (isotachs) around the eye of the hurricane ‘INEZ’ (Hawkins and Imbembo, 1976) The axisymmetric eyewall is shaded
3.2.2.2 Tangential Wind Figure 3.5 shows the zonal-vertical distribution of tangential winds (m s–1 ) at different distances from the center of the eye in a Pacific composite typhoon, after Frank (1977). Figure 3.5 shows cyclonic flow in the lower troposphere and anticyclonic flow in the upper troposphere. Similar features of the distribution of tangential winds are also revealed by a vertical cross-section for the tangential wind speed (m s–1 ) in hurricane INEZ studied by Hawkins and Imbembo (1976). Their report shows that at the radial boundaries of the eye wall the tangential wind attains hurricane speeds (>100 Kts). It also shows that the maximum speed inside the wall is higher in the rear of the storm than in front by about 10 knots in a westward-moving system.
66
3
Tropical Cyclones/Hurricanes/Typhoons
Fig. 3.4 Vertical cross-section of radial winds (m s–1 ) in the western Atlantic composite hurricane (Gray, 1979) Negative sign means inward flow
The distribution shows that starting from a distance of about 40 miles where the speed is about 50 knots the speed continually increases till it reaches the highest value of about 120 knots in the eye wall and then it drops rapidly to very low values inside the eye of the storm. Within the eye-wall, a weak vertical wind shear is maintained between about 900 and 400 mb by intense cumulus convection which transports cyclonic momentum upward. Outside the eye wall, the speed of the tangential wind decreases with distance from the center. Attempts have been made to estimate the tangential and radial components of the wind in a hurricane by making use of the law of conservation of angular momentum, given by the relations vλ r + fr2 /2 = Const
(3.2.1)
∂vr /∂r + vr /r = 0
(3.2.2)
where vλ , vr are the tangential and the radial components of the wind respectively, r is the radial distance from the center and f is the Coriolis parameter (taken constant in space). Since (3.2.1) would lead to infinite speeds at the center, Deppermann (1947) proposed a so-called Rankine vortex with velocity distribution, vλ /r = Constant, for the central area. The radial distribution of the tangential wind speed, vλ (r), outside the radius of the maximum wind, R0 , is often represented by the empirical relation
3.2
Observed Structure of a Tropical Cyclone
67
Fig. 3.5 Vertical distributions of tangential winds (m s–1 ) at different distances from the center in a Pacific composite typhoon (Frank, 1977). Positive denotes cyclonic
vλ (r) = vλ (R0 )(R0 /r)x ,
R0 < r < r0
(3.2.3)
where r0 is at the outer edge of the disturbed area of the hurricane at a radial distance of about 1000 km or so. The exponent x in (3.2.3) varies from 0.5 to 0.7 near the center of different storms (Miller, 1967). Sheets (1980) computed values of vλ (r) in hurricane ‘ANITA’ using (3.2.3) for values of x = 0.5 and 0.6.
3.2.3 Vertical Motion in a Mean Typhoon Frank (1977) computed vertical motion (mb day–1 ) in a mean Pacific typhoon, which is shown in Fig. 3.6. The radial-vertical section in Fig. 3.6 shows penetrative convection with a maximum of 400 mb day–1 (though it may be much stronger in an individual typhoon) in the eye-wall involving the whole troposphere, which appears to extend to about 4◦ latitude circle from the center. Beyond that distance, while light upward motion continues in the upper troposphere (above about 400 mb) up to a distance of at least 1000 km and replaced by downward motion farther away, alternate zones of subsidence and convection are found in the lower troposphere up to a distance of about 1000 km from the center. This brings us to an important question: Why is
68
3
Tropical Cyclones/Hurricanes/Typhoons
Fig. 3.6 A cross-section showing vertical motion (mb day–1 ) in mean typhoon (Frank, 1977)
this difference in the structure of vertical motion between the lower and the upper tropospheres? Perhaps, the answer lies with the stability properties of the two layers and the intensity of moist convection at low levels. The upper troposphere is statically stable, so a weak low-level convection is unable to pierce through the stability barrier above and its effect remains confined to the lower troposphere only with alternate zones of upward and downward motion as in a gravity wave motion. On the other hand, strong convergence in the eye-wall leads to penetrative convection involving both the lower and the upper tropospheres.
3.2.4 Pressure Distribution Barograph traces at shore stations lying in the path of a hurricane often give a measure of pressure fall in the core of a hurricane if the rate of movement of the storm center is known. Riehl (1954) quoted the barograph trace marking the passage of a hurricane in New Orleans, Lousiana, USA in September 1947, which is shown in Fig. 3.7.
3.2
Observed Structure of a Tropical Cyclone
69
Fig. 3.7 Barograph record marking the passage of a hurricane in New Orleans, Lousiana, USA in September 1947 (after Riehl, 1954). Pressure is in inches of mercury
It shows that the pressure fell extremely rapidly in the three hours before arrival of the eye. It rose equally rapidly after passage of the eye.
3.2.5 Temperature Distribution Hurricane ‘INEZ’ was a rather small-scale hurricane, though it was well-developed. It was found to possess a temperature structure very similar to that in many largescale hurricanes, as shown in Fig. 3.8. In Fig. 3.8 which shows the radial distribution of temperature anomalies (deviation from mean annual tropical atmosphere) in hurricane ‘INEZ’, a most striking feature of the temperature distribution is the appearance of two concentrated cores of warm air inside the eye wall, one in the upper troposphere with the maximum warm anomaly (deviation from mean annual tropical atmosphere) above the environment of about 16◦ C at 250 mb and the other in the lower troposphere with the maximum warm anomaly above the environment of about 11◦ C at 600 mb. The warm anomaly is minimal at about 500 mb between the maxima above and below. A strong temperature gradient exists across the eye wall at all levels, especially in the lower troposphere. The warm anomaly in the eye wall where heavy precipitation occurs is about half of what is inside the eye. In the environment outside the eye wall, alternate layers of cold and warm anomalies appear in the vertical. For the same radial-vertical section inside the hurricane ‘INEZ’, Hawkins and Imbembo also worked out the vertical distribution of equivalent potential temperature, θ e (not shown here), which showed that, by and large, the same type of thermal structure and vertical stability prevailed in the field of the hurricane as exhibited by the three well-known tropical soundings, presented in Fig. 3.9 (taken from Anthes, 1982, with permission of Am. Meteor. Soc.)
70
3
Tropical Cyclones/Hurricanes/Typhoons
Fig. 3.8 Radial-vertical distribution of temperature anomalies (◦ C) (from mean annual tropical atmosphere) through the center in a NW–SE direction in hurricane ‘Inez’ on 28 September 1966 (Hawkins and Imbembo, 1976)
Since the formation of an eye is crucial to the generation of extremely low pressures and high winds in a hurricane, it is interesting to consider how these low pressures found in the eye are created. Observations show that at the top of a hurricane, the horizontal pressure gradient almost disappears. This means that a lowering of pressure at surface is directly proportional to an increase in the temperature of the air column inside the eyewall. Malkus and Riehl (1960) showed that the warmest possible column of air rising from surface moist adiabatically with a value
3.3
The Eye and the Eye-Wall
71
Fig. 3.9 Vertical profiles of equivalent potential temperature, θ e for mean September tropical sounding, mean hurricane within 100 n miles of the center, and mean eye of storms with pressures between 939 and 949 mb
of θ e = 350 K at surface would yield a pressure of 1000 mb at surface. They showed that under moist adiabatic conditions, a change in equivalent potential temperature would be related to a change in mean sea level pressure, ps , by the empirical relation (3.2.4) ps = −2.5θe The relation (3.2.4) assumes that the whole atmospheric column has the same. Equivalent potential temperature as at the surface; an assumption which is seldom true in the real atmosphere.
3.3 The Eye and the Eye-Wall 3.3.1 General Considerations – Formation of the Hurricane Eye While it is likely that a high value of equivalent potential temperature and deep moist convection may explain a certain amount of lowering of pressure in the eye wall, the same cannot be said for the formation of the lowest pressure at the center
72
3
Tropical Cyclones/Hurricanes/Typhoons
of the hurricane eye which is found to be relatively dry and cloudless and much warmer than the eye wall. According to Gray and Shea (1973), the temperature excess over the environment in the eye is almost double of that in the eye-wall, an observation which is well supported by hurricane ‘INEZ’ data (Fig. 3.8). It is due to the temperature difference that pressure is considerably lower in the eye than in the eye-wall or the environment outside the eye-wall. Thus, the eye plays a fundamental role in the thermodynamics of the hurricane circulation by producing the warmest core with the lowest pressure at the center and the strongest hurricane winds and heaviest rainfall in the eye-wall. What then is responsible for the triggering and formation of a hurricane eye? Numerical studies (e.g., Eliassen, 1952, 1959; Kuo, 1959; Krishnamurti, 1961; Estoque, 1962; Willoughby, 1979; Smith, 1980) by computing the radial – vertical circulations in a symmetric hurricane using plausible assumptions in regard to friction, condensation heating, etc., have hinted at downward motion in the eye to explain the observed lowering of surface pressure ins ide the eye-wall. However, these studies leave many questions unanswered, especially that relating to the role of heat released by condensation in the evolution of the circulation, especially how it leads to such low pressure inside the eye. It is evident that a unified, comprehensive theory of development of a hurricane eye starting from a pre-existing feeble wave disturbance is yet to come. Our conceptual model of development of a cyclone from a pre-existing monsoon trough, presented in Sects. 2.4.2 and 2.4.3, emphasizes the role of condensation heating in producing the subsidence warming at the center by an indirect circulation which is required for supporting the energy-producing direct circulation between the eye-wall and its outer environment. The degree of development depends on the extent of the subsidence warming in the eye produced by condensation in the eyewall. The presence of such warm areas, one in the upper and the other in the lower troposphere in the eye region was well demonstrated by the temperature profiles in hurricane INEZ (Fig. 3.8). Initially, a trough of low pressure has only one sector of large-scale condensation at the location of the ITCZ and a subdued one at the TCZ. The heat generated at these convergence zones through condensation is, perhaps, just about what is required to maintain the trough in its existing form; it cannot lead to further development. It requires docking with an easterly or W’ly wave to initiate the development process by producing two active zones of penetrative convection, one on either side of the trough axis and a zone of subsidence warming and lowering of pressure in between. Thus, development proceeds by stages as the zone of condensation heating extends around the center of the low pressure trough by the tangential winds and it is only when it encircles the center and the disturbance becomes quasi-axi-symmetric that it attains maximum development and gets ready to form an eye. However, observations show that once formed, an ‘eye’ may not last long. Its continuance is subject to a delicate balance amongst the processes which create it. In some cases, with little variation in dimension and intensity, it may last for hours and days, or form a new after a collapse, especially when the system is over a warm ocean with favorable environmental conditions. In others, it may appear only momentarily and then vanish from the scene.
3.3
The Eye and the Eye-Wall
73
3.3.2 Circulation Inside the Hurricane Eye – Evidence of Meso-Scale Vortices Recent studies by Montgomery et al. (2006) and Aberson et al. (2006) and several others based on an observational analysis of a unique set of data inside hurricane ‘Isabel’ over the western part of North Atlantic, 12–14 September 2003, appear to throw light on several aspects of the low-level circulation inside its eye and its thermodynamics. The storm-relative data from multiple dropwindsondes, several in-situ visual and radar observations from aircraft flying at different levels below about 4 km, as well as high-resolution satellite imagery at different radial distances from the center and different heights inside and outside the eyewall during the period (12–14 September) were composited and the analysis was based on these composites. Figure 3.10 (the upper panel) is a satellite view of the hurricane eye at 1711 UTC on 13 September 2003, which reveals the presence of four meso-scale vortices near the inner edge of the hurricane and a pentagonal structure of the eye-wall. Isabel was a category-5 hurricane on the Saffir-Simpson scale during the period, 12–14 September 2003, with an azimuthally-averaged tangential wind speed maximum of 76 m s–1 located at an altitude of about 1 km above the ocean surface at a radial distance of about 42 km from the center. However, an extreme tangential wind speed maximum of 107 m s–1 and a vertical wind maximum of 25 m s–1 were observed by a dropwindsonde at an altitude of about 1.4 km a.s.l. at 40 km radial distance at 1752 UTC on 13 September 2003. It was also found that the tangential wind speed fluctuated greatly azimuthally at this radial distance. It is likely that these extreme wind speeds and their azimuthal fluctuations in the horizontal are related to the mesovortices that were observed at these locations. According to Montgomery et al. (loc. cit.), though the high radial and vertical resolutions of nearly 30,700 data points (approximately half of which were from dropwindsondes) instill high confidence in the retrieved axisymmetric features in the eyewall region, greater uncertainty exists about the axisymmetric features within the 30 km-radius of the center where data sampling was limited. The lower panel (b) of Fig. 3.10 shows a schematic representation of the likely low-level horizontal circulation inside the eye of hurricane Isabel with four identified mesovortices near the inner edge of its eyewall. Panel (c) shows a sketch of the likely radial-vertical secondary circulation inside the hurricane eye. Montgomery et al. (loc. cit.) used the data of the storm-relative tangential and radial winds as well as the equivalent potential temperature (θ e ), absolute angular momentum and transverse secondary circulation (vector) composited from the GPS dropwindsonde and the flight-level data between 1600 and 2300 UTC on 13 September 2003 to derive the radius-height mean structure of the circulation in the eye and the eyewall region. Some of their important findings are: (a) The tangential wind speed has a maximum at a radial distance of 40–50 km from the eye center at altitude 2–3 km a.s.l.; (b) The base of the eyewall appears lifted to a height of about 1 km, above the ocean surface, thereby allowing outside oceanic air to flow heavily into the eye
74
3
Tropical Cyclones/Hurricanes/Typhoons
Fig. 3.10 (a) A visible Moderate Resolution Imaging Spectroradiometer (MODIS) image (from NASA’s Aqua Satellite) of the eye of hurricane Isabel at 1711 UTC on 13 September 2003, showing the four whirls and vortices near the eye-wall and the pentangular structure of the eye-wall (after Montgomery et al., 2006 and Aberson et al., 2006). The lower panel (b) gives a schematic representation of the likely low-level horizontal circulation inside the eye with mesovortices near the eye-wall. The lower panel (c) is a schematic of the radial-vertical secondary circulation inside the eye. Shaded areas are zones of windspeed maxima. L denotes Low, V vortex
region to a radius of about 15 km. A low-level (0–250 m) radial inflow speed maximum of 20 m s–1 is located at a radial distance of 25 km from the center; (c) The equivalent potential temperature has a value of about 370 K at the eye center (0–15 km radius), 360 K at the nominal eyewall (40–50 km radius), 355 K in the outer core (about 200 km radius) and about 350 K in the ambient environment (300–1000 km radius) at the sea surface. Within the first 2 km above the ocean
3.3
(d)
(e)
(f)
(g)
The Eye and the Eye-Wall
75
surface, it varies little with height within the eye and the eyewall, but outside the eyewall it decreases steadily with height; The low-level inflowing air converges into the circulation which diverges out from the central region of the eye and the convergent currents flow into the base of the eyewall and rise in penetrative convection with heavy condensation inside the eyewall. However, condensation heating leads to strong divergence above about 1.5 km level, one branch turning inward towards the center of the eye at a height of about 2 km and above and the other turning upward and outward; The inward-turning branch of the diverging current above 2 km would appear to converge at the center of the eye and subside to lower levels only to diverge out near sea level towards the eyewall in some kind of a vertical secondary circulation; The secondary vertical circulation between the center of the eye and the eyewall adjusts itself suitably to accommodate a mesovortex if one happens to be present; As pointed out earlier, potential energy generated by subsidence of the warmer air at the center and lifting of the colder air in the eyewall in the secondary circulation would appear to be what provides for the kinetic energy of the hurricane circulation in which the warmer air rises in the eyewall and the colder air subsides in the outer environment.
3.3.3 Concentric Multiple Eye-Walls A tropical cyclone is a continually evolving system in which more than one eye-wall can form at different radial distances from the center, especially in intense cyclones. McNoldy (2004) reports a case of three concentric eye-walls which were observed in a hurricane named ‘Juliette’ in the eastern North Pacific during the period 23–27 September 2001. From details of observations by satellite as well as reconnaissance aircraft, McNoldy presents the radial distributions of tangential velocity and relative vorticity in the northwest quadrant of the hurricane at a height of about 3 km on 25 and 26 September, shown in Fig. 3.11. The hurricane developed off the coast of southern Mexico and moved almost parallel to the coastline to enter Baja California later in the month. The northwest quadrant was chosen for its higher resolution and better quality data on both days. The relevant features are the two peaks in tangential velocity at 9 and 58 km (with corresponding peaks in relative vorticity at 7 and 55 km) on 25 September and the three peaks in tangential velocity at 11, 56, and 90 km (with corresponding peaks in relative vorticity at 9, 54, and 82 km) on 26 September. In this simplified radial perspective provided by Fig. 3.11, the relative vorticity is produced solely by the radial gradient of the tangential wind. The steepness of the vorticity curves may play a role in the formation of moats. From the diagram, it is clear that the sharpest drop-off of relative vorticity occurs on the outer side of the innermost wall, that on outer side of the second and the third eyewalls being much less prominent.
76
3
Tropical Cyclones/Hurricanes/Typhoons
Fig. 3.11 Radial profiles of (top) tangential velocity and (bottom) relative vorticity through the northwest quadrant of Juliette on 25 and 26 September 2001 at 1819–1849 and 1722–1755 UTC, respectively (after McNoldy, 2004)
Once the initial eye-wall is formed, the positive relative vorticity falls steeply beyond the radius of maximum tangential velocity, establishing some kind of a moat beyond which deep convection is free to organize a new concentric eye-wall. Sometimes in just a few hours, the new outer eye-wall will dominate and the inner eye-wall will dissipate. The outer eye-wall will then move in to replace the inner eye-wall. The process may continue at various radial distances to produce more than one concentric eye-wall. Of course, it is the innermost eye-wall which will dominate, the intensity of the others falling off sharply with radial distance. As McNodly states: ‘Concentric eyewalls are ephemeral; once formed, they typically are not maintained for much longer than 12 h As the new outer wall forms, the original inner eyewall usually lacks the necessary inflow to maintain itself and
3.4
Spiral Bands Around the Eye-Wall
77
it gradually dissipates. In time, the dynamic processes cause the outer eyewall to contract, and the process can repeat itself.’ Black and Willoughby (1992) had called the process an eyewall replacement cycle. Observations of multiple concentric eyewalls in hurricanes are rather scarce, but from a study of western North Pacific tropical cyclones during 1969–1971, Willoughby et al. (1982) estimated that approximately 53% of intense tropical cyclones (windspeeds greater than 65 m s–1 ) exhibit concentric eyewalls. More often than not, individual convective cloud cells at the new eyewalls may move in bands spiralling inward to converge into the innermost eye-wall.
3.4 Spiral Bands Around the Eye-Wall 3.4.1 Structure Almost all mature tropical cyclones exhibit a number of (usually two or three) welldeveloped cloud bands spiralling inward into the central cloud mass (see Fig. 3.1). What are these cloud bands and how do they form? While there is considerable uncertainty about their origin, Senn et al. (1957) has shown that the geometrical structure of many of these bands follows a modified logarithmic structure Ln(r − r0 ) = A + Bλ
(3.4.1)
where r is radius, λ azimuth, r0 the radius of a limiting circle with center coincident with the cyclone center, and A is a constant that specifies the angular origin (orientation) of the spiral, The constant B determines the crossing angle α between the spiral path and the circle about the center according to the relation tan α = B(1 − r0 /r)
(3.4.2)
The constant B is evaluated by noting α at an infinite distance from the center, so that (r0 /r) tends to 0, and B = tan α. The spiral bands consist of large convective cloud cells associated with smallscale cumulus convection. The individual cells move with the mean wind in the layer in which they are embedded within the spiral bands which move at a much slower rate. The cloud cells which move through the bands have typical life spans of 20–40 min and last as long as they experience rising motion within the band in convectively unstable air. New cells form at the upwind (inner) side of the band, travel through the band, and dissipate at the downwind (outer) side of the band. The cells yield heavy precipitation while they move through the band. Some of the individual cloud cells within the band appear to move inward towards the center of the cyclone, while others move outward. However, the number of those that move inward appears to predominate. A schematic model based on several observational studies of the structure of spiral bands indicates that a meso-scale trough of low pressure of amplitude about
78
3
Tropical Cyclones/Hurricanes/Typhoons
1–2 mb occurs at the leading edge of the spiral band with heaviest rainfall occurring about a quarter wavelength behind the trough, the width of the spiral band being half a wavelength. This structure is consistent with the phase relationship of a shallow water gravity wave. The mean low-level winds show convergence into the trough of low pressure in front of the band and divergence in the rear. At the approach of a band, the mean wind veers by 10–20◦ and then backs towards the original direction after the passage of the band. Because of strong convection in the band, the surface wind often becomes highly gusty during its passage. Gentry (1964) analyzed many aircraft-observed profiles of wind, temperature, humidity and pressure (D-values) along and perpendicular to spiral bands and found that the vertical velocity and temperature perturbations in the bands are positively correlated, with warm updrafts and cold downdrafts, thereby producing kinetic energy within the bands which in magnitude was almost comparable to the kinetic energy produced by the large-scale circulation associated with the cyclone. It is likely, however, that much of the perturbation kinetic energy produced locally in the bands is offset by local dissipation due to a high degree of turbulence in the vicinity of intense convection.
3.4.2 Origin and Direction of Propagation Observational evidence regarding the origin and direction of propagation of spiral bands varies, although the predominant evidence supports an outward propagation of the band relative to a stationary storm center (propagation velocity of the band after the velocity of the storm has been subtracted). According to Senn and Hiser (1959) and several others, the bands originate near the eye and propagate radially outward. Fletcher (1945) hypothesized that bands of clouds from the intertropical convergence zone were drawn into the cyclone’s circulation and became the spiral bands. Wexler (1947) suggested that cloud streets which are common in the tradewind atmosphere coil into the cyclone. Tepper (1958) proposed that gravity waves were generated near the eye-wall and propagated outward to become spiral bands. A number of theoretical and simulation studies (e.g., Abdullah, 1966; Kurihara, 1976; Diercks and Anthes, 1976; Willoughby, 1977, 1978a,b) have also been carried out to explain the generation, structure and propagation of hurricane spiral bands. A detailed review of some of these studies has been presented by Anthes (1982).
3.5 Storm Surge 3.5.1 Introduction Cyclones disturb the ocean as well as the atmosphere. Out over the open ocean, the extremely low pressure under a cyclone raises the water level producing a mound, the height of which varies with the pressure anomaly. The high winds drag the ocean
3.5
Storm Surge
79
surface producing stresses and ocean currents and also mighty waves which are harmful to ocean-going vessels which might happen to be in its vicinity. But, by far, the greatest damage is done when it approaches an inhabited coastal belt and produces what is known as a storm surge at the coast. The water level may then build up to several meters which can inundate a whole coastal region inflicting death and destruction to life and property. It is estimated that the storm surge produced by the Bangladesh cyclone in November 1970 in the delta region of Ganga-Brahmaputra killed as many as 300,000 people apart from flooding extensive low-lying areas. The storm surge problem is, however, complicated, because it depends not only upon the characteristics of the cyclone but also upon the vertical profile of the ocean floor near the coast, such as the depth and slope of the sea-bottom, orientation of the coastline to the storm track, presence of a bay, peninsula, island or estuary, etc, in the path of the storm. Further, the possible effect of an astronomical tide on the rise of sea level also needs to be taken into consideration. The superposition of an astronomical high tide on the effects of the cyclone can magnify storm surge to dangerous levels. A recent example of the extent of damage and destruction that a storm surge can inflict upon a coastal community is that associated with the category-5 hurricane, Katrina, which on August 29, 2005, while crossing the coast of Louisiana from the Gulf of Mexico near Buras-Triumph broke three imposing concrete levees which protected the city of New Orleans and flooded the city killing at least 1420 people and causing a material loss to the tune of $75 billion dollars. It was one of the costliest hurricanes so far in American history.
3.5.2 Some General Aspects of Storm Surge The total rise of water level in a storm surge is produced by the superposition of several scales of wave motion in the ocean. The large-scale motion of period of several hours may be of the scale of the cyclone itself and measured by the radius of the maximum wind or any such length. The pressure drop associated with the cyclone and the wind stress averaged over horizontal scales of several kilometers and period of an hour or so are the forcing functions for the large-scale oceanic waves. Typical surge amplitudes associated with this scale may range from 2 to 6 m. Then, there are smaller scale fluctuations of water level produced by individual waves the amplitudes of which may vary from 1 to 10 m and the period of which may be a few seconds only. But, by far, the greatest rise of water level and for a somewhat longer duration is caused when the effects of local ocean-bottom and coastal topography as well as astronomical tides are superimposed on the responses due to large-scale and small-scale ocean waves. Because of the sensitivity of storm surge to all these varying factors, including the occasional tsunamis, it is difficult to present a standard model of the phenomenon. The simple picture described above is, however, strongly modified by coastal topography and configurations. A bay, for example, may double the amplitude of the storm surge. In some bay region, the ocean bottom slopes up to shallow waters a long distance before the coastline and a cyclone heading towards such a coastline
80
3
Tropical Cyclones/Hurricanes/Typhoons
produces disastrous storm surge effects. Such bay regions can be found over several parts of the globe, for example in the Ganga-Brahmaputra delta region of IndiaBangladesh at the northern boundary of the Bay of Bengal, or the Mississippi delta region of USA off the coast of Louisiana at the northern boundary of the Gulf of Mexico, where extensive storm surge flooding occurs whenever a severe cyclone crosses these coasts.
3.5.3 Mathematical Models of Storm Surge Since storm surge involves nonlinear interactions between the atmosphere, the ocean and bottom topography, it is impossible to devise a mathematical model which can be solved analytically. Only highly simplified numerical models may be tried to determine the approximate oceanic response under actual cyclone conditions. In a pioneering study, Jelesnianski (1965, 1966, 1967) developed a general numerical model of the storm surge and has used it to test various physical hypotheses regarding the effects of (a) pressure anomaly; (b) wind stress, and (c) intensity, size, speed and direction of motion of the cyclone in producing the storm surge. Jelesnianski’s storm surge models consider only the barotropic mode. The linearized equations of his model (Jelesnianski, 1966) consist of predictive equations for the depth-weighted velocity components U and V, and the continuity equation for h, the departure in height of the surface water from the undisturbed level. The equations are: ∂U/∂t = −gD(x,y)∂h/∂x + fV + {D(x,y)/ρ}∂Pa /∂x + ρ −1 τx (x,y,t)
(3.5.1)
∂V/∂∂t = −gD(x,y)∂h/∂y − fU + {D(x,y)/ρ}∂Pa /∂y + ρ −1 τy (x,y,t)
(3.5.2)
∂h/∂t = −(∂U/∂x + ∂V/∂y)
(3.5.3)
where D(x, y) is the initial (undisturbed) depth of the fluid, τ is the surface wind stress per unit mass, and Pa is the atmospheric pressure at the surface. Equations (3.5.1), (3.5.2), and (3.5.3) describe the linearized response of a shallow layer of the fluid to the forcings of atmospheric pressure gradients and surface wind stress, τ (x, y, t), on a rotating earth. Besides forced modes, the system allows inertia-gravity waves. A typical ocean basin is resolved as shown in Fig. 3.12. Finite-difference technique is adopted to solve the Eqs. (3.5.1), (3.5.2), and (3.5.3). The model is forced by specifying Pa (x, y, t) and τ (x, y, t) from the field of either an actual cyclone or a model cyclone. In Jelesnianski’s simulations, the pressure and stress fields are specified by empirically determined functions of space. Storm parameters included in the functions are the intensity (maximum wind), size (radius of the maximum wind), and an inflow (cross-isobar) angle. From the wind distributions, the surface eddy stress τs was estimated using the bulk aerodynamic relation
3.5
Storm Surge
81
Fig. 3.12 A typical one-dimensional ocean depth profile near a coast. A finite-difference grid with variable horizontal resolution S is shown. The curved bottom topography is approximated by a series of steps with constant depth Di (i = 1, 2, 3,. . .). Higher resolution may be introduced near the coast (Jelesnianski, 1966)
τ s = cD ρa |Va |Va where cD is the drag co-efficient and the subscript ‘a’ refers to the air. In some preliminary experiments, Jelesnianski (1965) investigated the ocean’s response to a variety of prescribed storm conditions, such as (a) a circular wind stress only, (b) a wind stress with a radial (inward) component introduced by a 30◦ cross-isobar angle, (c) an asymmetric wind stress associated with a moving storm, and (d) an atmospheric pressure gradient only. Two interesting results emerged from these preliminary experiments. The first was a 60% increase in the maximum surge at the coast by inward flow of water when a radial component to the stress was introduced. The second was a three times increase in the static height ‘h’ of the mound of water under the storm at the coast when the barometric pressure gradient alone was used to force the ocean (the static height is the height of the mound of water under the cyclone above the undisturbed level when it is far out over the ocean and is given by the relation, h = Pa /ρg, where ρ is the density of water). It turned out from the preliminary experiments that the two major forces driving the storm surge are the wind stress and the atmospheric pressure gradient. To test the relative importance of the two forcings, the model war run with each force separately and then with the combined forces. The results showed that the effect of the wind stress was approximately double of that due to the pressure gradient. This would be clear from Fig. 3.13 which shows the maximum coastal surge for the
82
3
Tropical Cyclones/Hurricanes/Typhoons
Fig. 3.13 Maximum coastal surge for forces associated with wind stress and pressure anomaly alone and combined forces as a function of speed of storm crossing coast at an angle of 90◦ . The dots are values from actual computer runs (Jelesnianski, 1966)
three combinations of forces when the storm crosses the coast at an angle of 90◦ with different speeds. Note that the peak storm surge obtained from the combined forces is not a simple sum of the surges obtained from the forces separately. Also, the maximum storm surge does not vary linearly with the storm speed, but reaches a maximum at an approximate speed of 30–40 knots. Jelesnianski (1966) simulated storm surge effects on the same coastal topography for many different model storms and constructed nomograms showing the quantitative effects of storm speed, intensity and direction of approach on the characteristics of the storm surge. The reader interested in further details of these effects and the nomograms may consult his original papers.
3.6 Prediction of Cyclone Track and Intensity From very early times, tropical cyclones have drawn the attention of meteorologists because of their power to destroy life and property wherever they occur. A demand for accurate forecast of these storms was, therefore, inevitable. Prior to the days of regular meteorological observations, the only guidance available in this regard was observations of movement of high clouds and sea swells in coastal areas. However, the deductions made from such observations were often misleading. It was found that plumes of high clouds diverging from the upper part of a cyclone did not necessarily move in the direction of motion of the storm to provide early warning of
3.6
Prediction of Cyclone Track and Intensity
83
its arrival. Further, sea swells were not exclusive to tropical cyclones. They could be caused by other agencies as well, such as underwater volcanoes and earthquakes (tsunamis). For example, the recent tsunami (high tidal wave) in the Indian Ocean area which killed nearly 175,000 people in the littoral countries originated from an underwater earthquake near the coast of Sumatra in Indonesia on 26 December 2004, not from a tropical cyclone.
3.6.1 Early Models – The Steering Concept With the availability of upper-air observations in the nineteen fifties and sixties, one of the methods used to predict cyclone motion with considerable success utilized what is known as the steering principle. According to this principle, small-scale cyclones tend to move with the large-scale currents in which they are embedded. Initially, a single level wind was used as the steering current but later it was realized that a layer-mean wind might be a better steering current than a single-level wind. The realization gave rise to other models. However, the basic concept of the steering current survived and found a place amongst the latest models of storm track forecasting. Since 1950s and 1960s, a number of objective methods have been developed to predict cyclone motion. These are either statistical or dynamical methods. Statistical methods seek to relate predicted movement to a number of cyclone parameters in an empirical way. The dynamical methods, on the other hand, use some form of the equations of motion to predict numerically the motion of the cyclone from an observed initial state of the atmosphere. In some objective methods, the output from a dynamical model is used in a statistical model. These are called hybrid models. Neumann and Pelissier (1981a,b) described a total of seven operational models which were in use at the National Hurricane Center (NHC) in USA for prediction of tropical cyclone motion in the North Atlantic Ocean area (Table 3.2).
3.6.2 Current Models During the eighties and nineties, there were continued efforts to improve the hurricane track forecast models in the light of experience gained during the earlier years. In 1988, SANBAR was replaced by the Quasi-Lagrangian Model (QLM) which had higher horizontal and vertical resolution and a larger model domain (Mathur, 1991). The QLM produced skillful forecasts during 1989–1993. Meanwhile, in 1992, the Geophysical Fluid Dynamics Laboratory (GFDL) developed a model which was adapted for real-time hurricane track forecasting. It included moving nested grids and a sophisticated initialization scheme. The horizontal grid spacing on the inner mesh was about 20 km, which was about half of that in QLM. The GFDL model outperformed all other models in terms of average track error when it was tried during the 1992–1993 hurricane seasons and replaced the QLM in 1995.
84
3
Tropical Cyclones/Hurricanes/Typhoons
Table 3.2 Operational models for the prediction of tropical cyclone motion over the North Atlantic area (after Neumann and Pelissier, 1981a,b) Model
Model type
Description (reference)
HURRAN
Statistical
CLIPER
Statistical
NHC-67
Statistical
NHC-72
Statistical
NHC-73
Statistical
SANBAR
Dynamical
MFM
Dynamical
Analog model based on tracks of all Atlantic tropical cyclones since 1886 (Hope and Neumann, 1970) Regression equation model, utilizing predictors derived from climatology and persistence (Neumann, 1972) Regression equation model utilizing predictors derived from climatology, persistence and observed geopotential height data (Miller and Chase, 1966) Regression equation model utilizing predictors derived from output of CLIPER model and observed geopotential height data (Neumann et al., 1972) Regression equation model utilizing predictors derived from output of CLIPER model, observed and numerically forecast geopotential height data (Neumann and Lawrence, 1975) Barotropic model based on pressure-weighted wind field averaged through troposphere and represented on a 154 km (at 22.5◦ N) space-grid (Sanders and Burpee, 1968, Sanders et al., 1975) Movable Fine Mesh (MFM) baroclinic model having 10 levels in the vertical and 60 km grid spacing in the horizontal (Hovermale and Livezey, 1977)
Since 1989, after withdrawal of SANBAR, a couple of other barotropic models have been tried. One of these was the VICBAR (Vic Ooyama Barotropic model) which has been run in NHC in an experimental mode in real time. The main advantages of VICBAR relative to SANBAR was that it obtained lateral boundary conditions from the National Meteorological Center (now National Centers for Environmental Prediction, NCEP) global forecast model and used a much more accurate numerical method. Another model which used the barotropic steering concept is the Beta and Advection Model (BAM). In this model, the track forecast is obtained by following a trajectory in the vertically averaged horizontal wind from the NCEP aviation global forecast model, after applying a correction for vortex drift. The BAM model which uses a layer-mean wind for steering has been used for shallow, medium and deep layers. All three versions are in use at NHC.
3.6
Prediction of Cyclone Track and Intensity
85
The large-scale regional and global models from NCEP have long been used as an aid in hurricane forecasting by providing forecasts of the hurricane environment. The skill of the global forecast models have improved over the hyears (Kalnay et al., 1994). This improvement can be attributed, inter alia, to improved model physics, increased horizontal and vertical resolution, and better data assimilation techniques. Beginning in 1992, a tropical cyclone bogussing system was implemented in the NMC operational global model for track forecasting (Lord, 1991). Currently, for cyclone track forecasting in the Atlantic, a number of models are in operation in NHC. These include: The simple analog model (HURRAN). CLIPER, the statistical-dynamical models (NHC-90), the barotropic model (VICBAR), BAM (3 versions), the GFDL regional baroclinic model, and the aviation model. All of these models, except CLIPER, use information from the aviation model. A version of the NHC-90 model, called UK-90, uses information from the UK Meteorological Office global forecast model. There were also efforts during this period to improve hurricane intensity forecasting. Experience showed that the intensity of a cyclone was very much influenced by environmental conditions, such as sea surface temperatures, upper-level flow patterns, temperatures near the tropopause, etc. These ideas were incorporated in a Statistical Hurricane Intensity Prediction Scheme (SHIPS) for the Atlantic basin which has been run in NHC on an experimental basis since 1990. Intensity prediction has also been attempted in the GFDL model, beginning in 1992 (Kurihara et al., 1993). Thus, a beginning has been made to use three-dimensional global circulation models to predict both track and intensity of hurricanes.
Part II
Tropical Monsoons over Continents and Oceans
Chapter 4
Monsoon over Southern Asia (Comprising Pakistan, India, Bangladesh, Myanmar and Countries of Southeastern Asia) and Adjoining Indian Ocean (Region – I)
4.1 Introduction – Physical Features and Climate The world’s largest and most powerful monsoon circulation develops over the region of Southern Asia and its adjoining Indian Ocean. There are several reasons for this development: The first is a most favorable land-sea distribution. The geographical location of the region (see a physical map in Fig. 4.1) with the Tropic of Cancer passing through almost the middle of the landmass of the region and the Tropic of Capricorn through the middle of the Southern Indian Ocean (indicated by thin dashed lines) appears to provide an ideal setting for a heat source to develop over the land and a heat sink over the ocean during northern summer and vice versa during northern winter. A second reason is orography. The lofty Himalaya Mountains and the Tibetan Plateau standing along the northern boundary of the region and rising steeply from the plains of northern India to peak heights of 8–9 km above sea level and then an average Plateau height of 4–6 km above mean sea level not only protect it from the icy cold winds of Central Asia during the winter but also help shape the structure of a well-defined summer monsoon circulation over the region and provide a natural barrier to the moisture-laden onshore winds from the Indian Ocean blowing into the region. The high mountain ranges all along the northwestern, northern and the eastern boundaries of the region as well as the coastal mountain ranges, such as the Western Ghats Mountain of peninsular India and the Arakan Yoma and the Tennaserim Ranges of Myanmar (erstwhile Burma) play important roles in shaping the structure of the monsoon circulation and the distribution of monsoon rains over the region. The geographical location of Southeastern Asia, sandwiched between the Bay of Bengal and the South China Sea, presents a complex array of narrow coastal mountain ranges and broad inland high plateaux with heights ranging from 1.5 to 2.5 km above sea level. Some of these ranges which are mostly north-south oriented rise steeply above the plains of Myanmar, Thailand and Malaysia. It will be shown later in this chapter that the geographical locations of these mountain ranges and high plateaux, both along the northern and the eastern borders of the Indian Subcontinent play crucial roles not only during advance and retreat of monsoon, but also in controlling the distribution of monsoon rains over the subcontinent. K. Saha, Tropical Circulation Systems and Monsoons, C Springer-Verlag Berlin Heidelberg 2010 DOI 10.1007/978-3-642-03373-5_4,
89
90
4 Monsoon over Southern Asia and Adjoining Indian Ocean
Fig. 4.1 Physical map of Southern Asia and adjoining Indian Ocean and its littoral countries
In the western sector, a chain of semi-arid and desert lands extends from Pakistan and Iran southwestward to Saudi Arabia and beyond to East Africa across Somalia, Kenya, Tanzania, Zimbabwe to Mozambique and South Africa against the backdrop of the high East African Mountains with heights ranging between 2 and 3 km asl. It will be shown in the present text that these varied land features along the western boundary of the Indian Ocean play crucial roles in advance and withdrawal of monsoon as well as in shaping the structure of the circulation and distribution of monsoon rains over the Western Indian Ocean and its littoral countries during both winter and summer. The seasons over the region which exercise great influence on climatic conditions are as follows: The winter season characterized by extremely cold and dry conditions over the northern parts of the region appears to cover a period of 3 months from December to February. Then comes a period of approximately 3 months, March to May, when the land surface warms up rapidly to very high temperatures and ‘heat lows’ develop over the land. This is the period of transition from winter to summer monsoon. Summer monsoon proper arrives towards the end of May or early June and covers the whole region including the northernmost parts of the Western Himalayas by end of August. It starts withdrawing from the northwestern part of the region from early September, but the process of withdrawal is not completed till end of November.
4.2
The Winter Season (December–February)
91
Thus, the total period of summer monsoon over the region, from its first onset to final withdrawal, is almost 6 months. It, therefore, stands to reason that in the present text, the period of summer (wet) monsoon be divided into two parts: the onset phase (June to August) and the withdrawal phase (September–November), in order to conform to ground realities. The transition from summer to winter monsoon conditions varies from place to place but is usually about a month. The winter monsoon is of shorter duration. It is important to bear in mind the great diversity of the seasonal weather and climate over the Indian Subcontinent arising largely from its topography and latitudinal extent from about 5 to 35◦ N. For example, in winter (December–February), while the northern parts of the subcontinent (north of about 25◦ N) shiver in nearfreezing temperatures at times, the Indian peninsula and a large part of SE Asia enjoys cool, salubrious climate. Conversely, during the SW monsoon season, cool, humid weather of the south and the southeast stands in sharp contrast to the hot, semi-arid desert climate of the NW India and Pakistan.
4.2 The Winter Season (December–February) 4.2.1 General Climatic Conditions In winter, the surface temperatures are generally low over the subcontinent and decrease with latitude. The lofty Himalayan mountain ranges along the northern boundary separate the region from Central Asia and protect it from the icy cold temperatures of that region. Relatively warmer temperatures prevail over the Indian Ocean, with the warmest temperatures being over the equatorial region where the ITCZ is located along about 15◦ S over the western Indian Ocean, but is much closer to the equator over the eastern part. In keeping with the above-mentioned temperature distribution, m.s.l. pressure is generally high over the cold Indian Subcontinent and low over the warm equatorial ocean. An interesting aspect of the pressure distribution is that the three land segments of Southern Asia, viz., the Arabian Peninsula, the Indian Subcontinent, and Southeastern Asia, develop their independent high pressure cells, with low pressure troughs located in between over the somewhat warmer oceans; one over the northern Arabian Sea and the other over the northern Bay of Bengal. The orographic influence of the Himalaya mountain ranges and other coastal mountain systems, such as the Western Ghats of India and the Arakan Yoma of Myanmar is also significant. Generally, across any of these mountain ranges of appreciable height, a trough of low pressure tends to form on the windward as well as the lee sides of the mountain range with a ridge of high pressure on the mountain itself. Accordingly, troughs of low pressure are observed to form along the foothills of the Himalayas; one over Pakistan in Western Himalayas and the other over West Bengal and Bangladesh in the eastern part of the subcontinent.
92
4 Monsoon over Southern Asia and Adjoining Indian Ocean
The prevailing quasi-stationary pressure distribution drives anticyclonic circulation around the high pressure cells and cyclonic circulations around the troughs of low pressure. As part of the anticyclonic circulations, low-level airflow is generally northeasterly over the northern Indian Ocean. However, near the equator, the airflow is cross-equatorial in the western part of the ocean (west of about 70◦ E), and more or less westerly in the eastern part. The equatorial westerlies in the eastern part of the ocean divide the warm low pressure over the area into two cells, one north and the other south of the equator, each having its own trough, viz., the North Equatorial Trough (NET) and the South Equatorial Trough (SET). According to rainfall map prepared by the India Meteorological Department (1943) (not shown), there is little rainfall over the subcontinent during the season, except areas affected by traveling disturbances.
4.2.2 Disturbances of the Winter Season 4.2.2.1 Western Disturbances (W.D.) These are large-scale wave disturbances which form in midlatitude-subtropical baroclinic westerlies and usually travel from west to east. During northern winter, they travel from the Eastern-Mediterranean region and enter the Indian region from November onward. Their arrival is heralded by appearance of first high clouds and then medium and low clouds with rain or snowfall in the mountains of western Himalayas. With advance of the season, their track continually shifts equatorward and more and more of the subcontinent are affected by them and experience rain or snowfall. In December, western disturbances cross the mountain ranges of Iran and Afghanistan and arrive over NW Pakistan and adjoining western Himalayas in a diffused state with distorted structure, but on arrival over the plains of Pakistan, some of them interact with the pre-existing orographically-maintained trough of low pressure and regain their frontal structure and identity. The interaction leads to strengthening of the wave disturbance. Those developing large amplitude may draw moisture from the Arabian Sea and strengthen further. The rejuvenated disturbance (which is popularly called a ‘Western Disturbance’ in the subcontinent) then generally moves eastward, often in an occluded state, affecting southern Himalayas and adjoining plains of northern India. Some of them travel as far as Bangladesh and Assam and the mountains of eastern Himalayas where they usually break up and disappear. While passing over the plains of northern India, they draw moisture from the Bay of Bengal and are often accompanied by thundery rain and squally conditions. The average lifetime of a western disturbance is 4–6 days and there may be 6–8 disturbances per month at the peak of the season. Farmers welcome winter rain, for it helps in the cultivation of winter crops, such as wheat, etc. There is also heavy snowfall in the western Himalayas, the depth of snow occasionally exceeding 15 metres and causing dangerous avalanches. In the wake of some of these disturbances, dense fog may appear over large tracts of northern India and persist for days, interfering with road, rail and aerial communications. During the passage
4.3
The Transition Season (March–May)
93
of these disturbances, the subtropical westerly jetstream may often reach a speed of 60–75 m s–1 at about 200 hPa over the latitude belt 25–30◦ N.
4.2.2.2 Easterly Waves and Cyclonic Storms of the Northern Indian Ocean A surge in the low-level E/NE-ly tradewinds that converge into the equatorial trough of low pressure over South China Sea often give rise to easterly waves which travel westward. Arriving over the Bay of Bengal, some of them develop into westward-propagating depressions and cyclones. Their formation and movement are clearly seen in day-to-day satellite cloud imagery. Large cloud clusters and rainfall associated with them are observed to move westward a few degrees north of the equator. During their movement, they may strike the coast of Sri Lanka and southern Tamilnadu. Once in a while, they may cross the peninsula and emerge into the Arabian Sea. But such occasions are rare. But not all surges in the tradewind easterlies develop into waves and cyclonic disturbances. Freeman (1948) draws an analogy between these surges and supersonic gas flows and concludes that surges lead to hydraulic jumps in the easterly flow and formation of cloud lines which move downstream over long disturbances.
4.2.2.3 Easterly Waves and Cyclones of the Southern Indian Ocean Westward-propagating easterly waves also form south of the equator when a surge in the low-level SE-ly tradewinds, or the monsoon northwesterlies, converges into the circulation around the SET and increases its cyclonic vorticity. The enhanced vorticity gives rise to westward-propagating cyclonic disturbances. After formation, these disturbances usually move westward as lows or depressions. However, a few of them which survive and enter the western Indian Ocean (west of about 80◦ E) recurve southwestward and develop into tropical storms and cyclones. Several islands of the region including Madagascar and Mauritius are badly hit by these cyclones almost every year. Beyond these islands they recurve further around the subtropical high pressure belt. Some of them are drawn into the circulation of the midlatitude wave disturbances which sweep across South Africa and move eastward across the Southern Indian Ocean towards Australia – New Zealand.
4.3 The Transition Season (March–May) Several important developments take place in the weather over the Subcontinent and the adjoining Indian Ocean during the transition season. These include: (1) Passage of WDs across the northern part of the Subcontinent; and (2) Development of powerful ‘heat lows’ over land and relatively ‘cold highs’ over neighboring oceans.
94
4 Monsoon over Southern Asia and Adjoining Indian Ocean
The equatorial Indian Ocean is, perhaps, one region which becomes most active during this period when summer monsoon withdraws from the southern hemisphere and enters the northern hemisphere. Varieties of interesting phenomena are observed over the equatorial region about this time.
4.3.1 Western Disturbances Western disturbances, described in Sect. 4.2.2.1, continue to be active over the northern part of the subcontinent during the transition season, though their track gradually shifts poleward with advance of the season. During their eastward passage, they draw additional moisture from the Arabian Sea as well as the Bay of Bengal and cause increased convective rainfall on the plains of northern India and extensive snowfall on the mountains.
4.3.2 ‘Heat Lows’ over Land and ‘Cold Highs’ over Ocean During March and April, ‘heat lows’ form over the Indian Subcontinent as well as other land areas bordering the Northern Indian Ocean, such as Somalia and Saudi Arabia in the west, and Malaysia, Thailand and Myanmar in the east, while the adjoining oceans, the Arabian Sea and the Bay of Bengal continue to be under ‘cold highs’. By mid-May, the heat lows over the land areas develop further and move further north to take up their northernmost summer locations. The heat low over the Indian Peninsula also deepens further with its trough at surface oriented in a direction almost paralleling the east coast of the peninsula. While low level convergence and convection occur in the boundary layer of these heat lows, strong subsidence and divergence prevails in the upper troposphere above them. With adiabatic cooling of the rising air and subsidence warming aloft, the result is a stable stratification in the vertical in these heat lows at this time. The scenario changes when, under the influence of the prevailing pressure distribution, low-level winds diverging from the neighboring oceanic high pressure cell inject cool, moist air into the heat low circulation over the land along the coastal belts of Southeastern India and Bangladesh. The injection of moisture at low levels in this manner makes the atmosphere over the region, especially the coastal belts, conditionally unstable.
4.3.3 Severe Local Storms Over central and eastern parts of the Subcontinent where a lot of latent instability energy is stored in the atmosphere due to influx of cool, moist air from the Bay of Bengal at low levels, widespread thunderstorms and hailstorms occur whenever a Western Disturbance affects the region. Locally, in Bengal, these storms are known
4.3
The Transition Season (March–May)
95
Fig. 4.2 Illustrating the formation of a Kal-Baisakhi: (a) A plan view showing locations of lowlevel troughs (double-dashed), (b) a vertical section showing superposition of a W’ly jetstream (J) on low-level trough. DV – divergence, CV – convergence
as ‘Kal-Baisakhis’ or deadly storms of the month of Baisakh (April–May). They are accompanied by thunder and lightning, squally winds, and heavy downpours. Many of them breed deadly tornadoes which inflict heavy loss of life and property. Figure 4.2 shows schematically a typical synoptic situation favorable for occurrence of a Kal-Baisakhi over eastern India and adjoining Bangladesh. The left panel of Fig. 4.2 shows the heat low troughs (Tr), to which cool, moist air from the Bay of Bengal converges and rises in strong penetrative convection when a upper-air W’ly trough approaches the region and has its divergent area associated with upper-level Jetstream (J) superimposed upon the pre-existing trough with lowlevel moisture convergence (right panel).
4.3.4 Developments over the Equatorial Indian Ocean During transition season, summer monsoon withdraws from the southern hemisphere and enters the northern hemisphere. The movement is marked by several interesting events over the equatorial Indian Ocean, including the following: (a) Cross-equatorial movement of Monsoon Circulation; (b) Formation of Equatorial Westerlies, Double Equatorial Troughs and Cloud Bands; (c) Increased Cyclonic and Anticyclonic Activity over the Equatorial Zone (a) Cross-Equatorial Movement of Monsoon Circulation: With increased warming of the earth’s surface and falling of pressure to the north of the equator and cooling and rising of pressure to the south and differential rate of heating between land and ocean, a gradient of pressure tendency develops between the two sides of
96
4 Monsoon over Southern Asia and Adjoining Indian Ocean
the equator, which enables the monsoon current of the southern hemisphere to cross the equator and move into the northern hemisphere to start the process of its advance towards the Indian Subcontinent. An example of this type of forcing for equatorial crossing and relocation of the monsoon current from the winter to the summer hemisphere may be seen in Fig. 4.3 which shows the distribution of mean sea level pressure at 12 GMT on 7 April 2008 when monsoon appeared to have crossed the equator in 2008 in the Western Indian Ocean. According to the isobaric field shown, a pressure gradient exists in the vicinity of the East African coastline not only between the two sides of the equator but also along the equator from west to east. Also, north of the equator, a pressure gradient exists between the ocean and the land across the coast of Somalia. These pressure gradients appear to provide a safe passage for a parcel of air approaching the equator from the south to move along a path indicated by a thick continuous line with arrow in Fig. 4.3. The new locations of the cold sector of the monsoon and related ITCZ (dashed) north of the equator are also shown. (b) Formation of Equatorial Westerlies, Double Equatorial Troughs: The low level winds and circulation over the equatorial Indian Ocean during the transition season show a broad band of equatorial westerlies between two troughs of low pressure, one in each hemisphere. Tradewinds diverging from the oceanic high pressures converge into these equatorial troughs forming penetrative convection and Cloud
Fig. 4.3 Map showing MSLP (mb) at 12 Z, 7 April 2008 over the Indian Ocean and littoral countries Thick continuous line shows how the cool humid monsoon current crosses the equator, with ITCZ (dashed) located to its east (NCEP Reanalysis)
4.3
The Transition Season (March–May)
97
Fig. 4.4 Wind vectors and streamlines over the equatorial Indian Ocean at 10 m above sea surface and ocean surface temperatures (◦ C) at 12 GMT, 14 April 2008. Thick continuous lines show the locations of the troughs (adapted from NCEP Reanalysis)
Bands, one on either side of the equator, during a short period of equatorial crossing of monsoon. Just how this happens is demonstrated by Figs. 4.4 and 4.5 which shows winds and streamlines at 10 m above mean sea level and 850 hPa at 12 GMT on 14 April 2008, approximately a week after equatorial crossing. The double equatorial convergence zones are characterized by formation of double cloud bands, one on each side of the equator, the presence of which is revealed in satellite cloud imagery (see Fig. 4.6). Double equatorial troughs and associated cloud bands are persistent features of the circulations over the equatorial Indian Ocean during the transition season. However, they are short-lived and usually observed around the time of equatorial crossing only. (c) Increased Cyclonic Activity over the Equatorial Eastern Indian Ocean: The spring transition season appears to be the period of maximum cyclonic activity over the equatorial eastern Indian Ocean as well as Arabian Sea (see Fig. 4.7) The reason why the equatorial eastern Indian Ocean is so cyclogenetic during the transition season is to be sought, inter alia, in the warm ocean surface temperature of the Bay of Bengal. Further, most of the cyclonic storms develop around monsoon troughs which interact with traveling disturbances such as E’ly and W’ly waves. After development, the cyclonic disturbances move northwestward and later recurve
98
4 Monsoon over Southern Asia and Adjoining Indian Ocean
Fig. 4.5 Wind vectors and streamlines at 850 hPa showing the locations of two equatorial troughs, one on either side of the equator at 12 Z, 14 April 2008
northward and then northeastward to strike the coasts of West Bengal, Bangladesh and Myanmar. However, a few of them may continue their northwestward movement to strike the coast of Andhra Pradesh before recurving. A few may even cross the peninsula to emerge over the Arabian Sea to affect the west coast of India. A few cyclones may also form and develop over the Arabian Sea itself.
4.4 Advance of Summer Monsoon to the Indian Subcontinent – General Remarks The India Meteorological Department (1943) has traditionally used certain criteria associated with rainfall to determine the normal date of onset of summer monsoon at a place. Ananthakrishnan et al. (1968) who proceeded on that basis and studied the onset of SW monsoon over Kerala found that the synoptic conditions over the Arabian Sea associated with the onset were, in their own words, as follows: (i) A disturbance in the Arabian Sea/Bay of Bengal. The most common initial form of the disturbance is a trough of low pressure in southeast Arabian Sea;
4.4
Advance of Summer Monsoon to the Indian Subcontinent – General Remarks
99
Fig. 4.6 Satellite Cloud Imagery showing double cloud bands, one on each of the equator, over the Equatorial Indian Ocean, at 12 GMT, 14 April 2008
(ii) Reports from ships and island stations in the South Arabian Sea, of heavy convection, squally weather and rough seas or swell from southwest with moderate to strong winds from some southerly to westerly direction; (iii) The strengthening and deepening of lower tropospheric west winds over extreme south peninsula and Sri Lanka and strengthening of upper tropospheric easterlies to 40 Kts for a few days at 14 to 16 km; at the time of onset, the easterlies reach a maximum speed of about 60 Kts. (iv) The tendency of the strong westerlies of the upper troposphere over north India to break up or shift northwards. (v) Persistent moderate to heavy clouding in the south Arabian Sea shown by satellite pictures and its tendency to shift northwards. In a note added to the criteria, the authors state that all these features may not always be present simultaneously. The organization of the circulation to the monsoon patterns extends over an interval ranging from a few days to 1 or 2 weeks. The above-mentioned study by Ananthakrishnan et al. (1968) is very significant in the sense that it marked a welcome re-thinking on the nature of monsoon and
100
4 Monsoon over Southern Asia and Adjoining Indian Ocean
Fig. 4.7 Tracks of cyclonic storms that formed over the Bay of Bengal and the Arabian Sea in May during the 70-year period, 1891–1960 (After Rao, 1981)
its onset. It appeared to recognize for the first time that monsoon was not simply observed rainfall, nor a particular type of wind, but something beyond that, and that we must consider some aspects of the circulation that is associated with the rainfall or the wind to determine the onset. So, the problem of onset of summer monsoon over the Indian Ocean still remains to be addressed. In this regard, in addition to qualitative analysis of monsoon onset in Chap. 1 and the preceding sections, the author examined data and analyses of several meteorological variables over the Indian Ocean at surface and lower and upper-tropospheric levels at 12 GMT daily over a 5-year period (2004–2008) (January–August), available from NCEP Reanalysis. For tracing the movement of the cool, humid air of the monsoon current, stress was laid on the analysis of temperature, pressure, specific humidity, and wind fields of the different layers extending from the ocean surface to 500 hPa. The results of this examination are depicted in Fig. 4.8. For a closer look at the problem of advance, the total period of advance was divided into three stages, as follows: Stage 1. Advance over the Southwestern Indian Ocean, the Arabian Sea, and the Bay of Bengal to the Indian Subcontinent (April–June); Stage 2. Advance over the Indian Subcontinent (June–July); and Stage 3. Advance from the Indian Subcontinent to the Western Himalayas and the Tibetan Plateau (July–August)
4.4
Advance of Summer Monsoon to the Indian Subcontinent – General Remarks
101
Fig. 4.8 Schematic showing advance of summer monsoon from the Indian Ocean. The thick continuous lines with arrows show the routes of the major cross-equatorial flows. Thin continuous lines with wave structure mark the approximate monthly locations of the advancing monsoon current Troughs are indicated by thick double-dashed lines. The monsoon circulation at the beginning of advance showing locations of ITCZ and TCZ (short thick lines) over the Southwestern Indian Ocean is indicated by dotted lines with arrows. ITCZ and TCZ over India are also shown
The following aspects of the advance of monsoon are highlighted by Fig. 4.8: 1. Locations of the main heat lows (L) and their troughs (thick dashed lines) in all the land sectors bordering the Southwestern Indian Ocean, the Arabian Sea, and the Bay of Bengal; 2. The (J-F) location of the equatorial trough of low pressure with cyclonic circulation (dotted lines with arrows) around it over the Southern Indian Ocean with locations of ITCZ and TCZ, before it starts its seasonal movement; 3. Three major monsoon currents (thick bold continuous lines with arrow) transporting cool, humid air of the winter hemisphere to the summer hemisphere, each being a divergent current from a heat sink or a high pressure area. These currents converge into the circulations around low pressure areas on either side, one being the heat low over the neighboring land and the other over the oceanic equatorial heat source (Note that the Continents of Africa, Australia and the Maritime Continent play important roles in this regard). A major cross-equatorial flow of monsoon current is influenced by the trough of the heat low over Peninsular India. It appears to cross the equator in the longitudes of Sri Lanka and flow along the western boundary of the Bay of Bengal close to the east coast of the peninsula;
102
4 Monsoon over Southern Asia and Adjoining Indian Ocean
4. Approximate northern boundary of the advancing monsoon current (thin continuous lines with wave structure) in the northern hemisphere at certain epochs of time (months/dates); in the southern hemisphere, the lines refer to the southern boundary of the retreating monsoon; 5. Advance of the monsoon current along the west coast of Peninsular India as well as the Arakan coast of Myanmar appear to be facilitated by movement of a cold sector of the monsoon wave over the respective region; 6. Locations of the ITCZ and the TCZ (indicated by short thick lines) on either side of the equatorial heat source after its final arrival over the northwestern part of the Indian Subcontinent.
4.4.1 Advance over the Indian Ocean (April–June) – Stage 1 4.4.1.1 Retreat from the Southern Hemisphere The successive monthly locations of the southern boundary of the monsoon wave during the period, January–April, shown in Fig. 4.8, illustrate how summer monsoon withdraws from the southern hemisphere before it enters the northern hemisphere. Note how the heat low over the land sector jumps from its January location over the island of Madagascar to the African mainland and then travels along the western boundary of the Southwestern Indian Ocean to southern Somalia in April and disappears from the southern hemisphere. The equatorial troughs over the ocean also shift northward. By April they move into the northern hemisphere to start the process of advance of monsoon towards the Indian Subcontinent. Summer monsoon advances over the Northern Indian Ocean from different parts of the equator, as shown in Fig. 4.8. However, the first move appears to be made from the side of the Eastern Bay of Bengal. So, we start from that side first. 4.4.1.2 Advance over the Bay of Bengal During late March (M) or early April (A), the increased seasonal warming of the northern hemisphere and cooling of the southern hemisphere in the Southeast Asia sector and differential heating between land and ocean in general deepens a newly-formed heat low over the Malaysia-Sumatra region, which forces the North Equatorial Trough (N.E.T) of low pressure over the Bay of Bengal to move northward along with it. As the heat low crawls slowly northward along the narrow landstrips of Thailand and Myanmar, it carries the W/SW’ly monsoon airstream with it on its southeastern side. Just about this time, under the powerful influence of the heat low over the Indian Peninsula, there is a major cross-equatorial airflow from the western side of the South Equatorial Trough towards the east coast of the Indian Peninsula, a branch of which turns eastward to strengthen the circulation around the northward-moving heat low over the mountain ranges of the Tennaserim coast. Early May, there is another cross-equatorial current from the side of Australia to enter the extreme western part of the South China Sea and flow northward along
4.4
Advance of Summer Monsoon to the Indian Subcontinent – General Remarks
103
the eastern coast of the Malaysia-Thailand peninsula, a branch of which turns northward under the influence of a powerful heat low over Central Myanmar, while the other branch turns eastward to flow around a heat low over Thailand. From here, the S-ly current comes under the influences of several mountain ranges and finally turn towards the plains of northern India to flow as a ESE-ly current around the heat low over India along the foothills of the Himalayas. Needless to state, the cool, moist winds converging at the mountain slopes all along the routes produce heavy rainfall over the mountain ranges. Thus, on the Bay of Bengal side, summer monsoon advances to the Indian Subcontinent via two main routes: (i) A strong cross-equatorial airflow by the side of Sri Lanka and parallel to the east coast of the Indian Peninsula, and (ii) a cross-equatorial airflow from the side of Australia and Indonesia. 4.4.1.3 Advance over the Arabian Sea The Arabian Sea during early summer when the monsoon current crosses the equator is relatively cold compared to surrounding land areas where heat lows form, and remains under a strong temperature inversion with a strong anticyclonic circulation prevailing at low levels. The conditions inhibit further northward advance of the monsoon current. Further, part of the southern-hemispheric tradewinds that cross the equator near the Somali coast at this time appears to be diverted westward by the heat low circulation over the Congo region of Equatorial Africa, leaving only a feeble narrow coastal current to circulate around the heat low over Somalia. But, the situation changes, though gradually at first, when a warming ocean surface and deepening heat lows over land allows a further northward movement of the cool monsoon current towards the Indian Subcontinent. The land-sea thermal contrast across the Somali coast intensifies resulting in development of the so-called Somali jet and intense upwelling along the Somali coast. Similar developments also occur along the coast of the Arabian Peninsula, further north. These developments help the cold monsoon current which binds the heat low over the land to the oceanic heat low to move further northeastward and cover more than half of the Arabian Sea by end of May. So, on or around 1 June, the equatorial trough of low pressure over the northeastern Arabian Sea is so placed as to have its associated ITCZ oriented in a more or less NW-SE direction over mid-ocean and TCZ near the coast of Kerala, the southernmost Indian State, where a S/SW-ly flow is enforced partly by the movement of the cold sector of the monsoon wave and partly by the local orography of the Western Ghats Mountains. From this stage onward, monsoon current moves rapidly northward under the influence of the heat low over India-Pakistan to cover most of the northern Arabian Sea and the Indian Peninsula by June 15. The progress of the monsoon over the ocean and the land during this period is indicated in Fig. 4.9. 4.4.1.4 Weather over the Northern Indian Ocean During Advance of Monsoon Summing up the preceding paragraphs, it may be stated that the conceptual or idealized model of the advance of summer monsoon over the Indian Ocean suggested in
104
4 Monsoon over Southern Asia and Adjoining Indian Ocean
Fig. 4.9 Dates of onset of summer monsoon over the Bay of Bengal and the Indian Subcontinent, as determined by India Met. Dept (Rao, 1981)
the Text appears to account qualitatively for several observed features of the weather over the Bay of Bengal and the Arabian Sea. These include cold SST anomaly, strong temperature inversion, low level jets with highly stormy seas and unusually strong ocean currents, little cloud and rain, and few cyclonic disturbances along and to the west of the cold monsoon current, as against warm SST anomaly, weak temperature inversion, light to moderate winds, heavy clouding and precipitation and a high frequency of cyclonic disturbances, such as depressions and cyclones, to the east. Both ITCZ and TCZ are characterized as zones of cloudy and rainy weather with relatively clear skies in between, confirming the wave structure of the monsoon circulation.
4.4.2 Onset over the Indian Subcontinent (June–July) – Stage 2 The India Meteorological Department (1943) worked out the normal dates of onset of summer monsoon over the Indian Subcontinent from climatological records of observed rainfall from coastal and inland stations. Figure 4.9 shows these dates by isolines over land and dashed lines over the Bay of Bengal. No isolines or dashed lines are drawn over the Arabian Sea. As stated in the preceding subsections, summer monsoon advances over the Indian Subcontinent in two broad airstreams; the Bay of Bengal stream from the southeast and the Arabian Sea stream from the SW, as shown in Fig. 4.8. In their
4.4
Advance of Summer Monsoon to the Indian Subcontinent – General Remarks
105
final position by end of June, these airstreams converge into the circulation around the heat low over northern India, forming the ITCZ on the equatorial side and the TCZ on the poleward side of the so-called monsoon trough. The progressive advance of the two streams to their final destinations over the Subcontinent is given in the map. From SE to NW, the Bay of Bengal branch of the monsoon advances to Bangladesh by 1 June, Bihar by 10 June, East Uttar Pradesh by 15 June, Punjab, Rajasthan, Pakistan and some parts of the State of Jammu and Kashmir by early July. The Arabian Sea branch approaching from the SW crosses the Indian Peninsula by 10 June, Madhya Pradesh by 15 June, and Rajasthan by 30 June. Monsoon remains established over the subcontinent till the end of August. During this period, moderate to heavy rain falls along the ITCZ to the S/SW of the monsoon trough and along the TCZ which runs along the foothills of the Himalayas to the N/NE. Rainfall appears to be deficient over the trough zone in between. However, as shown in the following subsection, the monsoon wave does not remain in this position for long, since it moves further north to Western Himalayas during late July or early August.
4.4.3 Advance to Western Himalayas (July–August) – Stage 3 After summer monsoon gets fully established over the Indian Subcontinent, an extraordinary development takes place further north, which shifts monsoon wave from the plains of northern India to the top of the Himalayan Mountain complex on account of the development of a series of heat lows to the north. It is well-known that during northern summer a heat low develops over the western part of the elevated Tibetan Plateau (e.g., Flohn, 1968; Yeh and Gao, 1979; Murakami and Ding, 1982; Luo and Yanai, 1984; Feng et al., 1984; Murakami, 1987a). Almost simultaneously, heat lows also develop to the north of the mountain complex over the extensive lowlands of Uzbekistan and Kazakhstan to the northwest, the extensive desert lands of the Sinkiang province of China to the north, and the Inner Mongolian region of Northeastern China to the northeast. All these heat lows have their warm anticyclonic circulations above them in the upper troposphere. But on a larger scale, they combine to form a powerful anticyclonic circulation centered over the Tibetan Plateau and extending from the Mediterranean Sea in the west to the Pacific Ocean in the east. It is the development of this giant anticyclonic circulation and its sudden poleward movement, which appears to draw the Monsoon and related Hadley circulations over the Indian Subcontinent within its fold. The poleward movement of the monsoon trough zone to the Himalayas in July–August causes a total re-organization of the associated Monsoon and Hadley circulation cells associated with it over the region. This movement implies a temporary bifurcation of the monsoon wave, one part remaining over the plains of Northern India with somewhat subdued activity, while the other part jumps over to the mountains to the north. The movement simply means a northward shift of the associated Monsoon and Hadley circulation cells, as shown schematically in Fig. 4.10.
106
4 Monsoon over Southern Asia and Adjoining Indian Ocean
Fig. 4.10 Schematic showing mean meridional-vertical circulations and their associated rainbelts (dotted) over the Indian Subcontinent before (upper panel) and after (lower panel) development of the elevated heat source over Western Tibet
4.4.4 Source of Moisture for Monsoon Rainfall Over most parts of the globe, the main source of moisture for monsoon rainfall in the summer hemisphere is the cool, moist tradewinds of the winter hemisphere which after crossing the equator in narrow longitudinal segments converge into the ITCZ of the summer hemisphere (e.g., Simpson, 1921; Findlater, 1969a,b; Saha, 1970). While traveling over the ocean, the tradewinds pick up additional moisture from the underlying ocean surface. A limited amount of moisture is also injected into the TCZ. The cross-equatorial origin of moisture-bearing tradewinds ushering in moisture for monsoon rainfall in the summer hemisphere is evident over several other parts of the globe as well, such as Eastern Asia, Australia, Africa and South America. For details, see the respective chapter on regional monsoon in the present text.
4.5
Disturbances of the Summer Monsoon during the Onset Phase
107
4.5 Disturbances of the Summer Monsoon during the Onset Phase 4.5.1 Onset Vortex over the Arabian Sea and the Bay of Bengal During advance of the summer monsoon wave over the Arabian Sea as well as the Bay of Bengal, the atmosphere becomes dynamically unstable, often resulting in the formation of a cyclonic vortex, whenever it is disturbed by a traveling wave. Such a vortex has come to be known as an ‘Onset Vortex’. In the Arabian Sea, a few studies (e.g., Krishnamurti et al., 1981; Saha and Saha, 1993a) have shown that both barotropic and baroclinic instability may be involved in the genesis of these vortices. After formation, these disturbances usually move in a northerly direction and accelerate the advance of the monsoon current along the West Coast of India. However, a few of them may move westnorthwestward and develop into cyclonic storms over mid-ocean before they hit the coast of Oman and then fizzle out over the sandy Arabian Desert. In the Bay of Bengal, traveling E’ly as well as W’ly waves by their interaction with the quasi-stationary monsoon wave play an important role in the formation of an onset vortex. The formation of such a vortex may upset the normal schedule of advance of monsoon by either accelerating or delaying it by a few days.
4.5.2 Monsoon Depressions and Cyclonic Storms During the monsoon onset phase, June to August, a number of depressions and cyclonic storms form in the monsoon trough zone over the Bay of Bengal and the Arabian Sea. A few also form over the land area adjoining the Head Bay of Bengal. Table 4.1 gives the total number of depressions and storms that formed over these areas during an 80-year period, 1891–1970 (After Rao, 1976). The tracks of these disturbances are shown in Fig. 4.11. After formation, most of the disturbances move in a WNW direction and yield heavy precipitation over their SW quadrant. Several States in India, such as Orissa, Andhra Pradesh, Southern Bihar, Jharkhand, Chhatishgahr, and Madhya Pradesh, receive a significant proportion of their annual rainfall during the passage of these Table 4.1 Number of monsoon depressions (D) and cyclonic storms (S) during the 80-year period (1891–1970) in June, July and August over different areas Area
Bay of Bengal Arabian Sea Land area
June
July
August
D
S
D
S
D
S
71 18 12
35 15 1
107 9 39
38 3 1
132 2 42
26 2 0
108
4 Monsoon over Southern Asia and Adjoining Indian Ocean
Fig. 4.11 Tracks of depressions and cyclonic storms that formed over the Bay of Bengal and the Arabian Sea in July during the 70-year period, 1891–1960 (after Rao, 1981)
disturbances. By contrast, little rain falls to the northeast of the trough axis. In the Arabian Sea sector, few lows and depressions form in this period. Those which form over the northeastern corner of the sea usually move in a northerly direction. Their contribution to annual rainfall is normally negligible. However, occasionally, midtropospheric cyclones form over the northeastern corner of the Arabian Sea and add significantly to coastal rainfall. More than 80% of all Bay disturbances during the onset phase (June–August) formed over the latitude belt, 17.5–22.5◦ N, and moved westnorthwestward.
4.5.3 Interaction of Monsoon with W’ly Waves During summer, large-amplitude W’ly waves moving across the Himalayan region, between about 30 and 50◦ N, interact with the circulation around the quasi-stationary monsoon trough over the Indian Subcontinent as well as with traveling monsoon disturbances, forming an extended trough between the two regions. On such occasions, a ridge of high pressure with anticyclonic circulations prevails over northwestern and Central India and weather remains dry over these areas (see Fig. 4.12). During the period of its eastward movement, the extended trough causes a relocation of the east-west oriented monsoon trough and its associated rainbelt which now lies across southern India. Thus, the interaction simply causes a redistribution of monsoon rainfall with two belts of heavy rain, one in the mountains in the north and other over Southern India in the south and a wide area of little or no rain in between over Central India. As the W’ly wave moves further eastward across the Himalayas taking the extended trough along with it, the belt of heavy rain also shifts eastward. A period of 3–5 days may be taken for the rainbelt to move across the mountains from west to east.
4.5
Disturbances of the Summer Monsoon during the Onset Phase
109
Fig. 4.12 Schematic showing interaction of a W’ly wave trough with: (a) the monsoon trough over the Indian Subcontinent, and (b) a Depression (D) over the Bay of Bengal. Heavy rainfall areas are hatched. Troughs are double-dashed
The W’ly wave troughs moving across the eastern Himalayas also interact with westward-propagating depressions over the Bay of Bengal forming an extended trough with it. In this case also, we have two areas of heavy rainfall, one in the north over the eastern Himalayas and the other over an extensive area of central and southern India. In fact, the ‘break monsoon’ situation over Central India almost disappears in this case. The windward slopes of the Western Ghats Mountain as well as the Arakan Yoma experience heavy rainfall on these occasions.
110
4 Monsoon over Southern Asia and Adjoining Indian Ocean
4.6 Rainfall over the Indian Subcontinent during the Onset Phase Several factors appear to contribute to the distribution of monsoon rain over the Indian Subcontinent during the onset phase. They include: (1) Orography, (2) Heat Lows, (3) Location of the monsoon trough, (4) Depressions and cyclones, and (5) Travelling wave disturbances. More than any other factor, orography appears to contribute the most to the distribution of monsoon rains over the Indian Subcontinent, as would be evident from Fig. 4.13, which records the highest concentrations of rainfall on the windward slopes of the mountains, and scanty rainfall on the leeside. According to Fig. 4.13, the Western Ghats Mountains of the Indian Peninsula, the Arakan Yoma and the Shan States of Myanmar, and the Mountains of Eastern Himalayas, especially in Bangladesh, Assam and Arunachal Pradesh experience heavy rainfall, while the leesides of these mountains have deficiency of rainfall. Along the foothills of the Himalaya Mountains, there appears to be a general decrease of rainfall from east to west. A rainfall maximum appears over Orissa and adjoining Central India to the southwest of the monsoon trough zone. While the eastern and southern parts of the Indian Subcontinent receive substantial rainfall during the onset phase of the summer monsoon, the northwestern part especially the Thar Desert area of Pakistan and India suffers from shortage of rain because of the presence of the heat low circulation over the region.
Fig. 4.13 Ten-year (1976–1985) mean July rainfall (unit: 10–5 kg m–2 s–1 ) over the Indian Subcontinent. The double-dashed line shows the location of the monsoon trough (after Saha and Saha, 1996)
4.7
Summer Monsoon – Withdrawal Phase (September–November)
111
The distribution of rainfall shown in Fig. 4.13 agrees substantially with that of normal summer monsoon rainfall during the period, June–September, as determined by the India Meteorological Department (not shown).
4.7 Summer Monsoon – Withdrawal Phase (September–November) 4.7.1 Dates of Withdrawal of Monsoon By and large, summer monsoon withdraws from the Indian Subcontinent and the Northern Indian Ocean by following the same route as it did during advance (Fig. 4.9), but in the reverse direction. After reaching its peak intensity during July–August, monsoon starts withdrawing from the Western Himalayas in early September. The process begins with the filling up of the heat low over the elevated Tibetan Plateau and the re-establishment of the heat low over the northwestern part of the Indian Subcontinent. However, the transition takes place very gradually and almost imperceptibly for a while. The withdrawal from northwestern India is marked by a weakening of the heat low and associated monsoon trough, reversal of the low-level wind from southerly to northerly and a decrease of rainfall. The change ushers in a regime of somewhat cooler and drier air from the north. Figure 4.14 gives isolines of the dates of withdrawal of the summer monsoon from India, as worked out by the India Meteorological Department on the basis of the climatological distribution of rainfall. According to Fig. 4.14, the SW monsoon pulls out of nearly half the subcontinent by 1 October when its northern boundary runs from the hills of Uttar Pradesh to the middle of the West coast of India. By 15th October, it moves further southeastward so as to have its western end over the middle of the Indian peninsula and the eastern end over central Myanmar or even further south. From this stage onward, its southward movement over land and ocean is very slow, and it is not until the end of November that monsoon totally withdraws from the Indian peninsula and reaches the latitude of Sri Lanka. The Myanmar branch of the monsoon continues to move south/southeastward to reach eventually its winter location as NET within about 5◦ of the equator in the eastern Indian Ocean.
4.7.2 Retreating Monsoon Rain over Tamil Nadu During the stage of withdrawal of summer monsoon from India, the area along the east coast of the Indian Peninsula, especially the state of Tamil Nadu, exposed to the moisture-laden ENE-ly winds at low levels, experiences retreating monsoon rains. In literature, this rain is often projected as winter monsoon rain. However, in reality, it is summer monsoon rain during its withdrawal phase. The mechanism of this rain is illustrated schematically in Fig. 4.15.
112
4 Monsoon over Southern Asia and Adjoining Indian Ocean
Fig. 4.14 Normal dates of withdrawal of the SW monsoon from the Indian subcontinent (after Rao, 1981)
Fig. 4.15 Illustrating the mechanism for formation of a low level moisture convergence zone and rain in Tamil Nadu during monsoon withdrawal phase
4.7
Summer Monsoon – Withdrawal Phase (September–November)
113
Following the withdrawal of the monsoon trough from northern India, two high pressure cells with anticyclonic circulations develop over land, one over Central India and the other over Myanmar. Between their anticyclonic circulations, a trough of low pressure forms over the North Bay of Bengal (see Fig. 4.15). Now, the wind diverging from the high pressure cell over central India has a limited sea travel before it approaches the coast of Andhra Pradesh and Tamil Nadu from N/NE-ly direction and its moisture content is low. On the other hand, the wind diverging from the high pressure cell over Myanmar has first to flow northwestward over North Bay and then turn cyclonically around the oceanic trough southwestward towards the Indian peninsula. Thus, it has a long sea travel and is fully saturated with moisture by the time it arrives at the Tamil Nadu coast where it converges into the circulation around the high pressure cell over Central India. The resulting moisture convergence along the Andhra Pradesh-Tamil Nadu coast (indicated by a thick continuous line) along with the prevailing upper-air divergence over the area is responsible for producing the rainfall over Tamil Nadu. Heaviest rainfall occurs at the mountains of Tamil Nadu facing the moist winds.
4.7.3 Disturbances of the Withdrawal Phase 4.7.3.1 Western Disturbances Closely following the southward movement of the equatorial trough, the belt of the subtropical westerlies shifts southward and W’ly waves (WDs) follow a more southerly track. Their influence on weather over the northern part of the Indian Subcontinent, especially western Himalayas, increases. They also influence the track of cyclonic disturbances which move up from the Bay of Bengal and the Arabian Sea towards the mountain.
4.7.3.2 Depressions and Cyclonic Storms During the monsoon withdrawal phase, there is a spurt in cyclonic activity over the Bay of Bengal and the Arabian Sea. A larger percentage of the depressions develop into cyclonic storms and their places of origin shift continually equatorward from September to November, as shown by statistics over a 70-year period, 1891–1970 (Rao, 1981). It is estimated that in November, 90% of the depressions and cyclones formed over a wide area bounded by latitudes 7.5 and 15◦ N and longitudes 77.5 and 100◦ E. This continued southward movement and the widening of the area of cyclonic activity would be evident from Fig. 4.16 which shows the tracks of these disturbances, when it is compared with Fig. 4.11 for July. These changes appear to be characteristic features of the monsoon withdrawal phase. During the same 70-year period, the number of depressions and cyclonic storms that formed in the Arabian Sea was 21 and 15 respectively. About 50% of them formed over the area bounded by latitudes, 10–12.5◦ N and longitudes, 62.5–75◦ E.
114
4 Monsoon over Southern Asia and Adjoining Indian Ocean
Fig. 4.16 Tracks of cyclonic storms that formed over the Bay of Bengal and the Arabian Sea in October during a 70-year period, 1891–1960 (after Rao, 1981)
In the Bay of Bengal, most of the depressions and cyclones after initial travel over the sea in a WNW direction recurved to a N/NE-ly direction after reaching the latitude belt, 15–20◦ N. Those amongst these which turned into severe cyclonic storms and entered land devastated many coastal belts. Low-lying river deltas which are particularly vulnerable in this regard suffered enormous losses due to high winds, storm surges and torrential precipitation. Table 4.2 lists a few of the deadliest cyclones on record which took heavy toll of lives through storm surge drowning during the monsoon withdrawal phase.
Table 4.2 Statistics of some killer tropical cyclones with record storm surge Cyclone
Date
Surge (m)
Death-toll
Bangladesh (Buckergunge) Calcutta (Kolkata) Bangladesh Andhra Pradesh
November 1, 1876 October 5, 1864 November 13, 1971 November 17, 1977
9.0–12.0 12.0 7.0 5.0
100,000 50,000 300,000 9000
4.8
Variability of the Indian Summer Monsoon Rainfall
115
4.8 Variability of the Indian Summer Monsoon Rainfall Observations show that the summer monsoon rainfall over India for the period, June–September, is not steady but varies on time scales ranging from a few days to a year or several years. Since large-scale variability often leads to disastrous floods and droughts which cause miseries to people and affect the economy of the country, the problem of variability has been studied extensively ever since the time of Blanford (1884, 1886) who first initiated the study. He was followed by Walker (1910a, b, 1914) and several scientists in India and abroad (e.g., Mooley, 1975, 1976; Hahn and Shukla, 1976; Kanamitsu and Krishnamurti, 1978; Bhalme and Mooley, 1980; Angell, 1981; Shukla and Paolino, 1983; Mooley and Parthasarathy, 1983, 1984; Rasmusseen and Carpenter, 1983; Parthasarathy, 1984). To date, there has been a large body of literature on the subject of variability of Indian summer monsoon rainfall. An excellent review of some of the recent studies has been provided by Mooley and Shukla (1987) as well as by Krishnamurti and Surgi (1987) and readers interested in details of these various studies may refer to the original papers mentioned in the references.
4.8.1 Interannual Variability For studying the rainfall variability, different workers used different sets of data. Mooley and Parthasarathy (loc. cit.) used data from a network of rain gauge stations which were fixed and evenly distributed over the country (one raingauge station per district) covering a period of 108 years from 1871 to 1978 (this was later extended to 114 years from 1871 to 1984) but by excluding the hilly stations (where rainfall depended upon height and was of a different pattern from rest of the stations) from the existing network. The season of rainfall considered in these studies is from June to September. On statistical tests, they found the series of rainfall data used by them to be homogenous. Mooley and Parthasarathy (loc. cit.) used the following statistical criteria for describing the various parameters of the variability.
4.8.1.1 Statistical Criteria
Mean or average (x) =
Standard deviation (SD) σx =
i=n i=1
i=n
xi /n
i=1
(4.8.1) 1/2
(xi − x) /(n − 1) 2
(4.8.2)
116
4 Monsoon over Southern Asia and Adjoining Indian Ocean
Co-efficient of variation (CV) = Standard deviation × 100/Mean = (σx /x) × 100 (4.8.3)
Inerannual variability =
i=n−1
[xi+1 − xi ]/(n − 1)
(4.8.4)
i=1
where xi denotes the monsoon season rainfall for the ith year, x (x underlined) is the average rainfall of the total number of years n, and σ x is the standard deviation of the monsoon rainfall. Mooley and Parthasarathy describe the monsoon rainfall in terms of a standard unit, which is equal to deviation from mean divided by the standard deviation. Figure 4.17 shows the all-India summer monsoon rainfall in standard units for each year of the period, 1871–1984. According to Mooley and Parthasarathy, a small deviation up to 5% on either side of the long-term average value can be considered as normal or average rainfall. Some of the statistical properties of the all-India summer monsoon rainfall and its long-term variability as given by Mooley and Parthasarathy (loc. cit.) are given in Table 4.3.
Fig. 4.17 All-India summer monsoon rainfall in standard units (deviation from normal divided by standard deviation) for each year during the period, 1871–1984 (after Mooley and Parthasarathy, 1984)
4.8
Variability of the Indian Summer Monsoon Rainfall
117
Table 4.3 Statistical properties of the all-India summer monsoon rainfall, 1871–1984 Property
Value
Property
Value
Mean Mean/annual Median Lower quartile Upper quartile
852 mm 78.1 864 mm 800 mm 908 mm
Highest Rainfall Deviation from mean Range Coefficient of variation Mean interannual variation as % of mean
1017 mm 19% 413 mm 9.7% 11.9
Standard deviation Lowest rainfall (1877) Mean interannual variation in terms of standard deviation Deviation from mean
83 mm 604 mm 1.22
–29%
4.8.1.2 Floods and Droughts As should be expected, large excess (deficit) of rainfall in a year leads to large-scale floods (droughts) in India. Mooley and Parthasarathy (1983) from the results of their study lay down the following criteria for occurrence of these events: Flood if [(xi − x)/σx ] > 1.28, Drought if [(xi − x)/σx ] < 1.28 Using the above criteria, they worked out the years of large-scale droughts and floods in India during the period, 1871–1984, and identified the following years as those of worst flood and drought years in India: Floods Droughts
1892, 1917, 1956, and 1961 1877, 1899, 1918, and 1972
4.8.2 Factors Likely Responsible for Interannual Variability Since the time of Blanford (1886), there has been frantic search on a global scale for parameters which likely influence the interannual variability of the Indian summer monsoon rainfall. Studies have pointed at several factors around the globe and even at extra-terrestrial ones, with varying degrees of correlation, but the relationship with most cases have turned out to be fragile and undependable. Only a few have passed the rigorous tests of confidence. To date, the following factors appear to be promising in this regard: (a) (b) (c) (d)
Dates of onset and withdrawal of summer monsoon; Eurasian snow cover; Sea surface temperature, El Nino and Southern oscillation; Soil moisture, vegetation and albedo of the earth’s surface
118
4 Monsoon over Southern Asia and Adjoining Indian Ocean
Fig. 4.18 Dates of onset of summer monsoon over southern Kerala, 1901–1984, showing the mean date and the limits of one standard deviation on either side of the mean (Mooley and Shukla, loc. cit.)
(a) Variability in Dates of Onset and Withdrawal of Monsoon: During a period of 84 years, 1901–1984, there was a large variability in the dates of onset of summer monsoon in Kerala, the southernmost state of the Indian peninsula, as may seen from Fig. 4.18 (Mooley and Shukla, loc. cit.). According to Fig. 4.18, the long-term mean date of onset of summer monsoon over southern Kerala is 2 June with a standard deviation of 8 days. Most of the dates of onset lie within two standard deviations. The extreme dates are 11 May 1918 and 18 June 1972. However, as Mooley and Shukla remark, ‘it is strange that both these extremes occurred in drought years. While a late onset in 1972 may be consistent with a drought in that year, the same cannot be said about an early arrival in 1918. This inconsistency only highlights the nature of the problem and the fact that an early or late arrival or departure of monsoon alone is not uniquely related to the overall behaviour of the monsoon during a season. Other factors may be involved in determining the observed variability of the rainfall.’ (b) Eurasian Snow Cover: Blanford (1884) was, perhaps, one of the first to point out that excessive snowfall in the Himalayas during the winter and spring was prejudicial to the subsequent monsoon rainfall over India. His observation was substantiated by Walker in 1910 and the inverse relationship was made use of in monsoon rainfall forecasts. However, on account of uncertainty in observations of snow cover, the use of this parameter was discontinued after 1950. The advent of earth-orbiting satellites which started observing the earth’s snow cover changed the situation. Wiesnet and Matson (1976) on the basis of snow cover data furnished by satellites commented that ‘the December snow cover for the northern hemisphere was a very good predictor of the following January–March snow cover’. Subsequent studies (e.g., Hahn and Shukla, 1976; Dickson, 1983, 1984) found an apparent inverse relationship of the Himalayan snow cover with the Indian summer monsoon rainfall.
4.8
Variability of the Indian Summer Monsoon Rainfall
119
In regard to the suspected inverse relationship between the two parameters, Mooley and Shukla (1987) observes: ‘An inverse relationship between the Eurasian snow cover and summer monsoon rainfall is understandable, since large and persistent Eurasian snow cover would substantially reduce the rate of heating of the concerned land masses during spring and summer and thus delay the onset of the monsoon and prevent normal monsoon activity. It may be mentioned that the snow cover area as measured by the satellite is not quite a representative parameter to assess the amount of snow. The snow cover area in two cases may be the same but depths may be quite different, and the impact on monsoon would be quite different in the two cases’. It may be hoped that with rapid development of satellite technology which would enable both horizontal and vertical extent of snow cover to be measured, a reliable correlation will be found between Eurasian snow cover and the Indian summer monsoon rainfall. (c) Sea Surface Temperature, El Nino and Southern Oscillation: Need for an intensive study of the variation of sea surface temperature (SST) in the context of the Indian summer monsoon rainfall came to the fore after Bjerknes (1969) demonstrated the linkage between the ocean and the atmosphere and interpreted the Southern Oscillation and the Walker circulation in terms of air-sea interaction and El Nino and La Nina events in the equatorial Pacific Ocean. He found that in general a positive SST anomaly in the equatorial western Pacific was associated with heavy rainfall and a negative anomaly with deficient rainfall over India. In a La Nina year, the equatorial eastern Pacific ocean is cold (SST anomaly negative), but the equatorial western Pacific including a part of the eastern Indian ocean, lying between about 70 and 160◦ E, is warm with a positive SST anomaly. With such a distribution of SST anomaly, the Southern oscillation index is positive and the Walker circulation has its ascending branch over the equatorial western Pacific and the descending branch over the equatorial eastern Pacific. The result is normal or heavy rainfall over the western Pacific including the India-Australia sector and drought condition over the eastern Pacific. The situation reverses in an El Nino year when the warm anomaly in the SST shifts to the equatorial eastern Pacific Ocean with a cold anomaly over the western Pacific. In such a year, the rainbelt shifts to the eastern side of the ocean and the India-Australia sector experiences drought conditions with abnormally deficient rainfall. Following the investigations of Bjerknes (loc. cit.), there was a spurt in studies of Indian summer monsoon rainfall in relation to the zonal anomaly of SST in the Pacific Ocean. Several studies (e.g., Angell, 1981; Mooley and Parthasarathy, 1983, 1984; Sikka, 1980a; Rasmussen and Carpenter, 1983) were undertaken to find the degree of relationship of the Indian summer monsoon rainfall with the southern oscillation and the El Nino events by working out the correlation between them. Each of these studies used data for different years and often different parameters. Angell (1981) who used data for the period 1868–1977 obtained a correlation coefficient around –0.6 between all-India monsoon rainfall and SST anomaly over the equatorial eastern Pacific ocean (0–10◦ S, 90–180◦ W) one to two seasons later. The relationship is highly significant. Mooley and Parthasarathy (1983) who used
120
4 Monsoon over Southern Asia and Adjoining Indian Ocean
a dataset for the years, 1871–1978, found a significant relationship between the monsoon rainfall over India and the El Nino events in the equatorial Pacific. They considered 22 moderate and severe El Nino events and found that in all the severe El Nino years, the all-India monsoon rainfall in standard units was less than –0.60, with the exception of 1884 when it was +0.92. However, on careful examination, they found that in that particular year, 20.4% of the country had, indeed, experienced drought conditions as would be expected in an El Nino year, but the deficiency was offset by heavy rainfall in the remaining parts of the country caused by other more influential factors. In another investigation (Mooley and Parthasarathy, 1984) using the same dataset examined the relationship between the Indian monsoon rainfall and the SST anomaly in the Pacific in 3-monthly periods preceding or following the monsoon season and found inverse relationships that were significant at the 1% level for each of the JJA, SON and DJF seasons, and at the 5% level for the MAM season. They found the relationship to be consistent and stable. Sikka (1980a) used a different set of data. He used the Line Islands precipitation data as indicator of El Nino events in the Pacific and sought to relate them with the monsoon rainfall over India. He found a general association of El Nino events with deficient rainfall over India. Rasmusson and Carpenter (loc. cit.) found that in 25 El Nino years, the Indian monsoon rainfall was below the median rainfall in 21 years and below the mean in 19 years. They thought that the association had some predictive value. Shukla and Paolino (1983) examined the relationship between the Indian monsoon rainfall and composites of normalized Darwin pressure anomalies (3-month running mean) for heavy monsoon rainfall years and deficient monsoon rainfall years during the period, 1901–1981, and found that the tendency of Darwin pressure anomaly before the monsoon season was a good indicator of the subsequent monsoon rainfall over India. According to them, a negative tendency between the DJF (December–February) season and the MAM season was found to be associated with good monsoon rainfall years and a positive tendency with poor monsoon rainfall years. The results of their investigation are shown in Fig. 4.19. They expressed the view that whenever the tendency showed a large negative value, non-occurrence of drought in India could be predicted with a very high degree of confidence. (d) Soil Moisture, Vegetation and Albedo of the Earth’s Surface: There appears to be a symbiotic relationship between the albedo of the earth’s surface on one hand and soil moisture, vegetation and rainfall on the other. The atmosphere over a region with high albedo tends to make up for the loss of solar radiation by large-scale subsidence which inhibits precipitation and perpetuates conditions which lead to high albedo. This is likely to happen particularly over the dry desert regions of the subtropics and the marginal lands where an increase of albedo caused by overgrazing or deforestation can lead to long-term reduction of precipitation, and continuation of dry desert conditions. On the other hand, a lowering of albedo by afforestation and vegetation can help to reduce subsidence and thereby promote more convection and precipitation to occur over the region. Afforestation and vegetation thus allowing long-term increased precipitation helps to increase soil moisture which in turn
4.8
Variability of the Indian Summer Monsoon Rainfall
121
Fig. 4.19 Composite of the normalized Darwin pressure anomaly (3-month running mean) for years of heavy (good) monsoon rainfall and deficient (poor) monsoon rainfall (after Shukla and Paolino, 1983)
reduces albedo and increase rainfall. The validity of these cyclic processes has been well demonstrated by a series of numerical experiments using GCMs by Charney (1975); Charney et al., (1977); Shukla and Mintz (1982); Sud and Smith (1985) and several others. Besides the global and regional factors mentioned above, attempt has been made to link the variability of the Indian summer monsoon rainfall with such extraterrestrial factors as solar activity as judged by the sunspot numbers. Some of the original studies in this direction were carried out by Gilbert Walker (1915a,b,c). However, the results of his studies showed no significant or consistent relationships with the mean annual rainfall over India.
4.8.3 Intraseasonal Variability The variability of monsoon rainfall on time scales ranging from a few days to several weeks are caused by the familiar westerly and the easterly waves that move across the Indian longitudes and interact with the quasi-stationary monsoon wave, and also by low-frequency intraseasonal oscillations of the atmosphere during northern summer. 4.8.3.1 Variability on Scale of 3–7 Days – Active and Break Monsoons IActive and break monsoon cycles occur frequently during the monsoon season. They have drawn the attention of meteorologists over a long time. Defining a
122
4 Monsoon over Southern Asia and Adjoining Indian Ocean Table 4.4 Statistics of break in summer monsoons, 1888–1967
Month
No. of breaks
No. of break (days)
July August 1 July–31 August
53 55 5
306 356 47
Most frequent Average duration Longest break duration (days) (days) (days) 5.8 6.5
17 20 21
4 3
monsoon break as a disappearance of the monsoon trough over India from mean sea level and 850 hPa maps for at least a couple of days at a stretch, Ramamurty (1969) catalogued the breaks in July and August from 1888 to 1967. His statistics are presented in Table 4.4. Ramamurty found from his long-period statistics that August is slightly more susceptible to ‘break days’ and longer breaks, particularly around the middle of the month. He also found that in the long record, there was no break in 12 years. Another important effect of the W’ly wave trough on the monsoon is a north/northeastward recurvature of the track of a monsoon depression or cyclone from its usual west/northwestward track if it happens to come under the influence of the W’ly trough. 4.8.3.2 Variability on 30–50 Day Time Scale Atmospheric oscillations on this time-scale are known as Madden-Julian Oscillations (MJOs) after the name of their discoverers (Madden and Julian, 1971, 1972) and are believed to be excited by large-scale tropical convection. All subsequent studies (e.g., Murakami, 1976; Yasunari, 1979; Sikka and Gadgil, 1980; Murakami, 1984) have suggested a strong relationship of these oscillations with the slow meridional movement of a zonally-oriented low-level trough of low pressure associated with penetrative convection, cloudiness and heavy rainfall from the equatorial region to higher tropical latitudes in the Indian monsoon region Krishnamurti and Subrahmanyyam (1982) found that active and break monsoon cycles over the Indian longitudes were closely coupled to the meridional movement of these low-level troughs associated with rainfall. Oscillations on this time scale over the equatorial latitudes which moved slowly towards the pole have also been noted in the relative angular momentum of the earth’s atmosphere from the datasets of the zonal wind by Rosen and Salstein (1981). Interestingly, they also found a strong signal on this time scale in the variations of the length of the day computed from lunar laser ranging observations.
Chapter 5
Monsoon over Eastern Asia (Including China, Japan, and Korea) and Adjoining Western Pacific Ocean
5.1 Introduction The study of monsoon and related weather phenomena over Eastern Asia has a long history. Prior to the 3rd century B . C ., it was mostly the farmers who watched the weather seriously and maintained some kind of an ‘agricultural calendar’ of climatic events in connection with agricultural operations. In some central parts of China, these agricultural calendars are still in vogue, though other parts have opted for more modern methods. The modern instrumental period may be said to have begun about the close of the 19th century, but the observing network was very limited in the beginning and confined mostly to densely populated areas. Vast areas were uncharted. It is only recently from about the middle of the twentieth century that the observational network over the region as a whole has improved. Since 1959, a network of surface and upper-air observing stations was established on the highly elevated plateau of Tibet. It is mentioned that during the period, 1949–1963, the number of meteorological observing stations in China increased 30-fold (Cheng, 1963). A Chinese national project on the meteorology of the Tibetan Plateau was reported upon by Yeh and Gao (1979) who along with their many colleagues carried out excellent studies of the heat budget of the plateau and other related problems of the high-altitude region. The observational network on the plateau was further improved upon during a special experiment known as the Qinghai-Xizang Plateau Meteorology Experiment (QXPMEX) during the summer of 1979 which was conducted by Chinese scientists as part of the Global Weather Experiment, 1978–1979. Earlier, during the winters of 1974 and 1975, a GARP field project under the leadership mostly of Japanese scientists had conducted an ‘Air Mass Transformation Experiment’ (AMTEX) over the sea areas southwest of Japan to learn more about the energy and momentum exchanges between the sea and its overlying atmosphere and meso-scale cellular convection and cyclogenesis that occurs when there is a cold air outbreak over the East China Sea and the Kuroshio current. Like the Indian Ocean, the South China Sea and the Western Pacific Ocean play important roles in monsoon circulation over Eastern Asia, the Maritime Continent and the Australian region, especially during advance and retreat of monsoon current across these ocean areas. In order to learn more about these K. Saha, Tropical Circulation Systems and Monsoons, C Springer-Verlag Berlin Heidelberg 2010 DOI 10.1007/978-3-642-03373-5_5,
123
124
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
roles, the International Community led by Chinese scientists mounted an impressive array of field experiment called the South China Sea Monsoon Experiment (SCSMEX) in 1998, spanning the period from 1 May to 31 August, for carrying out intensive observations of surface and upper air parameters relating to monsoons. A wealth of new information was collected from this experiment, which became available for further studies (e.g., Lau et al., 1998, 2000). The recent studies of monsoon over Eastern Asia are, therefore, based on an excellent coverage of data, though in some areas long-period data are still lacking. An excellent review of some of the recent studies which were carried out on the seasonal march of the East Asian Summer Monsoon has been provided by Ding (2004). It is well-known that during the peak summer months of July-August, monsoon over the tropical belt of Eastern Asia suddenly jumps to extratropical latitudes to cover such areas as Northern China, Northern Japan, Korea and Eastern Siberia with its poleward boundary near about 60◦ N. We study monsoon over this extratropical belt of Eastern Asia in the latter part of this chapter.
5.2 Physical Features and Climate It is not easy to delineate the southern boundary of Eastern Asia which includes some of the most heterogeneous elements of the global terrain and features practically the whole range of global climate from tropical to arctic. Here, along the southwestern boundary of mainland China, the mighty Himalaya mountain complex with the world’s highest mountain peak, Mount Everest, rising to an altitude of about 8.85 km a.s.l. and associated Tibetan plateau with a mean elevation of well over 4.6 km a.s.l. stand guard over the extensive lowlands of Northern and Eastern China which have several smaller high ground or hill ranges scattered all over the region. The Tibetan Plateau descends steeply both northward and eastward to the plains of China and this is clearly indicated by the direction of flow of water of the two mighty rivers, the Yantzekiang and the Hwang-Ho which flow in a zig-zag course eastward to the China Sea. Also, along the northwestern boundary of China lie a series of high-rise mountains, the Tien Shan and the Altay mountains, and several other lesser mountain ranges which extend northeastward to as far north as 60◦ N or even beyond. A relief map of Eastern Asia showing the above-mentioned topographic features is at Fig. 5.1. Another important physical feature of China which exercises great influence upon the climate of the region is an extension of the vast Central Asian desert lowlands from Sinkiang in the west to the Gobi desert or even beyond to Manchuria and Eastern Siberia in the east. The Korean peninsula lies over the northeastern part of the region and juts out southward so as to have the Yellow Sea to its west and the Sea of Japan to the east. The Korean Strait separates the peninsula from the Islands of Japan which lie to the south and east. Besides the mainland, several large and small islands belonging to China and Japan lie scattered over the western North Pacific Ocean.
5.3
The Winter Season over Eastern Asia (November–March)
125
Fig. 5.1 Relief map of Eastern Asia
The topography and the geographical location of Eastern Asia are responsible for a wide variety of climatic conditions in terms of temperature, pressure, airflow and rainfall. The seasons also are somewhat different here from those over the Indian Subcontinent. On account of more northerly location and greater continental and oceanic influences, the winter season starts early in November and lasts till the end of March. Summer monsoon starts in May and lasts till the end of September. The transition periods are usually April and October.
5.3 The Winter Season over Eastern Asia (November–March) 5.3.1 Temperature, Pressure, and Wind Mean air temperatures over Eastern Asia start falling rapidly from October onward and by January extremely low temperatures often dipping to a minimum of < –30◦ C may prevail over the Gobi desert of Outer Mongolia and adjoining eastern Siberia as well as over the Korean peninsula (Fig. 5.2). In response to the temperature distribution, an extremely high pressure cell builds up over the region with maximum pressure exceeding 1032 hPa centered over the
126
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
Fig. 5.2 Mean air temperatures (◦ C) over Eastern Asia in January (after Watts, 1969)
Mongolian region and a steep pressure gradient to the south and east to cover practically the whole of Asia and a good part of northwestern Pacific Ocean close to the coast of Eastern Asia. Side by side, a deep low pressure cell develops in the vicinity of the Aleutian Islands area with the subtropical high pressure cell of the Pacific Ocean lying to its south with ridge along about 25◦ N. Consistent with the prevailing temperature and pressure distributions described in the preceding para, there is strong anticyclonic circulation over a vast region of Eastern Asia and adjoining Pacific Ocean where it appears to merge with the subtropical anticyclonic circulation with its axis along about 30◦ N. A strong cyclonic circulation prevails over the Aleutian Islands area. However, the circulations change rapidly with height, with westerlies dominating the flow at 500 and 200 hPa over the subtropical and midlatitude belts. These aspects of the airflow at low levels and at 500 and 200 hPa over Eastern Asia and adjoining Pacific Ocean during January are shown in Fig. 5.3 (Crutcher and Meserve, 1970). Several studies (e.g., Yeh and Gao, 1979; Murakami, 1981a,b; Boyle and Chen, 1987) have emphasized the great mechanical and thermodynamical influence of the Himalayan Massif and Tibetan Plateau on the upper air circulation over Asia, especially Central and Eastern Asia where the winds are predominantly westerly above
5.3
The Winter Season over Eastern Asia (November–March)
127
Fig. 5.3 Streamlines showing mean atmospheric circulation over Eastern Asia and adjoining western Pacific Ocean during winter: (a) Low-level (925 mb), (b) 500 mb, and (c) 200 mb. Thick continuous line in (a) shows the NE–SW oriented convergence line over the Western North Pacific Ocean
the low-level E/NE-ly tradewinds. During winter, upper-level midlatitude westerlies migrate southward and blow around the Himalayan mountain complex and the Tibetan plateau. On striking the western side of the mountain barrier, the flow appears to divide itself into two parts, one flowing northward around the northern boundary of the mountain block and the other flowing southeastward around the southern boundary. The divided aircurrents appear to merge on the leeside over China, some distance away from the eastern side of the mountains. An interesting aspect of the midtropospheric circulation over the region is that it is the weakest (<10 m s–1 ) directly over the Tibetan Plateau at about 500 mb.
128
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
Murakami (1981a) attributes this feature to the frictional effects of the elevated Tibetan Plateau. By comparing the geostrophic wind with the observed wind, he estimated the height of the boundary layer over the plateau to be about 1.5 km above the plateau. Yeh and Gao (1979) showed that the height of the boundary layer over the plateau was substantially higher during summer than winter. Above the boundary layer, the westerly airstream over the plateau appears to strengthen with height. However, the speed appears to increase downstream to reach a maximum of 60 m s–1 or more on the leeside, about 1000–1500 km eastward of the plateau. This is shown by the 200 mb wind field. The existence of this high-speed Jetstream over the eastern part of Asia and adjoining northwestern Pacific Ocean has been known since the days of Second World War (1942–1945) when aerial missions flying high over Japan encountered these high winds.
5.3.2 Quasi-stationary Wave in Westerlies – Its Interaction with Traveling Waves – Cold Surges It is well-known that the low level flow around the northern boundary of the Himalayan mountain complex during the winter maintains a stationary wave in the airflow with a trough along about 70◦ E in the west and a trough along about 110◦ E in the east and a ridge of high pressure in between along about 90◦ E. This being a baroclinic zone, normally at midtropospheric height, there is convergence and descending motion to the west and divergence and rising motion to the east of the trough axis. Frequently, however, extraordinary developments take place in the stationary wave when eastward-propagating synoptic-scale baroclinic wave disturbances from western Asia move into East Asia and interact with the stationary wave trough. When the waves are in phase, the interaction leads to an amplification of the stationary trough, resulting in intense cold surge at low levels from the high pressure area in the rear of the trough. According to Lau and Chang (1987), it is customary to define a cold surge event by noting changes of one or a combination of the following indicators within a period of 24–48 h: (i) A drop in surface temperature at a station of 5◦ C or more; (ii) An increase of the surface pressure gradient between coastal and central China of at least 5 mbs; and (iii) A prevalent northerly surface flow over the South China Sea with speed exceeding 5 m s–1 . Although the above-mentioned criteria are somewhat arbitrary and several variations are also in use, they have been found to be of practical importance in defining a cold surge event.
5.3
The Winter Season over Eastern Asia (November–March)
129
The southward-moving cold surge usually converges into the circulation around the equatorial trough of low pressure over the Maritime continent where the moisture convergence produces heavy clouding and precipitation along the convergence zone (see Fig. 5.3a). However, occasionally, a branch of the southward-moving cold surge crosses the equator over several segments of longitude between about 105◦ E and the dateline and converges into the ‘heat low’ circulation over Australia. There is indication in upper-tropospheric flow that air rising in convection over the equatorial trough zone or the trough zone over Australia diverges in the upper troposphere and returns to the northern hemisphere after flowing around an anticyclonic circulation with its ridge along about 15◦ N at 500 mb and converges into the upper-tropospheric baroclinic trough over East Asia. The convergence appears to strengthen the pre-existing convergence on the western side of the trough axis, thereby leading to increased subsidence and inducing a fresh low-level cold surge. Also, adding westerly momentum to the flow because of increased Coriolis acceleration, results in strengthening the existing jetstream. The cold surge thus appears to constitute the equatorward-moving lower branch of a vertical circulation which links tropical convection over equatorial latitudes to midlatitude baroclinic waves. In case, the zone of tropical convection lies in the southern hemisphere, i.e., over Australia, the above circulation may be said to be a truly interhemispheric circulation. The importance of the meridional-vertical overturning associated with lowlevel cold surge in producing upper-level W’ly jetstream over northern China and
Fig. 5.4 Ten-year (1961–1970) average rainfall (mm) over Eastern Asia in January. Heavy dashed line indicates the 3 km-height (after Yeh and Gao, 1979)
130
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
adjoining Japanese islands has been pointed out by several studies (e.g., Murakami, 1981a, 1987b; Lau and Chang, 1987).
5.3.3 Winter Rainfall over Eastern Asia It appears that winter disturbances moving in the baroclinic westerlies around the Himalayan mountain complex produce appreciable rainfall on the windward slopes of the mountains (see Fig. 5.4).
5.4 Airmass Transformations and Cyclogenesis over the Oceans Observations indicate that quite a few of the W’ly wave troughs that arrive over Eastern Asia move out to the neighboring ocean areas and develop into intense cyclones under favorable conditions. New ones also develop over the oceans. For a long time, lack of adequate observations over the oceans hindered a thorough analysis and understanding of these disturbances. In 1974 and 1975, the World Meteorological Organization (1981), with the active participation of Chinese and Japanese scientists, carried out a comprehensive observational program, called the Airmass Transformation Experiment (AMTEX), over the East China Sea with the long-term objective of learning more about these oceanic cyclones. Figure 5.5, due to AMTEX (1973) Study Group, shows the distribution of mean ocean surface temperature (◦ C) off the eastern seaboard of Asia in February from 1953 to 1957, with a temperature maximum along the Kuro-shio current. When a cold surge flows out over a relatively warm ocean, there is increased evaporation from the ocean surface and a boundary layer of warm, moist airmass builds up downstream due to turbulent mixing. Further, the warming and moistening of the lower layer lead to increased convective instability and vertical growth of the moist layer and formation of cumulus and stratocumulus clouds and occasional light rain. But when the aircurrent enters the zone of the warm Kuroshio Ocean current or comes under the influence of an eastward-propagating disturbance, the development sometimes is explosive.
5.4.1 Cyclonic Disturbances over Eastern Asia and Neighboring Ocean Several studies (for example, Boyle and Chen, 1987; Lau and Chang, 1987) have traced the movement of W’ly waves from East Asia across the eastern seaboard of Asia during northern winter and shown how they develop into explosive cyclones when they move out over the East China Sea and the Sea of Japan and neighboring North Pacific Ocean.
5.4
Airmass Transformations and Cyclogenesis over the Oceans
131
Fig. 5.5 The mean sea surface temperature (◦ C) off the eastern seaboard of Asia in February from 1953 to 1957 (Study Group on AMTEX, 1973)
Using a Hovmoller diagram for high-pass-filtered (period less than or equal to 5 days) 500 mb height, Lau and Chang (1987) demonastrated how a W’ly wave disturbance from near Lake Baykal on entering the Sea of Japan developed into an explosive cyclone. Here, we present the results of two studies which throw light on different aspects of these explosive cyclones. These are shown in Fig. 5.6 (Trewartha, 1961), and Fig. 5.7 (Sanders and Gyakum, 1980). The study by Trewartha (1961) reports the frequency of cyclone formation over East Asia and the neighboring ocean areas, viz., East China Sea, Sea of Japan and the Northwestern Pacific ocean in the months October through April, 1932–1937. It shows that only a few waves develop into cyclones while they are over land, but the frequency of cyclone formation increases rapidly as they move out to the sea with a maximum over the latitudinal belt, 30–35◦ N, the highest maximum of near 28 being over the East China Sea. However, the frequency decreases further eastnortheastward where the ocean surface gets somewhat less warm. Explosive cyclones form mostly during the winter season and their effects are felt over a wide belt of latitudes poleward of about 25◦ N. They are almost non-existent during the summer when the region comes under the influence of a heat low over North China and the tracks of these disturbances shift northwestward over northern Siberia.
132
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
Fig. 5.6 Frequency of cyclone formation over ocean areas off the eastern seaboard of Asia. The value of any isoline at any point represents the number of cyclones that formed within a radius of 2.5 latitude degrees from that point in the months October through April, 1932–1937 (after Trewartha, 1961)
The study is by Sanders and Gyakum (1980) throws further light on the occurrence and distribution of explosive cyclones which they call bomb events. The results of their study pertain to two selected regions of the globe, viz., the oceanic areas off the eastern seaboards of Asia and North America, during three cold seasons. Sanders and Gyakum define an explosive cyclone as a cyclone in which the deepening rate is, on an average, the geostrophic equivalent of at least 1 mb(h)–1 over a 12-h period at 45◦ N. After analyzing several cases of such cyclogenesis, which they thought were like bomb explosions, in the northern hemisphere during the period September through May 1979, they found that the areas where they occurred with the highest frequency lay over oceans in the vicinity of the east coasts of the continents of North America and Asia, especially where the offshore winds were about to enter the zone of warm ocean currents. Raw non-zero frequencies appear in each 5◦ × 5◦ quadrilateral of latitude and longitude. Isopleths represent smoothed frequencies, obtained as one-eighth of the
5.5
Transition Period (April)
133
Fig. 5.7 Distribution of explosive cyclones in Pacific and Atlantic basins
sum of four times the raw central frequency plus the sum of the surrounding raw frequencies. The columns of numbers to the left and right of the heavy line along longitude 90◦ W represent, respectively, the normalized frequencies for each 5◦ latitude belt in the Pacific and Atlantic basins, using a normalization factor of (cos 42.5/cos ϕ). Heavy dashed lines represent the mean winter position of the Kuroshio and the Gulf Stream (after Sanders and Gyakum, 1980).
5.5 Transition Period (April) With gradual warming of the land surface from March onward, the rigour of the winter monsoon over Eastern Asia lessens and cold surges become less frequent. The NE’ly winds gradually withdraw poleward and are replaced by S/SW-ly winds which blow around a shallow low pressure that forms over southern China with incursions of warm, moist air from the southern and eastern oceans. In other words, this is the spring season over the region. However, a large part of the ocean surface is still cold compared to land surface, and the passage of the southerly winds which after flowing over the warm Kuro-Siou current blows over this cold water produces widespread stratus with drizzle and fog over the coastal waters, which is locally known as ‘crachin’. Crachin may persist over the coastal waters for days and
134
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
weeks. In the south, fogs are most frequent during March and gradually become less frequent as the ocean surface slowly warms up with the advance of summer. In April, the eastward-propagating baroclinic waves in the westerlies occasionally interact with the low pressure system developing over Southern China and cause light to moderate precipitation over Southern China during the period of interaction.
5.5.1 Development of ‘Heat Low’ over Eastern Asia Temperatures rise rapidly over Southern China from April onward resulting in the formation of a heat low over the region. However, the formation of a heat low over china is not an isolated event, nor is it a sudden development. It is part of a chain of heat lows that form during northern summer over all the land sectors of Southern Asia, extending from the Arabian Peninsula in the west to the Pacific Coast in the east. However, as explained in the case of the Indian Subcontinent, heat lows play the important role of shifting monsoon from the winter to the summer hemisphere. In this regard, Eastern Asia is no exception. However, the actual process in the case of Eastern Asia is explained in the following section.
5.6 Origin of Monsoon over Eastern Asia The summer monsoon over Eastern Asia appears to originate from movement of the equatorial trough of low pressure over the equatorial eastern Indian Ocean, in two distinct phases, the first from movement of the North Equatorial Trough (NET) during March-April, and the second from movement of South Equatorial Trough (SET) in May. At the beginning of summer, the NET is located within about 5◦ N of the equator, while the SET to the south is located farther away from the equator, leaning towards Australia where it links with the heat low over the continent (see Fig. 4.8). In the first phase, i.e., in March-April, the movement of the heat low over the landmass of the Sumatra-Malayasia region forces the NET to move northward along with it across the narrow landstrips of Malaysia–Thailand–Myanmar complex. As it reaches the northern part of the narrow landstrip, the circulation around it gets bifurcated into two airstreams, one heading northwestward towards a quasi-stationary heat low over Myanmar, and the other northeastward towards the heat low over Thailand and Vietnam and from there to Southern China. On arrival over China, the equatorial trough of low pressure merges with the trough of the seasonal heat low over Southern China and starts the first phase of monsoon season over Southern China during April. Monsoon activity over Southern China gets a boost in early May when in the second phase the cold, humid airstream associated with the SET arrives over the region under similar forcing.
5.6
Origin of Monsoon over Eastern Asia
135
Fig. 5.8 Schematic showing how summer monsoon advances to Southern China and other neighboring areas. Thick dark lines show the main supply routes for cool, moist airmass of the southern hemisphere. L denotes Low, H – High. Locations of ITCZ, TCZ, and Meiyu–Baiu front (thick lines) north of the monsoon trough
It is well-known that during southern summer, the SET extends southeastward over the southern Indian Ocean towards the heat low over Australia. It has then the cool NW-ly monsoon currents on its equatorward side and the cold SE-ly tradewinds of the southern-hemisphere on its poleward side, with the warm trough in between. With change of season, the heat low over Australia weakens and loses its hold on the SET which then starts shifting equatorward across the Indonesian islands to reach the equator by early May. As the heat low moves across the islands, it narrows
136
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
down and the orientation of its trough veers from NW-SE direction to almost W-E direction, allowing the cool, humid airmass of the winter hemisphere to flow first as equatorial westerlies and then cross the equator to blow as S/SW-ly winds of the South China Sea to usher in monsoon to Southern China. However, several studies (e.g., Findlater, 1969a,b; Saha, 1970; Wang and Leftwich, 1984; Sun, 2002) have emphasized that the main artery for supply of moisture for Southern Asia rainfall including that over Eastern Asia is that via the Somali Monsoon Current in the western Indian Ocean. Additional moisture supplies for rainfall over Eastern Asia are those from cross-equatorial airflow via routes over the Maritime Continent and from the tradewinds originating in the subtropical high pressure over the western Pacific Ocean, as shown schematically in Fig. 5.8.
5.7 Seasonal March of the Summer Monsoon As in the Indian Ocean region, climatological rainfall has been used as the main criterion for determining the dates of onset of summer monsoon over China. Figure 5.9, after Tao and Chen (1987), presents the dates of onset of summer monsoon over East Asia along with those over India by Rao (1976). According to Fig. 5.9, monsoon sets in over Eastern Asia a couple of weeks earlier than over the Indian Subcontinent. Starting on or about 10 May at the southern coast of China, it advances slowly and reaches Central China by mid-June. A faster advance appears to occur thereafter during July when monsoon advances to
Fig. 5.9 Dates of onset of summer monsoon over Eastern Asia (Tao and Chen, 1987) along with those over the Indian Subcontinent (Rao, 1981)
5.9
Meteorological Developments Associated with the Jump to Central China
137
Northern China and engulfs practically the whole of mainland China, part of Korea and the islands of Japan and the Kuerile Islands. However, in recent years, questions have been raised as to the justification for using rainfall as the sole criterion for determining the date of onset of monsoon, and additional parameters such as airmass properties have been used for this purpose. For example, Fong and Wang (2001) used an equivalent potential temperature value of 335 K and a SW-ly wind greater than 2.5 m s–1 at 850 hPa for this purpose. Earlier, Ding (1994) had used a value of 340 K for the equivalent potential temperature at 850 hPa. The trend to include airmass properties besides rainfall and wind direction for identification of monsoon would appear to be in the right direction, since monsoon in reality is not rainfall alone, nor is it only a reversal of the prevailing wind direction, as ordinarily believed. In the present text, we have defined monsoon as a perturbation of the tradewind circulation associated with the seasonal movement of the equatorial trough of low pressure, which converges into the circulation around the trough, producing rainfall along the ITCZ and the TCZ and several other changes in airmass properties.
5.8 Stationary States and Jumps Several studies (e.g., Fong and Wang, 2001; Wang and Lin, 2002; Sun, 2002; Ding, 2004) have identified three distinct stationary states and two significant jumps in the seasonal march of summer monsoon over Eastern Asia. The stationary states and their approximate periods are: (1) Southern China (approx. 18–25◦ N), from second week of May to middle of June; (2) The Yangtze River basin (25–30◦ N) (Meiyu front), from mid-June to mid-July; (3) North China (North of 30◦ N), from mid-July to mid-August The jumps are: (1) From Southern China to the Yangtze River basin, around the second week of June; and (2) From Yangtze River basin to North China, about mid-July.
5.9 Meteorological Developments Associated with the Jump to Central China After monsoon gets established over Southern China during April-May, two important developments occur in the heat budget of Eastern Asia, which appear to force a re-organization of the monsoon currents and shift monsoon activity further
138
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
northward. These are: (1) formation of large-scale heat source over the elevated Tibetan Plateau and (2) a further deepening and northward movement of the heat low from Southern to Central China. These developments result in the formation of a Plateau Monsoon over the eastern part of the Plateau on one side, and development of the Meiyu–Baiu Front over Central China and adjoining Pacific Ocean on the other.
5.9.1 Tibetan Plateau Monsoon The monsoonal characteristics of the boundary layer over the Tibetan Plateau has been brought out by several studies (e.g. Gao and Xu, 1962; Flohn, 1968; Gao, 1976, 1979; Yeh and Gao, 1979; Tang, 1979; Tang and Reiter, 1984). Tang and Reiter (1984) in a comprehensive study compared the monsoonal behaviour of the atmosphere over the Tibetan Plateau with that over the Western Plateau of North America and found that inspite of their altitude and latitude differences, the boundary layer of the two Plateaux behaved in very similar fashion. Both served as a heat source with low pressure, high absolute humidity and low static stability during summer and a heat sink with high pressure and relatively dry conditions during winter. These seasonal reversals had the effect of producing a well-developed warm, wet monsoon during the summer half of the year and a cold dry monsoon during the winter half. For their study, they selected 600 mb surface to represent conditions over the elevated Tibetan Plateau, and 800 mb for the Western Plateau. In the case of the Tibetan Plateau, the seasonal variation is well reflected in the distributions of temperature, pressure, absolute humidity, and atmospheric circulation over the plateau and its surrounding areas at the same altitude. As example of pressure variation, we present in Fig. 5.10 (after Yeh et al., 1979 and Gao et al., 1981), the distribution of monthly mean 600 mb height contours over a 10 year period, 1961–1970, during (a) winter, and (b) summer. It clearly reveals a high pressure system over the Plateau during January, as against a series of low pressure systems during July. Dashed lines are Troughs, solid lines-ridges, dotted lines – outline of Tibetan Plateau (after Yeh and Gao, 1979; Gao et al., 1981; Tang and Reiter, 1984) An important aspect of the pressure distribution during July is that between the troughs of low pressure systems over the eastern part of the Plateau and a welldefined ridge of high pressure system which extends eastward to Southern China, a southerly monsoon current enters the region and rises against the eastern slopes of the Plateau, causing extensive rainfall and other types of disturbed weather. A study by Tao and Ding (1981) shows that with the availability of additional moisture from the monsoon flow, the low pressure systems and their attendant shear lines over the Plateau have the effect of producing severe rainstorms, often accompanied by hail. After formation over the Plateau, many of these rainstorms are found to move eastward and generate new perturbations to affect low-lying regions, such as the Yantzekiang valley and adjoining areas.
5.9
Meteorological Developments Associated with the Jump to Central China
139
Fig. 5.10 Mean 600 hPa height contours (decameters) over the Tibetan Plateau and surrounding areas: (a) January and (b) July
5.9.2 The Meiyu (Plum Rain) Front over China 5.9.2.1 Formation of the Meiyu Front The merging of the equatorial heat low with the quasi-stationary heat low over Southern China towards the middle of May prepares the ground for development of the ITCZ on the equatorial side of the trough of the heat low and the TCZ on its poleward side. It is along these convergence zones that the tradewinds of the two hemispheres meet the heat low circulation over China. The TCZ which forms over the latitudinal belt 25–30◦ N and which runs along the Yangtze River basin constitutes the Meiyu front over China (see Fig. 5.8).
140
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
Monsoon activity with copious rainfall remains largely confined to the ITCZ over Southern China during the first stationary period from about the middle of May to middle of June. Relatively less activity is noticed along the TCZ during this period except when it interacts with eastward-propagating westerly baroclinic waves and produces heavy rainfall. However, the situation changes abruptly towards the middle of June when a branch of the Pacific Ocean anticyclonic circulation is drawn into the heat low circulation over Eastern Asia and it finds its way to the TCZ over Central China. The injection of cool, moist air from the Pacific Ocean now shifts the activity to the TCZ where heavy rainfall occurs whenever eastward-propagating disturbances in baroclinic westerlies interact with it. It marks the beginning of the Meiyu season along the Yangtze River basin. The other branch of the Pacific Ocean anticyclonic circulation veers eastnortheastward towards the islands of Japan, converging into the cyclonic circulation around the ‘Aleutian low’ pressure area. Thus, a field of deformation forms over Central China and adjoining East China Sea between the various aircurrents involved and the TCZ forms a part of the dilatation axis of this deformation field. It runs almost zonally along the Yantzekiang river valley and is characterized as a zone of heavy rainfall, or ‘plum rain’ during the period from mid-June to mid-July. The eastern part of the dilatation axis that extends towards Japan is known as the ‘Baiu’ front. At the Baiu front, the aircurrents that converge are the cyclonic circulation around the Aleutian Low and the anticyclonic circulation around the Pacific Ocean subtropical high pressure cell.
5.9.2.2 Interaction of the Meiyu Front with Traveling Disturbances The intensity of the Mei-yu front and associated rainfall fluctuate considerably when eastward-propagating baroclinic wave disturbances interact with it. The interaction leads to formation of vortices of different scales of motion which travel along the front and bring about heavy rainfall (Tao and Ding, 1981; Ninomiya and Akiyama, 1992; Chang et al., 1998, 2000; Ding, 1994; Fong and Wang, 2001; Wang and Lin, 2002). According to Tao (1980), it is the southeast quadrant of the vortex where heavy rainfall is concentrated. Several studies (e.g., Chen and Tsay, 1978; Tao, 1980) show that meso-scale vortices of wavelengths about 50, 80 and 150 km are very effective mechanisms for producing heavy rainfall in the Mei-yu front. Severe rainstorms in the Mei-yu front can cause devastating floods in the Yangtze River valley. However, there appears to be an interesting relationship between the strength of the SW monsoon current over China and Mei-yu rainfall. It is found that it is only in years of normal or weak monsoon that heavy rainfall occurs along the Mei-yu front during early summer. In years of strong monsoon, the moist monsoon current rushes to the northern part of China where it rains heavily, while the Yangtze River valley is deprived of its normal share of rainfall.
5.9
Meteorological Developments Associated with the Jump to Central China
141
5.9.2.3 Structure of the Meiyu–Baiu Front Chen and Chang (1980) studied the north-south temperature and humidity gradients and the vorticity budget of a Meiyu system over Eastern Asia and neighboring ocean during a period of 6 days, 10 through 15 June 1974, separately for its (a) mature and (b) decaying stages. For this purpose, they divided the front into three zonal sections: a western section (W) over continental China; a central section (C) over East China Sea; and an eastern section (E) over Southern Japan. They computed the temperature and vertical motion profiles across these sections and found that the structure of the Meiyu front over China (the W section), was substantially different from that of the Baiu front which extended over the oceans (the C and E sections) during both mature and decaying stages of the system. Their findings are in Figs. 5.11 and 5.12. According to their finding, the Baiu front had a typical baroclinic structure with cold air descending behind a sloping frontal surface and warm air ascending in front, as would be found in a typical midlatitude frontal system, while the Meiyu front behaved more like a tropical system. The results of their computations are presented in two vertical cross-sections across the fronts in Fig. 5.11 for the thermal structure, and Fig. 5.12 for vertical motion structure.
Fig. 5.11 Vertical cross-sections of the deviation of temperature (T’) from the cross-sectional mean at each level in sections, W (west), C (center) and E (east) during (a) the mature stage, and (b) the decaying stage. Dashed lines indicate the trough axes. The abscissa is the distance in degrees latitude from the trough axis (Chen and Chang, 1980)
142
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
Fig. 5.12 Vertical cross-sections of pressure velocity, ω (mb s–1 ), in sections W, C, and E during (a) the mature stage, and (b) the decaying stage. Dashed lines indicate the trough axes. The abscissa is the distance in degrees latitude from the trough axis (Chen and Chang, 1980)
The above-mentioned findings appear to be implicit in the deformation field leading to the formation of the Meiyu–Baiu front in Fig. 5.8. In the Meiyu front over Central China, three distinct aircurrents with different airmass properties are involved. These are: (1) the Cold, dry NW’ly winds blowing anticyclonically from the high pressure cell over Mongolia; (2) the cyclonic circulation around the heat low circulation over Southeastern China; and (3) the relatively cool, humid SE’ly tradewinds blowing out of the Pacific Ocean subtropical high pressure cell. While the cold NW’ly winds tend to converge into the circulation around the heat low producing the Meiyu front, the Pacific trades also converge into it producing a marked moisture tongue along the frontal zone. So, structurally, the Meiyu front is characterized by conditions which are partly tropical, and partly extratropical. On the other hand, the Baiu front appears to have different characteristics. Here, from East China Sea to Southern Japan and even further eastward, the airmass contrasts across the front appear to be well-defined and have the characteristics of an extratropical front.
5.10 Jump of East Asian Monsoon to Extratropical Latitudes As mentioned in Sect. 5.8, the activity of summer monsoon over Central China rapidly shifts to Northern China and adjoining land and ocean areas of the extratropical belt, including Korea, northeastern Siberia, and Japan, to as far north as
5.10
Jump of East Asian Monsoon to Extratropical Latitudes
143
near latitude 60◦ N during the months, July–August. It is interesting to note that this is almost the same period when the Indian Summer Monsoon suddenly shifts to the Western Himalayas. An examination of the NCEP Reanalysis climatology reveals that the jump involves a simultaneous movement of the monsoon current in meridional as well as zonal directions.
5.10.1 Evidence of Jump in Climatological Fields The movement is well reflected in the changes that occur in the distributions of Mean Sea Level Pressure, Mean Air Circulation at 925 hPa, and Mean Rainfall over the region between the periods DJF and JJA, shown in Figs. 5.13, 5.14 and 5.15 respectively.
Fig. 5.13 Mean sea level pressure over Eastern Asia and adjoining Pacific Ocean during: (a) DJF and (b) JJA (from NCEP Reanalysis)
According to Fig. 5.13, the advance is signaled by replacement of a high pressure cell of the winter season over Northeastern Asia by a low pressure cell. In fact, during JJA, two low pressure cells appear over the Chinese Mainland, one over Southern China and the other over Northern China.
144
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
The field of low-level circulation shown in Fig. 5.14 appears to be consistent with the seasonal change in pressure.
Fig. 5.14 Mean wind circulation (Streamlines-isotach analysis) at 925 hPa over Eastern Asia during: (a) DJF and (b) JJA
Figure 5.15 shows that there is more rain over the ocean than over the land in winter (DJF). But the field reverses during summer (JJA) when the rainfall maximum lies over the land with lesser amounts over the ocean.
5.10
Jump of East Asian Monsoon to Extratropical Latitudes
145
Fig. 5.15 Mean precipitation rate (mm day–1 ) over Eastern Asia and adjoining Pacific Ocean during (a) DJF and (b) JJA
5.10.2 Zonal Anomaly in Seasonal Variations Whereas the northward jump of the summer monsoon activity from Central to Northern China has been documented in considerable detail by Chinese and Japanese meteorologists (e.g., Lau and Chang, 1987; Ninomiya and Murakami, 1987; Tao and Chen, 1987; Ding, 2004), the evidence furnished in the preceding section suggests that the jump also involves a zonal movement. This is clearly suggested by Fig. 5.16 which shows the distribution of seasonal mean sea level pressure P and rainfall rate R during DJF and JJA along latitude 42◦ N across North Korea and China.
146
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
Fig. 5.16 Zonal anomaly (deviation from annual mean) of MSLP (P) and precipitation rate (R) along latitude 42◦ N across the coast of Eastern Asia
5.10.3 Climatological Rainfall over Eastern Asia During July In the early stages of summer monsoon over China, rainfall occurs largely along the Yangtze River valley. After monsoon shifts to Northern China in mid-July, the main rainbelt shifts to areas north of the Huang-Ho River valley. Also, during June-July, there is a north-south oriented belt of heavy rainfall along the eastern slopes of the Tibetan Plateau caused by Plateau Monsoon. These aspects of monsoon rainfall appear to be brought out by distribution of 10-year mean July rainfall, presented in Fig. 5.17.
Fig. 5.17 Ten-year (1961–1970) mean July rainfall (mm) over Central and East Asia. The heavy dotted line indicates the 3-km topographic contour height (from Yeh and Gao, 1979)
5.11
Monsoon over Japan
147
5.11 Monsoon over Japan 5.11.1 Geographical Location and Climate Japan consists of a series of islands, of which four, viz., Kyushu, Shikoku, Honsu and Hokkaido are large and the rest are small. Situated in close proximity of the eastern seaboard of Northern China and adjoining Eastern Siberia at the northwestern corner of the Great Pacific Ocean, it is separated from Mainland Asia by the Sea of Japan to the north and the Yellow Sea and East China Sea to the west and south. In view of its geographical location, it is affected by the land-sea thermal contrast between mainland Asia and the Pacific Ocean in both the seasons. In winter, it comes under the influence of the cold, relatively dry NW-ly winds from the land, while in summer it is affected by the relatively cool, moist S/SW-ly winds of the Pacific Ocean. The islands are affected by the eastward-propagating baroclinic wave disturbances during the winter, while it is the movement of the Baiu front and typhoons that brings weather to the islands during the summer.
5.11.2 The Baiu Front – Its Seasonal Movement and Activity The monsoon over Japan is associated with the movement and activity of the Baiu front which forms the extended part of the Meiyu front over the northwestern part of the Pacific Ocean in the vicinity of Southern Japan. As stated earlier in this chapter, this part of the front coincides with the dilatation axis of a deformation field over Eastern Asia and the adjoining Pacific Ocean. During winter, the Baiu front lies way out over the ocean south of the main islands of Japan and the rainfall activity and cyclogenesis are mainly over the open ocean where they affect a few small islands only. But in summer it rapidly moves northwestward towards the heat low over North China and adjoining Eastern Siberia and affects both Japan and Korea. According to observation, the southwestern part of Japan experiences Baiu rainfall about a month earlier than the central and northeastern parts. Heavy rainfall occurs along this front whenever extratropical disturbances of different scales moving eastward interact with it. During the period from mid-June to mid-July, the disturbances move across Japan and are active over the islands between latitudes 30 and 40◦ N. The seasonal movement of the front from south to north across Japan is revealed by the dates of maximum rainfall averaged over a 30-year period, 1951–1980, at a few selected stations, as given by Ninomiya and Murakami (1987). It is evident from these rainfall figures that as the front moves northward from the oceanic island station of Naha towards mainland Japan, the amount of rainfall increases, becoming maximum at Kagoshima which is located in the southwestern part of Japan and then rapidly decreases at stations further north. It is during this period that the southern parts of the Korean peninsula also get their summer monsoon rainfall.
148
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
Several observational studies (e.g., Matsumoto and Ninomiya, 1971; Ninomiya and Akiyama, 1971, 1974; Matsumoto, 1972; Ninomiya,1978; Ninomiya and Yamazuki, 1979; Ninomiya et al., 1984) find that the low-level southwesterly jetstream associated with a well-developed medium-scale depression in the Baiu front is significantly intensified in the southeast quadrant where the vertical motion is strongly upward and rainfall is maximum. Ninomiya (1978) illustrates these developments in the Baiu front in the case of a medium-scale disturbance which developed into a depression at 09 LST on 27 June 1972 near 35◦ N, 132◦ E, shown in Fig. 5.18.
Fig. 5.18 Southwesterly low-level jetstream: (a) temperature (dashed lines), geopotential height (full lines), and winds (arrows) at 850 mb. (b) Mixing-ratio (full lines) and winds (arrows) at 900 mb at 9 LST, June 27, 1972 (from Ninomiya, 1978)
5.11
Monsoon over Japan
149
The sequence of development of the depression is shown in Fig. 5.19(a, b, c). It is evident from Fig. 5.19 that the southerlies inject a large amount of moist oceanic air into the frontal zone, which when lifted by the strong upward motion appears to lead to the observed heavy rainfall on 27 June 1972.
Fig. 5.19 Time series of several meteorological parameters from 21 LST, June 25, to 3 LST, June 28, 1972: (a) The vertical p-velocity (ω) and the zonal wind component relative to the depression’s movement (ur ); (b) Meridional component (v) of the wind, with shading indicating regions of southerlies; and (c) Hourly rainfall (after Ninomiya and Yamazaki, 1979)
150
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
5.12 Monsoon over Korea 5.12.1 Historical Background Early meteorological records of Korea were maintained by the Japan Meteorological Agency, since the land was under occupation of Japan. But from 1953 onward, the peninsula became independent and was divided into North Korea and South Korea, each having its own meteorological service. The dividing line was along about 38◦ N parallel. In 1963, the Meteorological Service of South Korea at Seoul published a detailed Climatological Atlas of Korea, based on earlier observations from 14 South Korean stations and 11 North Korean stations up to 1960. This Atlas remains the main source of information regarding the climate of the peninsula.
5.12.2 Physical Features and Climate The climate of Korea (here we mean both the Koreas) is shaped largely by its geographical location and physical environment. To the north and northwest of the peninsula, the varied landforms consisting of deserts, mountain ranges and extensive Plains undergo large seasonal variations in temperature, being extremely cold during the winter and hot during the summer. On the south side, the peninsula is surrounded by ocean, i.e., the Sea of Japan to its east, the Yellow Sea to the west and the East China Sea to the south. The narrow Strait of Korea separates it from the islands of Japan. Compared to the seasonal variations of temperature over the land, those over the ocean are small. So, it is the differential heating between land and ocean which plays a significant role in atmospheric circulation over Korea.
5.12.3 Winter Monsoon over Korea During northern winter (DJF), extremely cold conditions with mean air temperatures often dipping to –40◦ C and a powerful high pressure cell prevail over Eastern Siberia and Mongolia to the north and northwest of the peninsula (see Fig. 5.2). Strong anticyclonic winds blowing around the high pressure cell (Fig. 5.3a) advect extremely cold conditions from the north and northwest, which lower air temperatures over the peninsula to sub-freezing levels, as shown by isolines of mean January air temperatures on the basis of observations from 150 stations spread over a period of about 20 years by McCune (1941) who divided the peninsula into 10 climatic regimes (Fig. 5.20).
5.12
Monsoon over Korea
151
Fig. 5.20 Mean January air temperatures in Korea delimiting 10 climatic regimes (after McCune, 1941)
The sub-freezing temperatures lead to extensive snowfall over Korea between October and April when the peninsula comes under the influence of eastwardpropagating baroclinic wave disturbances in the westerlies. Starting from the north in October, the snow-belt extends southward till by December it covers the whole peninsula. It starts retreating in March and the process is completed by end of April. The mean dates of first and last dates of snowfall are given by the Climatic Atlas of Korea, published by the Central Meteorological Office, Seoul, in 1962.
5.12.4 Summer Monsoon over Korea – Changma Season The Meiyu–Baiu front which moves northward across China and Japan in late June and early July also affects the Korean Peninsula and produces heavy rain over the southern parts of the Peninsula, ushering in the Changma season in Korea. The mean annual precipitation (mm) of Korea, as given by the Climatological Atlas of Korea (1962), is presented in Fig. 5.21.
152
5
Monsoon over Eastern Asia and Adjoining Western Pacific Ocean
Fig. 5.21 Mean annual precipitation (mm) of Korea
5.12.5 Korea’s Climatic Zones (After McCune, 1941) McCune’s ten climatic divisions, shown in Fig. 5.20, are characterized by the following: 1. An isolated mountainous region with sub-freezing temperatures during winter which may last for 5 months, and a short warm, moist summer; 2. The northeastern coastal belt has a cold winter with 3 months at sub-freezing temperatures and a mild summer;
5.12
Monsoon over Korea
153
3. The northern west – a forest-covered hilly land – has a very cold dry winter with mean January temperatures below –8◦ C, and a warm wet summer; 4. The central west Korea has a mean January air temperature between –6 and – 8◦ C; 5. The southern west Korea has a mean January air temperature between –3 and –6◦ C and an annual rainfall varying between 860 and 1370 mm; 6. The southern area which has a mean January air temperature between –3 and 0◦ C, has a mean annual rainfall which varies from 890 mm in the east to 1500 mm in the west; 7. The southeastern coastal belt has a mild winter with mean January air temperature between 0 and –3◦ C, and no dry months; 8. The southern coastal belt has a mild winter with mean January air temperature above 0◦ C and a long hot summer; rainfall exceeds 1500 mm at places, especially in the mountains; 9. Quelpart Island (Cheju Do) has a warm, moist marine climate with mean January air temperatures above 4◦ C and annual rainfall of about 1400 mm. 10. Dagelet Island (Ullung Do) has a mean January air temperature of 1◦ C, heavy winter rainfall and annual rainfall of 1500 mm. Cook (1964), in his survey of the climate of the Korean Peninsula, draws attention to the extraordinarily long duration of sub-freezing air temperatures over North Korea during winter. His work largely corroborates McCune’s conclusions regarding the climatic divisions of the peninsula and described above.
Chapter 6
Meteorology of the Maritime Continent (Region – III) (Comprising Philippines, Indonesia and Equatorial Western Pacific Ocean)
6.1 Introduction The region which we are about to study is not really a continent in its usual sense consisting mainly of land, but a continent in which the land surface may be looked upon as fragmented into thousands of islands, large and small, with water masses separating them so as to combine the effects of a continent and an ocean over a vast geographical area lying between the continents of Asia and Australia and between the Indian and the Pacific oceans. It covers a wide area of land and ocean extending longitudinally from about 100◦ E to the dateline or even beyond, and latitudinally from about 10◦ S to about 20◦ N (see Fig. 6.1). Along its northwestern borders lies a part of Southeast Asia. It has within its borders two great archipelagos, Philippines and Indonesia, and a large number of oceanic archipelagos, such as Micronesia, Polynesia, Melanesia, besides innumerable small islands or island groups in the heart of the Pacific Ocean, both north and south of the equator. The Philippines has more than 7000 islands, of which three (Luzon, Visayas and Mindanao) only are large. Indonesia also has within its borders thousands of islands, spread out along and about the equator between the Indian and the Pacific oceans. Here also there are only a few large islands (viz., Sumatra, Djawa, Borneo, Sulawesi, Moluccas, Timor and Irian Barat) and the rest are all tiny islands. It is estimated that more than two-thirds of the total area of the Maritime continent may be occupied by water masses and only one-third or even less by land. Apart from land-sea configuration, the topography of the Maritime continent plays an important role in determining the climate of the region. High mountain ranges, many of which are volcanic, are strewn all over the region and, depending upon their location and alignment in relation to the heat sources and sinks and prevailing winds, they play important roles in the meteorology of the continent, especially the distribution of fog, clouds, thunderstorms and rainfall. The Maritime continent occupies a centerstage in the monsoon circulations between Asia and Australia and between the Indian and the Pacific oceans. It experiences a monsoon type of climate, mostly warm and humid, practically throughout the year. Here, the NE monsoon alternates with the SW monsoon in a seasonal cycle, as the intertropical convergence zone moves north and south across the equator. Further, the tradewinds of the two hemispheres meet and K. Saha, Tropical Circulation Systems and Monsoons, C Springer-Verlag Berlin Heidelberg 2010 DOI 10.1007/978-3-642-03373-5_6,
155
156
6 Meteorology of the Maritime Continent
Fig. 6.1 Map of the Maritime continent showing its approximate boundaries and main islands
produce near-equatorial troughs and convergence zones including the South Pacific Convergence Zone (SPCZ). The Maritime continent acts as the principal heat source for the tropical circulation, involving the monsoon, Hadley and Walker-type circulations. It appears to play a crucial role in El Nino years when large-scale shift occurs in the distributions of equatorial heat sources and sinks and associated global climate systems. It is also the region of development of tropical waves into summertime typhoons of the NW Pacific and tropical cyclones of the SW Pacific. Because of the importance of the Maritime continent as a diabatic heat source for regional and global circulations, the meteorology of this continent has been studied by several workers ever since the colonial days. Historically, climatic studies in the Maritime continent started with two of its largest archipelagos, viz., the Philippines (Flores and Balagot, 1969) and Indonesia (Sukanto, 1969). A brief survey of some of the studies is presented in the sections that follow.
6.2 Climate of the Maritime Continent The important climatic elements for this continent are: pressure, temperature, humidity and rainfall, besides its wind systems.
6.2.1 Pressure The annual variation of surface pressure is very small over the Maritime continent, but not quite so over the continents which lie to its north and south. During northern winter (December–February) when maximum insolation is in the southern hemisphere, a ‘heat low’ develops over the continent of Australia and an extended trough of low pressure over the SW Pacific in the region of New Guinea,
6.2
Climate of the Maritime Continent
157
while a series of intense high pressure cells develop over the subtropical belt of the continent of Asia. These pressure distributions bring about a pressure gradient across the Maritime continent from northern to southern hemisphere. During northern summer (June–October), the pressure field is reversed with development of a series of deep ‘heat lows’ over Asia and an intense high pressure cell over Australia. The pressure gradient is then from southern to the northern hemisphere. The impact of the above-mentioned large-scale seasonal changes of pressure, however, appears to be limited to only the northern and southern parts of the Maritime continent. Near the equator, the pressure field is rather flat with only a weak gradient from the winter to the summer hemisphere.
6.2.2 Temperature The atmosphere over the Maritime continent is generally warm throughout the year. However, some seasonal variations do occur, mostly over the islands which lie in the extreme northern and southern latitudes, which come under the direct influence of the large-amplitude oscillations of temperature from one season to the other. An example of this influence is given in Table 6.1 which presents the average monthly temperatures during bi-monthly periods of the year in the main islands of the Philippines, viz., Luzon, Visayas and Mindanao. It may be seen from Table 6.1 that the seasonal variation of temperature in Luzon Island which is closer to East Asia mainland is much larger than that of Mindanao which is farther away. Similar differences in seasonal variation of temperature may also be seen between Indonesian Islands which are closer to the landmass of Australia than those located farther away. The seasonal variation of temperature in the Indonesian island of Sumatra is found to be much larger than that of any of the other islands of Indonesia, except, perhaps, Irian Barat for which no data are available. But it is well-known that during northern winter, the island of New Guinea as a whole being near the equator registers very high temperatures, similar to those over Northern Australia. Table 6.1 Average monthly temperatures (◦ C) during bi-monthly periods for 42 stations in the Philippines, averaged over the islands of Luzon, Visayas and Mindanao Number of stations averaged in each island is given in bracket Month Island
J-F
M-A
M-J
J-A
S-O
N-D
Luzon (19) Visayas (14) Mindanao (9)
25.1 26.3 26.3
27.0 27.5 27.2
28.2 28.1 27.4
27.5 27.5 27.0
27.1 27.4 27.0
25.7 26.8 26.7
158
6 Meteorology of the Maritime Continent
6.2.3 Relative Humidity and Cloudiness A high level of relative humidity, varying from 75 to 85%, prevails over the Maritime continent throughout the year. The seasonal variations are small. Mean monthly cloudiness over the continent is also moderate to high, varying from 4 to 6 oktas, with little seasonal variation.
6.2.4 Rainfall Rainfall is by far the most important climatic element of the Maritime continent. Its space-time distribution depends upon several climatic factors which we shall discuss in the next section. Tables 6.2 and 6.3 show the average monthly rainfall during bi-monthly periods in the main islands of the Philippines and Indonesia respectively. Table 6.2 Distribution of average monthly rainfall (mm) during bi-monthly periods in the Philippine islands (number of stations averaged for each island is indicated in bracket). Yearly total is given at the end Months
Luzon (20) Visayas (14) Mindanao (10)
J–F
M–A
M–J
J–A
S–O
N–D
Year
107 149 167
89 95 140
214 189 202
384 240 215
315 257 221
252 266 230
2724 2392 2350
Table 6.3 Average monthly rainfall (mm) during bi-monthly periods at selected island/stations in Indonesia, 1931–1960 Months Island/stations Sumatra/ Medan Padang Palembang Djawa/ Djakarta Surabaja Borneo/ Balikpapan Band-jarmasin Sulawesi/ Menado Timor/ Kupang Moluku/ Ambon
J–F
M–A
M–J
J–A
S–O
N–D
Year
134 306 260
140 382 297
173 315 142
190 300 94
232 516 144
217 563 304
2174 4764 2480
388 303
171 250
105 104
56 30
85 27
175 203
1760 1837
178 283
234 264
242 150
252 108
174 135
218 258
2597 2393
396
270
255
208
211
336
3352
327
176
27
3
5
110
1297
128
193
541
537
210
172
3457
6.3
Factors Affecting the Climate of the Maritime Continent
159
From the above tables, it is found that many of the island stations in Indonesia have more rainfall during southern summer than winter. Notable among these stations are Palembang in Sumatra, Djakarta and Surabaja in Djawa, Bandjarmasin in Borneo, Menado in Sulawesi and Kupang in Timor. Exceptions are Medan and Padang in Sumatra, Balikpapan in Borneo and Ambon in Maluku which experience more rain during southern winter than summer. An interesting aspect of rainfall at Padang which is located on the western coast of Sumatra facing the Southeastern Indian Ocean and at Palembang located on the eastern coast is that both experience two peaks in rainfall, once during northward movement of the monsoon and the other during the southward movement. According to Table 6.3, Padang in Sumatra appears to have the highest recorded annual rainfall of 4764 mm at a station in Indonesia. But, of course, no record exists for rainfall in Irian Barat where rainfall may be much higher, since its mountains are affected by the Southwest Pacific Convergence Zone (SPCZ) almost throughout the year. In the Philippines (Table 6.2), more rain falls during northern summer than during winter in all the main islands.
6.3 Factors Affecting the Climate of the Maritime Continent Several factors appear to affect the climate of the Maritime continent. Important among these are: 1. 2. 3. 4. 5.
Geographical location and topography; Ocean currents; Equatorial trough, ITCZ and Monsoons; SPCZ; Depressions and cyclones/typhoons.
6.3.1 Geographical Location and Topography Geographically, the Maritime continent is located in a more or less NW-SE orientation between the two large continents of Asia and Australia. However, being the larger continent, Asia appears to have somewhat greater influence on the climate of the Maritime continent than Australia during both winter and summer. It is also influenced by conditions in the world’s two largest oceans, the Pacific Ocean to the north and east and the Indian Ocean to the west and south. Almost every large island of the Maritime continent has mountain systems, many of which are volcanic. Through their thermal and mechanical effects, they not only control the direction of airflow but also influence the distribution of rainfall.
160
6 Meteorology of the Maritime Continent
6.3.2 Ocean Currents For a fuller understanding of the climate of the Maritime continent, it is necessary to know the impact of ocean currents on the circulation of the atmosphere over the region. According to oceanographers (e.g., Sverdrup, 1943), the warm North Equatorial ocean current driven by the northeast tradewinds on arriving at the Philippine coastal basin divides itself into two parts: one moving northwestward along the east coasts of Visayas and Luzon islands and turn into the well-known Kuroshio current and the other turning southward around the coasts of Visayas and Mindanao and then eastward to form the eastward-flowing warm Equatorial countercurrent. On the western coasts of these islands also, there is a similar division of the ocean currents driven by the monsoon winds, one turning northward along the coast of Luzon and the other turning south or southeastward along the coasts of Visayas and Mindanao islands. The above-mentioned ocean currents keep the average ocean surface temperature around the Philippine islands relatively high at 27.3◦ C, as compared to the average m.s.l. air temperature of 26.9◦ C at landstations in the area (Selga et al., 1931). Observations also reveal that the variability of water temperature in the area decreases from north to south. For example, the annual range of variation of water temperature in the Bashi and Balintang channels north of Luzon is 7◦ C, but it drops to 5◦ C in the South China Sea and about 4◦ C in all the seas south of Luzon. The warm sea surface temperature around the Philippine islands maintains high convective activity over the region. It is also believed to be responsible for frequent occurrence of instability thunderstorms, cyclogenesis and high frequency of development of typhoons over the area, and in general increase summer monsoon rainfall over the Maritime continent.
6.3.3 Equatorial Trough, the ITCZ and Monsoons 6.3.3.1 The Equatorial Trough and Its Movement It is well-known that the equatorial trough of low pressure which acts as a heat source undergoes a seasonal movement following the seasonal movement of the Sun. The seasonal movement of the sun creates differential heating and an isallobaric gradient in the meridional direction. The equatorial trough of low pressure which is usually located over the ocean is forced by the background isallobaric gradient to move in the direction of the gradient. Now, the isallobaric gradient may occur between two neighboring parts of land, or ocean, or between land and ocean. In the Maritime continent, because of the distribution of land and ocean, the only viable route for movement of the Equatorial trough of low pressure from one
6.3
Factors Affecting the Climate of the Maritime Continent
161
hemisphere to the other is across the landmasses of Southeast Asia, i.e., via the western Indonesian islands of Djawa and Sumatra, Malayasia, Thailand, Vietnam and mainland China. 6.3.3.2 ITCZ Since the tradewinds of the two hemispheres converge into the circulation around the equatorial trough of low pressure, forming the ITCZ on its equatorial side and the TCZ in the poleward side (see Fig. 1.4), the ITCZ and its associated monsoon current will follow the same land route in moving from one hemisphere to the other as the equatorial trough. This is well borne out by Fig. 6.2, due to Flores and Balagot (1969), showing the mean monthly locations of the ITCZ during its northward and the southward movements across the equator. It may be noted that the ITCZ reaches its southernmost position near the north coast of Australia sometime in February and northernmost position in the East China Sea in July. Starting from the south, its movement is slow till about April when it lies evenly poised along the equator. From May onward, however, a dramatic change appears to occur in its location. Its western part (west of about 120◦ E) comes under the influence of a series of ‘heat lows’ over southeast Asia and moves rapidly northward through Thailand and Vietnam to the southern part of China, while the eastern
Fig. 6.2 Mean monthly locations of the ITCZ over the Maritime continent and neighborhood during its (a) northward advance between January and July, and (b) southward retreat between July and December (after Flores and Balagot, 1969)
162
6 Meteorology of the Maritime Continent
part moves slowly over the ocean toward the Philippine Islands. The differential movement continues till July when the ITCZ on the two fronts reaches its northernmost location. The return journey appears to be traversed very much by the same route, though somewhat tardily. It is back to its equatorial location in November. From there, its western part moves rapidly across the Indonesian islands to converge into the circulation around the ‘heat low’ over the Australian continent during January-February.
6.3.3.3 Monsoons The Maritime continent experiences monsoon wind systems almost throughout the year, with wind directions reversing between winter and summer, as shown in Fig. 6.3(a) and (b) respectively. (a) Northern Winter Monsoon: During northern winter, N/NE-ly winds diverging from the strong anticyclonic circulation around the Siberian high pressure area, converge into the northeasterly tradewinds which blow over the tropical North Pacific Ocean along a line oriented in a NE-SW direction in the vicinity of the Philippine Islands. However, over most parts of the western North Pacific Ocean, diverging winds from the Asiatic anticyclone appear to cross the equator and blow as the NW-ly tradewinds and converge into the heat low circulations over Northwestern Australia and New Guinea area. On account of their long travel over warm equatorial oceans, both the aircurrents pick up enough heat and moisture from the underlying ocean surface before they arrive at the convergence zones. At the convergence zones they produce penetrative convection, heavy clouding and precipitation, which are easily detectable in satellite visible and infrared cloud imagery. Islands of the Maritime Continent with large mountain ranges facing these winds experience heavy rainfall on their windward sides. (b) Northern Summer Monsoon: From about late-April onward, a series of ‘heat lows’ develop over the landmasses of Southeast Asia, from Malayasia north-northeastward via Indo-China peninsula to Southern China. A heat low also develops over the Philippine islands area. The aircurrents which converge into these heat low circulations are basically from the equatorial Indian Ocean, both north and south of the equator and the western Pacific Ocean. They travel over the Maritime Continent to converge into these heat low circulations and produce intense rainfall along the ITCZ and light to moderate rainfall along the TCZ.
6.3.3.4 SPCZ The Southwest Pacific Convergence Zone (SPCZ) is a narrow but prominent convergence zone between the SE tradewinds of the southern hemisphere and the heat low circulation over the New Guinea area, extending eastsoutheastward across the Coral Sea in equatorial SW Pacific Ocean region (see Fig. 6.3). It is associated with a warm trough of low pressure over the oceanic area. It appears as a prominent
6.3
Factors Affecting the Climate of the Maritime Continent
163
Fig. 6.3 Low level tradewind circulations and convergence zones over the Maritime continent and neighborhood during northern (a) winter, and (b) summer
cloud band in satellite cloud imagery almost throughout the year and constitutes a quasi-permanent heat source for the tropical circulation. However, its location and intensity vary with season, being in a more southerly location and more intense and active during southern summer. As of to-day, no completely satisfactory explanation has been offered regarding the origin and movement of this convergence zone. Kiladis et al. (1989) who carried out a GCM experiment on the origin of the SPCZ using the ECWMF T21 model found that the removal of the Australian or the South American continent or both had no effect on the location and orientation of the
164
6 Meteorology of the Maritime Continent
SPCZ, though it caused a loss of intensity of the convergence when the continental influences were removed. A plausible explanation offered by them regarding its orientation was that ‘the eastsoutheastward orientation of the SPCZ into the SW Pacific Ocean is primarily due to its interactions with the midlatitude circulation rather than to the presence of continental boundaries or oceanic upwelling in the eastern South Pacific’.
6.3.3.5 Tropical Cyclones/Typhoons Every summer, a number of tropical cyclones/typhoons develop in the ITCZ which lies close to the Philippines over the western North Pacific and affect the climate of the Maritime continent. Easterly waves moving into this area undergo explosive cyclogenesis/development for the reasons discussed in 2.10 and 2.11. After development, most of them move northwestward towards the east coast of China. But a large number of them recurve northeastward from near the island of Taiwan and affect Korea and the islands of Japan and North Pacific Ocean. Some cyclones developing over the South China Sea move westward across Vietnam and countries of Southeast Asia where they cause high winds and extremely heavy rain. Storm surges associated with most of these disturbances inundate large coastal belts and floods in many areas. The word ‘typhoon’ in Chinese means the ‘Big wind’. It is the typhoons that cause maximum loss of life and damage to property by their hurricane-speed winds, torrential flood-producing rain and coastal inundation by storm surge. Several accounts of the ferocity and dangers associated with these typhoons are available (e.g., Algue, 1904; Gherzi, 1930). For example, according to Tannehill (1927), the number of people who died at Haiphong in the typhoon that hit the East coast of China in 1881 was 300,000. Tropical cyclones usually weaken after landfall and few persist after moving beyond about 500 km from the coast. However, remnants of a few westwardpropagating ones may survive to emerge over the Bay of Bengal. An interesting study by Chin (1958) compiled the statistics of the number of tropical cyclones that passed through each square of 2.5◦ latitude–longitude over the northwest Pacific and the China seas during August of a 70-year period, 1884–1953, presented in Fig. 6.4. A monthly breakup of the average number of tropical cyclones within 5–30◦ N and 105–150◦ E during the same 70-year period is given in Fig. 6.5. Figure 6.5 which includes tropical cyclones of varying intensity, ranging from weak depressions to mature cyclones (or typhoons), shows that annually about 22 tropical cyclones form over the region of northwestern Pacific and the China seas. Of these, most occur between July and October. With improvement in observation, a more reliable statistics has been available in recent years. For example, during the period 1947–1964, the average number of cyclones detected per year over the same area was found to be 31 with a maximum of 45 in 1961 and a minimum of 20 in 1951.
6.4
The Maritime Continent – A Heat Source
165
Fig. 6.4 Number of tropical cyclones passing through each square of 2.5◦ latitude–longitude over the northwest Pacific and the China seas during August in a 70-year period, 1884–1953. Broken lines are axes of maxima (after Chin, 1958)
Fig. 6.5 Monthly distribution of the average number of tropical cyclones
6.4 The Maritime Continent – A Heat Source Attempts have been made to identify the Maritime continent as a heat source or sink by using one or more of the well-known characteristics associated with the ITCZ. Flores and Balagot (1969) used the low-level convergence of airflow and
166
6 Meteorology of the Maritime Continent
precipitation for the purpose. Upper level divergence derived from computed fields of velocity potential has been used by a number of workers (e.g. Ding and He, 1984; Krishnamurti and Surgi, 1987; Murakami, 1987a). Chen and Li (1981) used the energy balance equation to compute the distribution of heat sources and sinks over the Asia-Australia region. They used the expression, HS = R∞ − R0 − SC + LH where HS denotes heat source, R∞ is the net radiation at the top of the earth’s atmosphere, R0 is the radiation balance at the earth’s surface, SH is sensible heat and LH is latent heat. Their results for the month of July are shown in Fig. 6.6. Murakami (1983) used the 3-hourly infra-red radiation data from a GMS-1 geostationary satellite which was positioned directly over the Maritime continent (0, 140◦ E). The IR data were compiled in terms of equivalent blackbody temperature, TBB . They were further edited on every 1◦ latitude-longitude square mesh so that a histogram could be prepared to yield a mean value and a standard deviation, σ B , to take account of the spatial variability of the original TBB observation (about 5 km in resolution) within the mesh. He empirically proposed that a value of σ B larger than 5 K can be regarded as a convective area. To measure convective activity, he used an intensity index Ic , based on cloud-top temperature and temperatures at two standard atmospheric levels. This index enabled him to distinguish between the radiation emitted from tops of large convective clouds and that from low level clouds or cloud-free areas.
Fig. 6.6 Distribution of mean July heat sources (+) and sinks (–) over the Maritime Continent and adjoining continents and oceans (Chen and Li, 1981)
6.4
The Maritime Continent – A Heat Source
167
He defined the index, Ic , by Ic = (Tc − T400 ) × 10/(Ttrop − T400 ) where Tc denotes the mean cloud-top blackbody temperature, and T400 and Ttrop are atmospheric temperatures at 400 mb and tropopause level respectively. It is easy to see that the value of Ic is zero at 400 mb but increases directly with the height of the cloud top above the 400 mb surface. The horizontal distributions of monthly mean values of Ic for December 1978 and June 1979, as computed by Murakami (1983, 1984), are shown in Figs. 6.7 and 6.8 respectively. Murakami’s computed values of Ic for December 1978 (Fig. 6.7) reveals three prominent heat sources over the Maritime Continent, one in the near-equatorial region between longitudes 110 and 120◦ E, the second in the Irian Jaya area between 130 and 140◦ E, and the third over the Southwestern Pacific Ocean east of about 160◦ E. It appears that these heat sources may be related to the three regions of cross-equatorial monsoon flows into the southern hemisphere in December (see Fig. 7.8). The distribution of Murakami’s computed values of mean heat sources and sinks for June, 1979, appears to be in substantial agreement with that computed by Chen and Li for July. Further, Murakami’s results show the seasonal migration of heat sources and sinks between Australia and Eastern Asia.
Fig. 6.7 Mean monthly distributions of Ic over the Maritime continent and neighborhood for December 1978. Contours are drawn every 0.5 unit, starting from 1.0. Values greater than 2.0 are shaded. The hatched area roughly denotes the plateau above 3000 m (after Murakami, 1983, 1984)
168
6 Meteorology of the Maritime Continent
Fig. 6.8 Mean monthly distributions of Ic over the Maritime continent and neighborhood for June 1979. Contours are drawn every 0.5 unit, starting from 1.0. Values greater than 2.0 are shaded. The hatched area roughly denotes the plateau above 3000 m (after Murakami, 1983, 1984)
Murakami’s study reveals a prominent heat source in the region of the SW Pacific Convergence Zone (SPCZ) during both summer and winter. Another such heat source appears over the South Equatorial Trough (SET) area of the eastern Indian Ocean. The most intense heat source appears to be located over the northeastern part of the Bay of Bengal during the northern summer.
6.5 The Maritime Continent and the ENSO It is not always that the Maritime continent and the adjoining equatorial eastern Indian Ocean act as a heat source with a positive zonal SST anomaly, while the equatorial eastern Pacific Ocean or the equatorial western Indian Ocean acts as the heat sink. In an El Nino year, which comes usually after every 2–7 years, the roles are reversed and the Maritime continent acts as a heat sink with a negative SST anomaly, while the region of positive SST anomaly switches over to the eastern side of the Pacific ocean near the coast of South America and the western side of the equatorial Indian Ocean. The southern oscillation index over the Maritime continent and neighboring regions becomes highly negative in such a year. The change has disastrous effects on atmospheric circulation and associated rainfall over a vast region of the western Pacific and adjoining Australia, India and China where large deficiencies of rainfall lead to widespread droughts and famine conditions. Such reversals in temperature anomaly between the western and the eastern parts of the equatorial ocean are observed in other equatorial ocean basins as well.
6.5
The Maritime Continent and the ENSO
169
In an interesting study using the IIOE (1963–1965) data, Ramage (1968) found that in January 1963, the Maritime continent served as an active heat source as indicated by excessive rainfall and thunderstorm activity over the continent. The heat generated over the continent was efficiently transported by the Hadley circulation to cause an unusually strong subtropical jetstream over East Asia. By contrast, January 1964 was a drought year for the continent with deficient rainfall, inefficient poleward transport of heat and a weaker meridional circulation. According to him, ‘most winters over the western Pacific and southeast Asia fluctuate between situations typical of January 1963 and January 1964.’
Chapter 7
Monsoon over Australia (Region – IV)
7.1 Introduction – Location and Physical Features Situated in the southern hemisphere between latitudes about 10 and 43◦ S and longitudes 113 and 153◦ E and surrounded by oceans, the continent of Australia experiences its summer monsoon from about December to March and winter monsoon from May to October. A map showing the geographical location and physical features of the continent and surrounding areas is presented in Fig. 7.1. The oceans around Australia are: the Indian Ocean in the west, the Pacific Ocean in the east, the Great Australian Bight in the south, and a series of seas, such as the Timor, Arafura, and Coral Seas in the north. The Gulf of Carpentaria which lies between the Northern Territory and the York Peninsula also lies in the north. The land-sea configurations of the northern and the southern coasts of Australia maintain a quasi-stationary wave in the fields of temperature, pressure and circulation along these coasts, especially during Australian summer. Orography plays an important role in the climate of Australia. The Great Dividing Range which runs more or less parallel to its eastern coast divides the moderately cool oceanic climate on one side from the dry desert climate on the other. However, the southern part of the mountains including the Australian Blue Alps experiences moderate rainfall almost throughout the year. The mighty MurrayDarling River rises in these mountains and flows westward to make the southeastern part of the continent fertile and abundantly habitable. A significant impact on the continent’s weather and climate is made by synopticscale disturbances in the form of depressions and cyclones. They develop in the quasi-stationary waves when traveling E’ly or W’ly waves of the southern hemisphere interact with them. Most of them form over the warmer waters of the oceans around Australia. In the north, the oceanic areas which are most likely to breed these disturbances are the Timor and Arafura seas, the Gulf of Carpentaria and the Coral Sea. The southern parts of the continent are affected by the eastward-propagating subtropical/midlatitude baroclinic waves. The continent is also affected by ENSO events, though irregularly, once every 2–5 years.
K. Saha, Tropical Circulation Systems and Monsoons, C Springer-Verlag Berlin Heidelberg 2010 DOI 10.1007/978-3-642-03373-5_7,
171
172
7 Monsoon over Australia
Fig. 7.1 Geographical location and physical features of Australia Territorial boundaries are indicated by long-dashed lines, deserts by dots, mountains by hats, and depressions and cyclones by open circles
7.2 Early Studies The early studies of Australian summer monsoon date back to the sixties and early seventies of the last century (e.g., Troup, 1961; Berson and Troup, 1961; Gentilli, 1971). But the same were conducted with limited surface and upper-air data. The data situation improved during the FGGE Northern Hemisphere Winter Monsoon Experiment (WMONEX), 1978–1979, following which there was a spurt in research activity with studies undertaken on such diverse topics as onset and structure of the summer monsoon, divergent circulations, active and break monsoons, intraseasonal oscillations, tropical midlatitude interactions, depressions and cyclones, effect of ENSO on Australian rainfall and other related weather phenomena. The results of some of these studies (e.g., Sumi and Murakami, 1981; Murakami and Sumi, 1982; Nicholls et al., 1982, 1984a,b; Davidson et al., 1983, 1984; McBride, 1983; McBride and Nicholls, 1983; Love and Garden, 1984; Pittock, 1984; Love, 1985a,b; Holland and Nicholls, 1985) are available in an excellent review by McBride (1987). Holland (1984a,b,c) made a special study of the climatology and structure of the tropical cyclones which form in the Australian/SW Pacific region. Westward-propagating tropical disturbances often
7.3
Climate of Australia and Surrounding Oceans
173
recurve into higher latitudes where they come under the influence of eastwardpropagating midlatitude disturbances of the southern hemisphere. All the above-mentioned studies contributed significantly to our knowledge and understanding of the Australian summer monsoon. A more recent review of Australian summer and winter monsoons has been provided by Saha and Saha (2001a).
7.3 Climate of Australia and Surrounding Oceans A major portion of the continent lies over the subtropical belt, so it is mostly the northern part (north of about 20◦ S) that experiences the impact of a monsoonal-type of climate. The rest of the continent, barring the coastal regions, located over the subtropical belt, experiences generally dry and desert-like climate almost throughout the year. In fact, some of the world’s extensive deserts, viz., the West Australian desert, the Gibson Desert, and the Simpson Desert, are all located over this continent. However, a narrow belt along the southern coast of Australia, especially the southwestern portion of Western Australia and the southeastern States of Victoria, New South Wales and Southern Queensland which jut out into the southern oceans enjoy mild climate. These latter areas are affected by eastward-propagating midlatitude baroclinic wave disturbances of the southern hemisphere, more frequently during winter than summer, and experience occasional spells of cool, rainy weather. In this respect, the southeastern states which extend to higher latitudes feel the impact of these disturbances to greater extent and enjoy much milder climates with heavier rainfall. The island of Tasmania which lies further poleward has cool, rainy climate almost throughout the year.
7.3.1 Ocean Surface Temperature (SST, C) Figure 7.2 shows the mean ocean surface temperature around Australia during (a) February and (b) August (Courtesy: NCEP Reanalysis). In February (Fig. 7.2a), two prominent areas of warm SST (≥28◦ C, shaded) stand out; one rather a narrow zone over the equatorial eastern Indian ocean extending from near equator southeastward to a wider area near the northwestern coast of Australia, and the other a very extensive area of equatorial western Pacific Ocean near New Guinea. These warm areas extend poleward to about 20◦ S. The warm zone near New Guinea area covers a wide area across the Coral Sea and extends eastsoutheastward over the SW Pacific Ocean. The SST off the coast of Western Australia is much lower than that off the coasts of Queensland and New South Wales at the same latitude. In August, the whole thermal pattern appears to have shifted northward by a few degrees of latitude, so that most of the warm areas now lie north of the continent. The large warm SST area near New Guinea has also shifted northward.
174
7 Monsoon over Australia
Fig. 7.2 Climatological (1971–2000) SST (C) around Australia: (a) February and (b) August
7.3.2 Air Temperatures Mean air temperatures at two pressure surfaces, viz., 925 and 300 hPa, over Australia and surrounding oceans during February and August, obtained from NCEP Reanalysis, are shown in Fig. 7.3(a,b) respectively. In February, the land surface and the lower troposphere over the continent is very warm with a pronounced temperature maximum centered over Western Australia. Two warm ridges may be seen, one over the southern part of Western Australia and the other over Queensland, New South Wales and Victoria areas. Temperatures drop rather slowly towards the equator, but steeply poleward. However, in the nearequatorial Indonesian region, there appears to be a gradual increase of temperature from west to east, resulting in a well-marked warm area over New Guinea and
7.3
Climate of Australia and Surrounding Oceans
175
Fig. 7.3 Mean air temperature (◦ C) over Australia and surrounding regions at 925 and 300 hPa (a) February and (b) August (Courtesy: NCEP/NCAR Reanalysis Project)
adjoining Coral Sea area which appears to extend further eastward to the dateline or even beyond. In the upper troposphere (300 hPa) during February, a belt of warm area is located over the northern part of the continent with a temperature maximum along about 15◦ S. The temperature gradient towards the equator is almost negligible, but that to higher latitudes appears to be quite steep. The east–west gradient of temperature in the lower troposphere almost disappears in the upper troposphere. In August, which represents the Austral winter season, the whole thermal field in the lower and the upper troposphere appears to have moved northward and temperatures over the continent as well as the adjoining oceans to west and east are much lower. Further, there appears to be no real temperature maximum now over the continent except a weak warm area over the ocean close to its northern coast.
176
7 Monsoon over Australia
7.3.3 Atmospheric Pressure (Isobaric Height) Consistent with the distribution of temperature shown in Fig. 7.3(a,b), the distribution of geopotential height (gpm) at 925 hPa during February and August is shown in Fig. 7.4. According to Fig. 7.4(a), a deep low pressure trough oriented in a WSW-ENE direction is located at 925 hPa over northwestern Australia. It is the summer
Fig. 7.4 Isobaric height (m) fields at 925 hPa over Australia and surrounding areas during (a) February and (b) August (Courtesy: NCEP Reanalysis)
7.3
Climate of Australia and Surrounding Oceans
177
‘heat low’ (H.L.) trough over Australia. Other well-marked troughs of low pressure at 925 hPa include: (i) The equatorial trough over the eastern Indian Ocean to the west of the Australian heat low; (ii) An approximately N–S oriented trough over Western Australia; a poleward extension of the ‘heat low’ trough; (iii) A north-south oriented trough of low pressure over eastern Australia. This trough is often called the ‘Cloncurry trough’, since it is associated with a low pressure over the Cloncurry area of Queensland; Two additional troughs of low pressure appear over the oceans. These are: (iv) A prominent trough of low pressure over the Southwest Pacific Ocean, extending from New Guinea area eastsoutheastward across the Coral Sea region; (v) A low pressure trough off the east coast of Australia. In February, the ridge of the subtropical high pressure appears to lie along about 35◦ S, with a high pressure cell each over the Great Australian Bight and the Tasman Sea. A ridge of the Tasman Sea high pressure cell appears to extend equatorward along the Great Dividing Range. In February, a strong pressure gradient exists at 925 hPa between the heat low over Australia and the cold high pressure areas over the oceans to west, east and south. Pressure generally increases northward with only a small gradient over the equatorial region. In the equatorial region, pressure appears to decrease generally from west to east. In august, with change of season, the ridge of the subtropical high pressure over Australia appears to have moved equatorward by about 7–10◦ of latitude and runs along about 28◦ S, with much steeper pressure gradient towards the pole than towards the equator. A pressure minimum appears over the equatorial region.
7.3.4 Wind and Circulation Figure 7.5 shows the mean wind field and circulation over Australia and neighborhood at 925 and 300 hPa: (a) February, (b) August. The February wind field at 925 hPa shows the following features: (i) A well-defined cyclonic circulation around the ‘heat low’ over northwestern Australia; (ii) A broad band of strong cross-equatorial flow from northern to southern hemisphere over the longitudes of Indonesia; NE-ly tradewinds, after crossing the equator, turn anti-clockwise and blow as W/NW-ly tradewinds. These are the deflected monsoon winds; (iii) The deflected northwesterlies converge into the cyclonic circulations around the heat lows over Australia and New Guinea area, producing well-defined
178
7 Monsoon over Australia
Fig. 7.5 Mean wind field and circulation over Australia and surrounding regions at 925 and 300 hPa (a) February and (b) August (Courtesy: NCEP/NCAR Reanalysis Project) but for wind and circulation
ITCZ and TCZ. It is the TCZ which extends eastsoutheastward from New Guinea area across the Coral Sea region which has come to be known as the SW Pacific Convergence Zone (SPCZ). At 300 hPa in February, the windfield shows an anticyclonic circulation over northern Australia with its ridge running along about 18◦ S. This means that the low level W/NW-ly trade winds are overlain by upper-air easterlies over the region to the north of Australia. Poleward of the ridge, the flow is generally westerly. The August wind field at 925 hPa shows the following circulation features (i) A cross-equatorial flow from the Australian region to the northern hemisphere (Note the reversal of the flow direction from SE to SW);
7.3
Climate of Australia and Surrounding Oceans
179
(ii) A general northward shift of the axis of the subtropical anticyclone over the Great Australian Bight from about 35◦ S to about 28◦ S over Australia; (iii) Strong E/SE-ly tradewinds sweeping across most of northern Australia, with westerlies to the south of the subtropical ridge. At 300 hPa in August, the axis of the anticyclonic circulation appears to have shifted northward to the extreme northern part of Australia with strong easterlies to its north and westerlies to the south. The westerlies attain high jet speeds during southern winter. Two aspects of the atmospheric circulation over the Australian region are noteworthy. These are: (1) Cross-equatorial airflow of the lower troposphere; and (2) The upper-tropospheric subtropical jetstream.
7.3.4.1 Low-Level Cross-Equatorial Airflow Figure 7.6 presents the equatorial distribution of monthly mean v, the meridional component of the wind, between 105 and 150◦ E: (a) February and (b) August. The February distribution (Fig. 7.6a) reveals the existence of three longitudinal segments of northerlies, one extending from the eastern part of the Bay of Bengal to about 105◦ E, the other from about 110◦ E to about 135◦ E and a third from about 140◦ E further eastward. Of these, the first two are strong and deep, extending from surface to almost 300 hPa, while the third has a layer of southerlies between about 700 and 500 hPa. All the segments have southerlies above about 300 hPa. The August distribution (Fig. 7.6b) shows that barring a few exceptions, crossequatorial flows occur over almost the same longitudinal segments as during February, but in the reverse direction, that is, from southern to the northern hemisphere. Here also, more than one layer of cross-equatorial flow are involved. Strong southerlies blow below about 700 hPa over all the segments, but the layer is overlain by a shallow layer of northerlies around 600 hPa. But a layer of southerlies reappears in midtroposphere, with a deep layer of northerlies above. 7.3.4.2 Jetstreams Loewe and Radok (1950) computed the meridional profile of the zonal component of the wind from the distribution of temperature by assuming geostrophic approximation. They found that in both the seasons, the W’ly jetstream occurred at a height of about 12–14 km, but the speed of the jetstreams during winter (July) was much higher than that during the summer (January). Their computed results appear to be in substantial agreement with actual observations obtained later from various sources, such as rawins, etc. The accuracy of their computations is also confirmed by IIOE datasets presented by Ramage and Raman (1972) and later observations.
180
7 Monsoon over Australia
Fig. 7.6 Vertical profile of v, the meridional component of the wind (m s–1 ), along the equator between 105 and 150◦ E: (a) February and (b) August (Saha and Saha, 2000)
7.3
Climate of Australia and Surrounding Oceans
181
Fig. 7.7 Jetstreams over Australia during (a) January, (b) July (after Loewe and Radok, 1950)
The subtropical jetstreams over Australia, first computed by Loewe and Radok (1950), are presented in Fig. 7.7: (a) for January and (b) for July, in which the isolines show the latitudinal distributions of temperature (◦ C) (dashed) and zonal geostrophic windspeed (m s–1 ) (continuous lines) at various pressure surfaces from 1000 to 50 mb over the southern latitudes from the equator to 50◦ S. The distributions
182
7 Monsoon over Australia
show that the winter jetstream is not only stronger but also occurs much closer to the equator than the location of the summer jetstream. Such seasonal movements of the subtropical jetstreams are also observed in the northern hemisphere and the winter jetstreams which occur closer to the equator are much stronger than the summer jetstream.
7.4 Monsoon over Australia 7.4.1 Onset of Monsoon According to Troup (1961) who studied the onset of summer monsoon at Darwin, a change in the gradient-level (900 hPa) wind direction to northwesterly marks the beginning of the monsoon season at the station. He found that the change is followed by a spurt in convective activity and precipitation. He also found that the intensity of monsoon rainfall was highly correlated with the intensity of the cross-equatorial flow from the northern hemisphere. The thermal wind over the region in Australian summer being easterly, the northwesterly flow at low level changed over to easterlies at some level in midtroposphere. While the above view of onset of summer monsoon over Australia appears to be widely accepted, Davidson et al. (1983) put forward an alternative view that the onset of summer monsoon over Australia is strongly influenced by synoptic events in the subtropics of the southern hemisphere. They observed that a dramatic increase in the intensity of convective activity and precipitation at Darwin occurred during the movement of eastward-propagating large-amplitude midlatitude baroclinic wave disturbances across the Australian region. They defined monsoon onset by the firsttime appearance of the gradient-level westerly wind at Darwin, following Troup (1961). The actual date of onset was taken to be the date at which there was largescale increase in convective activity, as long as it occurred within 5 days of the appearance of the gradient-level northwesterly wind. Their conclusion was based on the results of a study of monsoon onset during the 6 years 1971, 1973, and 1976–1979 in which they found that in each of the years examined the enhancement of tropical convection could be attributed to an interaction between a largeamplitude subtropical/midlatitude baroclinic wave and the monsoon trough and its eastward movement along latitude about 10◦ S towards Darwin. In the present text, we are led by Fig. 7.6(a) to put forward a hypothesis that monsoon advances and sets in over Australia in the same manner and retaining the same wave structure as it does over other parts of the tropics, for example, the Indian Subcontinent, Eastern Asia, Africa, and South America. In the case of Australia, the advance of the monsoon wave is spearheaded by three cross-equatorial currents of strong northwesterlies which after flowing over the Maritime Continent converge into the heat lows of the Australian region. While the main heat low lies over the Australian continent, the low pressures over the equatorial eastern Indian Ocean, and the New Guinea area, are the other two heat
7.4
Monsoon over Australia
183
Fig. 7.8 Schematic showing the principal cross-equatorial aircurrents (thick continuous lines with arrow) involved in monsoon onset over Australia and other neighboring regions Streamlines (thin continuous lines with arrows) show directions of air motion around troughs (thick dashed lines) of low pressure. L denotes Low Pressure; H, High pressure
lows. The circulation features around these heat lows are shown by a schematic in Fig. 7.8. Figure 7.8 shows how the principal aircurrents, after crossing the equator near longitudes 105, 130 and 150◦ E first diverge and then converge into the circulations around the heat lows over the different regions. Note that two principal aircurrents converge into the circulation around the heat low over the Australian mainland. It is the arrival of the principal aircurrents that appears to signal the onset of the summer monsoon with its convective activity and rainfall over a region.
7.4.2 Co-existence of Monsoon and Hadley Circulations – Interhemispheric Movement The monsoon circulation as a perturbation in the tradewind circulation over the Australian region and its co-existence with the Hadley circulations of the two hemispheres stands out in a meridional-vertical section through the heat low over Australia, shown schematically in Fig. 7.9. The resultant streamlines, shown in Fig. 7.9, were arrived at by using computed values of v (the meridional component of the wind) and ω (the vertical
184
7 Monsoon over Australia
Fig. 7.9 Resultant streamlines (constructed from computed v and ω values) along 135◦ E, showing the linkage between the circulations of the two hemispheres and the Monsoon circulation over Australia during Australian summer. The arrow shows the direction of airflow. The location of the monsoon trough zone is indicated by a thick dashed line sloping equatorward with height (after Saha and Saha, 2000, 2001a)
p-velocity) from an 18 year (1979–1996) mean NCEP/NCAR reanalysis, after suitably scaling the ω-values, along 135◦ E, a meridian passing through the heat low over Australia. The results shown in Fig. 7.9 clearly reveal the interhemispheric structure of the monsoon circulation in co-existence with the Hadley circulations of the two hemispheres. During Australian summer, the cool, humid monsoon winds diverging from the subtropical high pressure of the northern hemisphere cross the equator and converge into the circulation around the heat low over Australia. The field is reversed during Australian winter when similar winds diverging from the subtropical high pressure belt of the southern hemisphere cross the equator and converge into the circulations around the heat lows over Eastern Asia and the Indian Subcontinent. The direction and magnitude of cross-equatorial fluxes of air in the two seasons are presented in Table 7.1. For comparison, similar information available in respect of the Western Indian Ocean region and the equator as a whole is also included in the table.
7.4
Monsoon over Australia
185
Table 7.1 Magnitudes of estimated and computed cross-equatorial fluxes of air (Unit: 1012 metric tons (day)–1 ) (Plus sign indicates S’ly, minus N’ly) Equatorial sector
Tropospheric layer
January/ February
July/ August
(i) Australian section (105–150◦ E) (Saha and Saha, 2000) (ii) Western Indian Ocean (42–75◦ E) (Findlater, 1969a, b) (Saha, 1970) (iii) Estimated total across the whole equator (5◦ N–5◦ S) (Rao, 1964)
Lower (sfc-300 hPa) Upper (300–50 hPa)
–3.14 +1.93
+2.68 –4.00
Lower (sfc-400 hPa) Lower (sfc-400 hPa) Lower (sfc-500 hPa)
–2.28
+7.68 +5.03 +16.20
–18.49
Table 7.1 highlights the magnitudes and direction of the seasonal movements of airmasses between the hemispheres during both Australian summer (January/February) and winter (July/August). According to Table 7.1, during February, the total cross-equatorial flux amounts to –3.14 in the lower troposphere and 2.68 in the upper troposphere. In August, the figures change to 1.93 and –4.00 respectively. It is evident from Fig. 7.9 and Table 7.1 that during Austral summer (February) the lower-tropospheric monsoon trough where the circulations from the two hemispheres converge slopes equatorward with height up to about 500 hPa and that the Hadley circulation of the winter hemisphere makes considerable inroads into the summer hemisphere both in the lower and the upper troposphere. It is mostly around this sloping line between the equator and latitude 20◦ S that low-level convergence supported by upper-air divergence leads to penetrative convection and precipitation over Australia. The involvement of the Hadley circulations of the two hemispheres in the formation of monsoon circulation over Australia stands out in Fig. 7.9. It connects the monsoon circulation over Australia with that over Asia.
7.4.3 Summer Monsoon Rainfall over Australia The distributions of mean February rainfall (mm day–1 ) and observed outgoing longwave radiation (OLR) over Australia and surrounding oceans, as obtained from Reanalysis, are presented in Fig. 7.10(a,b) respectively. They show concentration of heavy rainfall exceeding 6 mm day–1 along the northern and northeastern coasts of the continent as well as over the adjoining oceanic areas. The area of heavy rainfall extends northward to about 5◦ N. There are several pockets of heavier rainfall exceeding 10 mm day–1 along the equatorial zone of Indonesia as well as the oceanic area to the north and northeast of Australia, especially the New Guinea area. The rainfall rates are well supported by OLR values. In general, low OLR values (≤220 Wm–2 ) indicate penetrative convection and high rainfall rates. Three areas
186
7 Monsoon over Australia
Fig. 7.10 Distribution of February (a) Mean rainfall (mm day–1 ) and (b) Outgoing Longwave radiation (OLR) (Wm–2 ) over Australia and surrounding areas (from NCEP/NCAR Reanalysis)
7.5
Annual Rainfall of Australia and Its Seasonal Variability
187
of low OLR with values of 190 Wm–2 or less appear over the equatorial region of eastern Indian Ocean, Southern Borneo, and the New Guinea area. By contrast, very high values of OLR exceeding 260 Wm–2 prevail over the subtropical belt especially Southern and Western Australia. An extensive area of the eastern Indian Ocean off the coast of Western Australia is an extremely dry area within the subtropical belt with high OLR values.
7.5 Annual Rainfall of Australia and Its Seasonal Variability 7.5.1 Annual Rainfall The inadequacy of water resources in Australia which is a major problem of the continent is well reflected by the distribution of its annual rainfall, shown in Fig. 7.11.
Fig. 7.11 Mean annual rainfall (mm) of Australia during period, 1911–1940 (after Bureau of Meteorology, 1962)
188
7 Monsoon over Australia
According to an estimate by Gentilli (1971), the annual rainfall of Australia leaves 37% of the land with less than 250 mm, 57% with less than 375 mm, and 68% with less than 500 mm of rain (cumulative percentages). Statewise, on the average, 83% of Southern Australia, 58% of Western Australia, 25% of Northern Territory, 20% of New South Wales, and 13% of Queensland receive less than 250 mm of rain in the year. Only 0.5% of South Australia, 5.5% of Western Australia, 16% of New South Wales, 17% of Northern Territory, 23% of Queensland, and 27% of Victoria receive more than 750 mm of rain in the year. Only in Tasmania, rainfall appears to be plentiful with more than half the island receiving more than 1000 mm in the year. The driest area in Australia is located in Southern Australia, around Lake Eyre, where the average rainfall is less than even 125 mm in the year. On account of extremely low rainfall, vast tracts of the continent suffer from frequent droughts
7.5.2 Seasonal Variability A study by Andrews (1932, 1933) reveals the seasonal variation, i.e., the percentage of the total annual rain that falls in a particular season. His findings for the four seasons, viz, Summer (A), Autumn (B), Winter (C), and Spring (D), are shown in Fig. 7.12.
Fig. 7.12 Seasonal concentration (%) of mean annual rainfall over Australia: Summer (a), Autumn (b), Winter (c), and Spring (d) (after Andrews, 1932, 1933)
7.7
Tropical Disturbances in the Australian Region – Depressions and Cyclones
189
7.6 Variability of Australian Rainfall with ENSO A possible relationship between Australian rainfall and the Southern Oscillation has been the subject matter of much research in recent years (e.g., McBride and Nicholls, 1983; Allan, 1983). The studies conducted so far reveal that years of extreme values of the Southern Oscillation Index (SOI), coincide with those of widespread very high or very low values of rainfall over the Australian tropics. McBride (1987) quotes an example from two contrasting years of rainfall, namely, 1983 and 1974, when the January SOI was highly negative and positive respectively. In January 1983, an SOI of –29.8 corresponded to a distribution of widespread much below average rainfall over northern Australia, whereas in January 1974 with a value of +21.7 for SOI, rainfall was very much above average over a much wider area of tropical Australia. Here, the SOI which is usually given by the difference in sea level pressure between Tahiti and Darwin divided by the standard deviation of that quantity is a measure of the seesaw type oscillation in surface pressure between the equatorial eastern Pacific Ocean and the Western Pacific-Eastern Indian Oceans, as originally defined by Sir Gilbert Walker and called the Southern Oscillation. However, the relationship appears to exist in years of highly extreme values of SOI only. Over a period of many years in a row, the relationship appears to be weak.
7.7 Tropical Disturbances in the Australian Region – Depressions and Cyclones In Australia, interest in studies of atmospheric disturbances began quite early last century and has continuously grown since then, as demonstrated by several studies (e.g., Bureau of Meteorology, 1956, 1978; Keenan, 1981, 1982; McBride and Keenan, 1982; Holland, 1984a,b,c; Lajoie and Butterworth, 1984; Nicholls, 1984b), especially after the MONEX, 1978–1979. The contributions made by these studies have thrown light on several aspects of the formation and behaviour of monsoon depressions and tropical cyclones in the Australian region and adjoining the SW Pacific Ocean. However, notwithstanding great advances made, uncertainties remain in several areas relating to these disturbances, especially their development and movement that need further study. With the availability of satellite data and global data analysis, we have now much greater opportunity to observe and study these disturbances than ever before. Paterson and Bate (2001) who carried out a detailed study of tropical cyclones over the South Pacific and Southeast Indian Ocean during the cyclone season (November–April), 1999–2000, found that the number of tropical cyclones formed during the season to the west of 105◦ E, between 105 and 165◦ E, and to the east of 165◦ E was 10, 11 and 5 respectively against a climatological average of 12.7, 9.6 and 5.6 over the respective basins. Figure 7.13 shows the tracks of the cyclones that formed between longitudes 100 and 180◦ E during the season.
190
7 Monsoon over Australia
Fig. 7.13 Tracks of named tropical depressions and cyclones which formed during the summer season (December–April) of 1999–2000 between 100 and 180◦ E. Name of the disturbance is given at the location of its first detection. The arrow shows the direction of its movement (Paterson and Bate, 2001)
Paterson and Bate furnished particulars of these cyclones stating the date each was first identified as a low, the date it turned into a cyclone, the date it reached maximum intensity and the date it ended its tropical cyclone phase. Those interested in these details may look up their original paper. As shown in Fig. 7.13, after formation, most of the cyclonic disturbances tend to move with the dominant airstream in which they are embedded, and gradually recurve towards the south as they move towards the belt of the midlatitude westerlies. But a small percentage of them do not recurve but keep moving in their original direction till they move over a cold ocean or come under the influence of some other disturbances. McBride and Keenan (1982) who carried out a case-by-case study of tropical cyclone development over a period of 5 years found that in 84% of the cases examined, the precyclone cloud cluster when it first appeared was located on the gradient-level monsoon trough or shear line. The cloud cluster associated with the developed cyclone also was fully developed. As a fully developed cloud also, 97% of them were on the monsoon shear line. This close association of monsoon depressions and cyclones with the monsoon trough suggests that the development,
7.8
Tropical-Midlatitude Interaction in the Australian Region
191
intensification and movement of the cyclone are probably governed by fluctuations of the aircurrents that converge at the troughline. McBride and Zehr (1981) and McBride and Keenan (1982) found that a strengthening of monsoon westerlies equatorward of a disturbance was favorable for its development into a tropical cyclone in many cases. According to Love (1985a,b), the strengthening of the westerlies usually followed cold surges in the South China Sea. Studies reveal that a monsoon disturbance may undergo several transformations during its life period. In most cases, it starts off as a low or depression in or around a monsoon trough and takes a day or two to develop into a tropical cyclone over a warm ocean surface. As long as it remains over a warm ocean and other environmental conditions continue to be favorable, the cyclone rages in full fury but on entering land or a cold SST anomaly area it rapidly transforms itself back again into a depression or low pressure system. A reverse transformation of a low or depression into a tropical cyclone appears to occur when a low or depression after extensive traveling over land enters a warm ocean. Studies by Paterson and Bate (2001), referred to earlier in this section report several cases of such transformations in the life of a tropical disturbance in the Australian region.
7.8 Tropical-Midlatitude Interaction in the Australian Region During Australian summer, wave disturbances in midlatitude baroclinic westerlies of the southern hemisphere which usually move along latitudes south of about 35◦ S at surface during Austral summer often develop large amplitudes and interact with the monsoon circulation over the continent and adjoining oceans. An example of such an interaction during onset of monsoon was discussed by Davidson et al. (1983). These wave disturbances also interact with tropical depressions and cyclones which may come under their influence dung their eastward movement along the southern parts of the continent. There are several cases on record of such interactions in the past (see, e.g., McBride, 1987; Saha and Saha, 2001a). The interactions affect tropical convection and lead to enhanced rainfall in certain sectors, especially over the southwestern and southeastern parts of the continent. Saha and Saha (2001a) discuss the life history of two tropical disturbances, ‘BOBBY’ and ‘JASON’, particulars of which including their dates of initial formation and movement and later recurvature are given in Fig. 7.14. Of the two tropical disturbances, Bobby had a chequered career. Starting life as a simple low pressure wave in the monsoon trough zone (TCZ) near Darwin in Arnhem Land on 18 February 1995, Bobby continued to move southwestward as a closed ‘Low’ along the northwestern coast of Australia but on emerging over the adjoining ocean after 2 days of land travel it rapidly developed into a depression on 21 February and a tropical cyclone the following day. From 22 February onward, it gradually came under the influence of a midlatitude W-ly trough which was approaching Australia from the west. On 24 February, the westerly trough reached
192
7 Monsoon over Australia
Fig. 7.14 Dates, approximate locations, central pressures, intensities, and tracks of two tropical disturbances in the Australian region (Courtesy: NCEP/NCAR Reanalysis)
the extreme southwestern part of the continent and interacted directly with Bobby which was then centered near Onslow at about 21◦ S, 115◦ E. The interaction led to a coupling of the two wave disturbances and for the following 2 days they moved together over the sandy deserts of Western Australia. However, the umbilical cord was soon broken and Bobby moved away southeastward across the southern coast of Australia. A satellite view of Bobby when it was located over the desert area is in Fig. 7.15. The coupling between Bobby and the quasi-stationary wave is clearly suggested by the analyses in Fig. 7.16(a,b) respectively. Through massive warm air advection in the southeast and cold air advection in the northwest of the coupled trough, an isallobaric gradient forced the tropical cyclone to recurve and move in a southeasterly direction as the trough in the midlatitude westerlies in the south moved away eastward in the following 4 days.
7.8.1 Northerly and Southerly Bursters During the movement of midlatitude W’ly waves across Southern Australia, cyclonic and anticyclonic circulations associated with these waves follow each other in quick succession as they move eastward, with gale-force winds blowing from the
7.8
Tropical-Midlatitude Interaction in the Australian Region
193
Fig. 7.15 NOAA-12 satellite view of the tropical cyclone ‘BOBBY’ over the desert of southwestern Australia on 26 February 1995
hot desert lands to the southern oceans ahead of a cyclonic circulation and cool humid winds blowing from ocean to land in its rear when the cyclonic circulation is replaced by anticyclonic circulation. Locally, these winds are known as Northerly and Southerly Bursters respectively, because of the suddenness with which they set in, or one replaces the other. They can occur in any part of the year, but are more common during summer. The southeastern parts of Australia and Africa are particularly vulnerable to occurrence of bursters. According to the Meteorological Glossary (Second Edition) of the American Meteorological Society (2000), a southerly burster is ‘a sudden shift of wind to the southeast in the south and southeast parts of Australia, especially frequent on the coast of New South Wales near Sydney in summer. It occurs in the rear of a trough of low pressure that is followed by the rapid advance of an anticyclone from west Australia. After some days of hot, dry northerly wind, cumulus clouds approach from the south, the wind drops to calm and then sets in suddenly from the south, sometimes reaching gale force. Temperature at Sydney has fallen from 38 to 18◦ C in 30 min. The average summer frequency of bursters at Sydney is 32. Similar winds are experienced in the east of South Africa, especially near Durban.
194
7 Monsoon over Australia
Fig. 7.16 Analyses of (a) temperature and (b) wind at 925 hPa during interaction of tropical cyclone ‘Bobby’ with an eastward-propagating midlatitude W’ly trough disturbance over the Australian region, 12 GMT, 24 February 1995
Chapter 8
Monsoon over Africa (Region – V)
8.1 Introduction Africa is a continent of seemingly great ecological diversity. Here, in the central part of the continent, there are the equatorial rainforests, blessed with perennial heavy precipitation, which sustain life and activities of the bulk of the African population, whereas, on either side of this green belt, barring a few coastal areas, environmental conditions are too dry and harsh with little or no rain to sustain any large-scale vegetation or population. To the north of the equatorial belt lies the great Saharan desert and to the south the great Kalahari-Namib desert. Away from the equatorial zone, only the southeastern part of Africa is blessed with some good amount of rainfall during southern summer and sustains a sizable population. Though the people of Africa have come to terms with this diversity of climatic conditions, the importance of monsoons as sources of much-needed rainfall for their livelihood and activities was realized a long time ago. The erstwhile colonial powers set up a sparse network of observatories in their territories, which has been expanded in modern times, but the density of meteorological observations over the continent as a whole still leaves much to be desired. From studies carried out so far (e.g., Kendrew, 1953; Thompson, 1965; Griffiths,1972), we have some knowledge and understanding of the climates of Africa and the atmospheric circulations associated with weather patterns over different parts of the continent (e.g., Soliman, 1958; Johnson and Morth, 1960; Flohn, 1960; Saha and Saha, 2001b). Attempts have also been made to understand the origin, development and movement of disturbances that not only affect the weather and climate of Africa but also in many cases, on emerging over the Atlantic ocean, develop into depressions and hurricanes (e.g., Carlson, 1969a, b; Rennick, 1976; Reed et al., 1977; Norquist et al., 1977; Saha and Saha, 2002). In the present text, we have used an 18 year (1979–1996) mean climatological dataset and NCEP/NCAR Reanalysis to look at the monsoons of Africa and some of its associated problems over the northern as well as the southern parts of the continent. Since the concept of summer and winter changes with the hemisphere, we have taken February to represent summer monsoon over Southern Africa and August to represent summer monsoon over Northern Africa.
K. Saha, Tropical Circulation Systems and Monsoons, C Springer-Verlag Berlin Heidelberg 2010 DOI 10.1007/978-3-642-03373-5_8,
195
196
8 Monsoon over Africa
8.2 Physical Features and Environment Important among the factors which shape the climate of Africa are its geographical location, land-sea configuration, oceanic environment, and topography (see Fig. 8.1). Situated across the equator, the continent of Africa extends latitudinally from about 35◦ S to 37◦ N. However, it is only the eastern part (east of about 12◦ E) which has this stated latitudinal dimension. The western part lies wholly in the northern hemisphere and extends from about 5 to 37◦ N. Longitudinally, the continent as a whole extends from about 15◦ W to 50◦ E. But it is only the northern part (north
Fig. 8.1 Map of Africa showing topographic heights (m), main river systems, and ocean currents around it heights above 500 m a.s.l. are shaded. Double-dashed arrows show the direction of the ocean currents. Rivers are denoted by the letter R
8.2
Physical Features and Environment
197
of about 5◦ N) which has this longitudinal dimension. The southern part extends from about 12 to 45◦ E and gradually narrows down to Cape of Good Hope at its southernmost point. The difference in latitudinal-longitudinal dimension between the western and the eastern parts of Africa is significant, since it restricts monsoon to the northern hemisphere in the western part, while allowing it to oscillate between the hemispheres in the eastern part. The oceans around the continent consist of the Atlantic Ocean in the west, the Indian Ocean in the east, and the Mediterranean Sea to the north. Some relatively minor waterways separate the African continent from Asia. An important aspect of Africa’s oceanic environment is the presence of cold and warm ocean currents that through air-sea interaction condition the airmasses that participate in the monsoon circulation over Africa. These include the cold Benguela current of the South Atlantic which causes intense upwelling along the coast of southwestern Africa and the cold Canaries current of the northern Atlantic that washes the shores of northwestern Africa. These ocean currents form parts of the subtropical anticyclonic gyres of the Atlantic Ocean on either side of the equator. A warm equatorial ocean current flows eastward in the vicinity of the Guinea coast of western Africa almost throughout the year. In the Indian Ocean, the warm Agulhas current flows southwestward through the Mozambique Channel by the side of southeastern Africa during southern summer, while, a strong cold current flows northeastward off the Somali coast of East Africa during northern summer. The topography of Africa varies widely between its western and eastern parts (see Fig. 8.1). To the northwest of the great Saharan desert lies the lofty Atlas mountain range. Several low mountain ranges and high grounds appear in different parts of the Saharan desert and along the coasts of Upper and Lower Guinea. Over the subSaharan belt of Sahel lying between about 5 and 20◦ N lie a series of mountains interspersed by extensive desert lowlands. Well-known among these are the Guinea and Nigerian mountains the average height of which varies from 1.0 and 1.5 km a.s.l., the Marra mountain rising to a peak of about 3 km a.s.l along the western border of the Darfur region of Sudan, and the Ethiopian mountains which rise to well over 4 km. Between the Nigerian and the Marra mountains lies the vast Lake Chad lowlands and between the Marra and the Ethiopian mountains lies the extensive lowland of Sudan with River Nile flowing through it. The Ethiopian mountains are divided in the middle by the well-known Denakil desert of Eritrea and the Great Rift Valley of East Africa, while the mountains descend steeply in the east to the plains of Somalia desert. The high-rise Mountain Range of Ethiopia extends southward all the way to South Africa with only a few gaps and rift valleys. The average altitude of the eastern mountain ranges that lie between longitudes 30 and 35◦ E is about 2– 3 km a.s.l. In Southern Africa, most of Angola, Namibia and Botswana are located on high plateaux, the average altitude of which varies from 1.5 to 2.0 km a.s.l. To the north of these extensive plateaux and west of the eastern high mountain ranges lie the vast lowlands of Congo-Zaire which extend northward across the equator into the northern deserts.
198
8 Monsoon over Africa
The importance of the high plateaux of Angola, the lofty mountains of East Africa, and the coastal high grounds of Guinea to the economy of Africa can hardly be over-emphasized, for they form the source regions of Africa’s main river (R) systems: (i) R. Nile (both Blue and White) which flows northward through Sudan and Egypt; It is, indeed, a Nature’s bounty that the world’s longest river (6695 km), is able to survive a most tortuous and harsh journey through the dry desert to reach its ultimate destination, the Mediterranean Sea. (ii) R. Congo-Zaire which originates from the Angolan Plateaux and flows through the lowlands of the western part of equatorial Africa; (iii) R. Zambesi which descends from the Angola-Katagan plateaux and flows through Zambia, Zimbabwe and Mozambique; (iv) R. Niger which rises from the mountains of Guinea and after passing through the desert States of Mali and Niger enters Nigeria to flow into the Gulf of Guinea. Besides these main river systems, there are several others, large and small, not mentioned, the origin of which can be traced to orography. Some of the well-known lakes in Africa are fed by rivers draining from nearby mountains.
8.3 Climates of Africa and Surrounding Oceans 8.3.1 Sea Surface Temperature and Wind The salient features of the continent’s oceanic environment are revealed by mean Sea Surface Temperature (SST) and wind at 10 m above ocean surface, presented in Figs. 8.2 and 8.3 for February and August respectively. The cross-isotherm tradewinds reveal the oceanic origin of the cold airmasses which enter into the monsoon circulation over the continent. But for a northward seasonal shift of the thermal field, the pattern of thermal advection during August is very similar to that during February, as shown by Fig. 8.3.
8.3.2 Air Temperature Figures 8.4 and 8.5 presents the mean air temperatures over Africa and surrounding areas at 925 hPa (proxy for surface) during February and August respectively. In February, two latitudinal zones of temperature maximum appear over the continent, one along about 5–10◦ N and the other along 25◦ S, which are separated by a zone of temperature minimum over the equatorial zone. The minimum over the equatorial zone is clearly the result of cold air advection from nearby cold oceans. The August temperatures are shown in Fig. 8.5. In August (Fig. 8.5), the seasonal movement of the temperature field is wellmarked. The February maximum temperature belt along 10◦ N has moved to about 25◦ N over western Africa, and to about 15–20◦ N over Central Africa. In the Southern hemisphere, the belt of February maximum temperature along 25◦ S has
8.3
Climates of Africa and Surrounding Oceans
199
Fig. 8.2 Mean SST (◦ C) and 10 m wind (m s–1 ) above surface over oceans around Africa during February. Areas with SST 27◦ C and above are shaded
shifted to about 10◦ S. The two hemispheric belts of maximum temperatures are separated by a belt of minimum temperature along the equator. The temperature field in the upper troposphere, say 200 hPa (not shown), is similar to that at surface during both February and August with two well-marked zones of maxima appearing above the low-level maxima, one in each hemisphere, separated by a zone of minimum over the equatorial belt.
8.3.3 Isobaric Height (gpm) Consistent with the distribution of surface temperature, well-defined troughs (dashed lines) of low pressure appear over the warmest zones and ridges (continuous lines) of high pressure over the coldest zones in the geopotential fields at 925 hPa, shown in Figs. 8.6 and 8.7, for February, and August respectively. In February (Fig. 8.6), the trough in the northern hemisphere is located along latitude about 5–10◦ N, while that in the southern hemisphere lies along about 18◦ S. A ridge of cold high appears along about 5◦ S. A height maximum is located over northern Africa with its ridge along about 25◦ N. Over the ocean, heights are maximum over the subtropical belts of the two hemispheres and minimum near the
200
8 Monsoon over Africa
Fig. 8.3 Mean SST (◦ C) and 10 m wind (m s–1 ) above surface over oceans around Africa during August. Areas with SST 27◦ C and above are shaded
equator. A low pressure area appears over the Mozambique Channel and adjoining Madagascar area off the southeastern coast of Africa. A ridge of the subtropical high of the Southern Indian Ocean lies between the Mozambique Channel low pressure and the seasonal low pressure over Southern Africa. It is remarkable that no such ridge of high pressure appears over the South Atlantic Ocean off the Namibia coast. The geopotential field in August (Fig. 8.7) reveals the usual northward seasonal shift in the locations of the lows and highs in the geopotential height field. For example, the trough along 10◦ N has moved to about 20◦ N, while that along 25◦ S has moved to about 5–10◦ S. The ridge in between the two troughs has moved to the equator. In the upper troposphere, the height field at 200 hPa (not shown) shows a height maximum over the surface ‘heat low’ and a height minimum over the surface ‘cold high’ in both February and August. In August (Fig. 8.7), there has been a general northward shift of the height fields. However, as in February, two troughs of low heights appear at 925 hPa; one in each hemisphere, separated by a ridge of height maximum along the equator. The northern-hemispheric trough is oriented in a more or less east-west direction along
8.3
Climates of Africa and Surrounding Oceans
201
Fig. 8.4 Mean air temperatures (◦ C) over Africa and surrounding oceans at 925 hPa during February
about 20◦ N, while the southern hemispheric trough is oriented more or less meridionally along about 20◦ E. The subtropical high over both the South Atlantic and the Indian Ocean have intensified and moved northward. At 200 hPa in February (not shown), a height maximum appears over the lowlevel ‘heat low’ over Southern Africa. The height gradient to the north of this maximum is small, but is steep poleward of about 30◦ S. At 200 hPa in August (not shown), an intense height maximum appears over Northern Africa with its ridge between 20 and 30◦ N, more or less above the ‘heat low’ below. Height gradient poleward of the ridge is steep, but slacken towards the equator.
8.3.4 Wind and Circulation The climatological winds and circulations over Africa and surrounding oceans at 925 hPa are shown in Fig. 8.8 for February and August. They appear to be consistent with the temperature and height fields, shown in the preceding sections.
202
8 Monsoon over Africa
Fig. 8.5 Mean air temperatures (◦ C) over Africa and surrounding oceans at 925 hPa during August
In February (Fig. 8.8), at 925 hPa, the tradewinds diverging from the subtropical anticyclones over the oceans are very strong. They appear to converge into the equatorial trough of low pressure over the ocean as well as into the heat low circulations over the continent. North/northeasterly winds diverging from the winter anticyclone over northern Africa with its ridge along about 27◦ N appear to converge into the trough of low pressure along about 10◦ N. A number of low level airstreams from the Indian Ocean appear to converge into the ‘heat lows’ over Southern Africa (Namib area) and adjoining Southwestern Indian Ocean on either side of the Mozambique Channel. These include the tradewinds from the northern as well as the southern Indian Ocean. The tradewinds from the southern Atlantic subtropical anticyclone also converge into the heat low over the Namib area.. The hatched areas between latitudes 15 and 20◦ S mark the locations of the ITCZs over Southern Africa and the Southern Indian Ocean during February. At 200 hPa during February (not shown), two powerful anticyclonic circulations appear over Africa, one in each hemisphere; that over South Africa with ridge along about 20◦ S, and that over North Africa with ridge between 5 and 10◦ N. There is
8.3
Climates of Africa and Surrounding Oceans
203
Fig. 8.6 Isobaric height (gpm) at 925 hPa over Africa and adjoining oceans during February
strong cross-equatorial flow from the southern to the northern hemisphere anticyclone. Westerlies blow poleward of the ridge of the anticyclone in each hemisphere. The strongest westerlies with speeds ≥50 m s–1 appear in the northern hemisphere along about 25◦ N. In August (Fig. 8.9 ), the low level circulation is very similar to that in February, except that there has been an appreciable northward shift of the circulation systems in keeping with the seasonal movement of the sun. The tradewinds diverging from the subtropical anticyclones converge into the circulation around the equatorial trough of low pressure located along mean latitude about 18◦ N, producing the ITCZ (hatched). In the southern hemisphere, tradewinds diverging from the subtropical
204
8 Monsoon over Africa
Fig. 8.7 Isobaric height (gpm) at 925 hPa over Africa and adjoining oceans during August
high pressures over the Atlantic as well as the Indian Ocean converge into the circulation around a trough of low pressure over the Congo/Zaire region. Tradewinds diverging from the subtropical high pressure belt of the Southern Indian Ocean after crossing the equator converge into the circulation around the equatorial trough of low pressure over India and usher in the monsoon to the Subcontinent. The maps at 200 hPa show strong anticyclonic circulations above the lowertropospheric cyclonic circulations in both the hemispheres. They are stronger over Southern Africa during February and over Northern Africa during August. Strong easterlies appear over the equatorial belt during August.
8.3
Climates of Africa and Surrounding Oceans
205
Fig. 8.8 Field of wind (m s–1 ) and streamlines (lines with arrows showing direction of airflow) over Africa and surrounding oceans at 925 hPa during February. Areas with isotachs ≥5 m s–1 are shaded. ITCZs are hatched
8.3.5 Rainfall over Africa Figure 8.10 shows the climatological (1971–2000) mean precipitation rate over Africa and surrounding oceans during (a) February, and (b) August. In February (Fig. 8.10a), the rainfall over Africa is mainly concentrated in the southern hemisphere, stretching southeastward from near-equatorial west coast of southern Africa, across Angolan and Katangan plateaux to the East Coast mountains in Zambia and Mozambique. The belt joins up with an area of extremely heavy rainfall (≥13 mm day–1 ) over Madagascar area. In the northern hemisphere, there is little rain in February.
206
8 Monsoon over Africa
Fig. 8.9 Field of wind (m s–1 ) and streamlines (lines with arrows showing direction of airflow) over Africa and surrounding oceans at 925 hPa during August. Areas with isotachs ≥5 m s–1 are shaded. ITCZs are hatched
During August (Fig. 8.10b), the whole rainbelt appears to have shifted to the northern hemisphere, leaving only a small area of moderate rainfall over the central Congo basin south of the equator. The main concentrations of rainfall in the northern hemisphere appear to be in the mountainous regions of Upper Guinea, Nigeria and Cameroons. A belt of heavy rain also appears along the ITCZ over the North Atlantic. The rainbelt over the land appears to be limited to latitudes between the equator and about 18◦ N. There is little rain over the vast Saharan desert. There is little rain in the Western Indian Ocean, except on the windward sides of the mountains of Madagascar and some other islands. Similarly, there is no rain over the South Atlantic. But a rainbelt appears over the subtropical oceans of the southern hemisphere, which appears to be due to the activity of the eastward-propagating disturbances of the midlatitude baroclinic westerlies.
8.4
Equatorial Westerlies over Africa
207
Fig. 8.10 Climatological (1971–2000) mean precipitation (mm day–1 ) over Africa during (a) February and (b) August (Isohyets ≥2 mm day–1 are shaded)
The above-mentioned distribution of seasonal rainfall over Africa and surrounding regions is in substantial agreement with that of the outgoung longwave radiation (OLR) for (a) February and (b) August (Not shown here for lack of space).
8.4 Equatorial Westerlies over Africa Low level equatorial westerlies over Africa have been referred to in literature (e.g., Flohn, 1960), as part of a broad westerly airstream extending from Africa to the Indian Ocean along the equator. But, in reality, the equatorial westerlies, are the zonal components of the tradewinds which after diverging from the subtropical anticyclones of the two hemispheres blow towards the ‘heat lows’ over the continent, after crossing the equator. But the question is: how far do they extend eastward? To find out the realities about these winds, we suitably combined the zonal component of the wind with vertical motion along the equator obtained from NCEP/NCAR reanalysis and arrived at the resultant streamlines which are shown in Fig. 8.11 for (a) February and (b) August (Note that the meridional component of the wind was ignored to arrive at the resultant streamlines). The February circulation (Fig. 8.11a) shows that after rising over the coastal Guinea Mountains, the westerlies descend on the lowlands of Zaire. They rise again as westerlies at the Kenya-Ethiopian Mountains only to merge with the ascending easterly flow from the Indian Ocean. However, in August (Fig. 8.11b), the upperlevel equatorial easterlies over the Indian Ocean descend strongly over its western part and return as equatorial westerlies.
208
8 Monsoon over Africa
It is, therefore, evident that the equatorial westerlies over Africa are confined to a shallow layer below about 800 hPa and have no connection with those over the Indian Ocean.
Fig. 8.11 Equatorial circulation over Africa during (a) February and (b) August
8.5
The Equatorial Trough over North Africa – Its Zonal Anomaly
209
8.5 The Equatorial Trough over North Africa – Its Zonal Anomaly The trough of the ‘heat low’ over North Africa during summer has a zonal variation, being located farther north in its western and eastern sections than the central section between 10 and 30◦ E where it lies along about 18◦ N. Soliman (1958) who studied the problem almost half a century ago correctly discounted the idea of a single convergence zone over Africa. He regarded the latitudinal extent of the heat low circulation over Africa as a thermal belt and proposed two frontal zones, one along the southern boundary of the thermal belt which he called the Intertropical Front (ITF) and the other along the northern boundary which he called the Subtropical Front (STF). The approximate locations of these fronts, as deduced from Reanalysis, are marked in the mean sea level pressure map for August in Fig. 8.12. According to Soliman, both the fronts are baroclinic zones, the baroclinicity of which varies with season; that at ITF becoming maximum and strongly correlated with the upper-air easterly jetstream in summer and that at STF becoming maximum and strongly correlated with the upper-air westerly jetstream in winter. Orography adds a new dimension to the zonal anomaly of the equatorial trough of low pressure over North Africa by giving it a stationary wave structure, as would be evident from Fig. 8.13 which gives the zonal anomaly (deviation from the zonal
Fig. 8.12 Climatological (1979–1996) mean sea level pressure (hPa) distribution over North Africa during August showing the locations of the Subtropical Front (STF) and the Intertropical Front (ITF), as suggested by Soliman (1958)
210
8 Monsoon over Africa
Fig. 8.13 Mean zonal anomaly (deviation from the zonal mean) of msl pressure P’ (mb), 925 hPa temperature T’ (◦ C), 925 hPa specific humidity H’ (g/kg), and rainfall R’ (mm day–1 ) along the latitudinal belt, 10–15◦ N, over Africa between longitudes 10◦ W and 40◦ E in August (Topographic height (m) is shown by hatching)
mean) of MSL pressure (P’), temperature (T’) and specific humidity (H’) at 925 hPa, and Precipitation (R’) over the latitudinal belt, 10–15◦ N, during August. An association of the zonal stationary wave with the mountain ranges of central and eastern North Africa stands out in Fig. 8.13, the mountains being associated with high pressure, low temperature, high specific humidity and high rainfall and the intervening valleys or lowlands with low pressure, high temperature, extreme dryness and scanty rainfall. The stationary wave appears to have average wavelength of 1500–2000 km and amplitude 1.5 hPa in the pressure field and 1.5◦ C in the temperature field.
8.6 Structure of the Circulation Associated with the Equatorial Trough The structure of the vertical circulation across the quasi-stationary monsoon wave over Central North Africa is considered in two sections: a zonal section along latitude 12.5◦ N and a meridional section along 17.5◦ E.
8.6
Structure of the Circulation Associated with the Equatorial Trough
211
8.6.1 Zonal Circulation The zonal section in Fig. 8.14 shows the resultant streamlines constructed from mean values of the zonal component (u) of the wind and the vertical p-velocity (ω) along latitude 12.5◦ N during August. Figure 8.14 shows the low level westerlies rising gently over the Nigerian mountains (for topography, see Fig. 8.13), sinking over the Lake Chad area, and then rising strongly over the slopes of the high Marra mountains with peak at 3071 m. Sinking strongly on the leeside of the mountains and after flowing over the Nile valley of Sudan the airflow rises again steeply over the Ethiopian mountains with peak at about 4620 m above mean sea level. There is little doubt that the monsoon stationary wave over Africa is maintained by the joint action of heating and topography of the desert surface.
Fig. 8.14 Mean resultant streamlines based on (u, ω) values along latitude 12.5◦ N latitude over Central North Africa during August
8.6.2 Meridional Circulation Figure 8.15 shows the resultant streamlines based on the mean values of the meridional component (v) of the wind and the vertical p-velocity (ω) along longitude of 17.5◦ E during August.
212
8 Monsoon over Africa
Fig. 8.15 Mean resultant streamlines based on (v, ω) values along a mean longitude of 17.5◦ E over Central North Africa during August
It may be noted that the structure of the mean meridional circulation along meridian 17.5◦ E which passes through almost the middle of Central Africa differs greatly from that along the Greenwich meridian, presented earlier in Fig. 1.12 in Chap. I. The circulation along the Greenwich meridian shows as many as three latitudinal zones of upward motion, one to the south of the equatorial trough zone (ITF) which constitutes the rising branch of the monsoon circulation with divergence aloft, the second representing a zone of low-level convection over the middle of the Saharan desert between about 15 and 22◦ N evidently due to intense surface heating, and the third over the coastal belt north of 30◦ N where the cool, moist air from the Mediterranean Sea converges into the ‘heat low’ circulation. In contrast, the circulation along 17.5◦ E shows a well-defined zone of strong upward motion in the equatorial trough zone south of about 18◦ N with strong descending motion above and a zone of feeble convection where the cool, moist Mediterranean Sea air converges into the desert heat low circulation
8.7 Origin of African Wave Disturbances 8.7.1 Early Studies Ever since it had been suspected that many of the wave disturbances that develop into hurricanes over the Atlantic Ocean emerge from the African continent, there has been frantic effort to trace the place of their origin over Africa and study their dynamics.
8.7
Origin of African Wave Disturbances
213
Carlson (1969b), who first noted that the majority of them originated in the mountainous region of east Central Africa thought that they might arise from the effect of afternoon heating of the mountain surfaces. Frank (1970) put forward the view that the mechanical lifting of the easterly airflow from the Indian Ocean by the mountains of Ethiopia might be responsible for them. However, Burpee (1972) who carried out a systematic power- and cross-spectrum analysis of available upper-air data discounted both these speculations, though his study supported the idea of origin of African wave disturbances somewhere over east Central Africa. He sought to explain the formation of these waves by applying Charney and Stern (1962)’s quasi-geostrophic theory of an internal baroclinic jet to the case of the midtropospheric easterly jet over Africa. He sought to prove that the horizontal and vertical shear of the mean zonal wind associated with the jet satisfied the condition of dynamic instability of the flow under certain strict boundary conditions only. One of these conditions stipulates that there should be no horizontal transports of heat and momentum from neighboring regions to the midtropospheric jet through appropriately-placed vertical boundaries at the poleward and equatorward limits of the influence of the jet. In the absence of adequate data, he just assumed the validity of the restrictive boundary condition and attributed the origin of African wave disturbances to the dynamic instability of the midtropospheric easterly jet on the basis of the above-mentioned Charney-Stern theory. Rennick (1976) emphasizes a barotropic mechanism related to the existence of the midtropospheric easterly jet, while others (Burpee, 1972; Norquist et al., 1977) emphasize almost equal contributions from barotropic and baroclinic energy conversions for the formation of African wave disturbances. Observations, sparse as they may be, however, have a different story to tell. They show that the boundary condition stipulated for the instability of the easterly jet is frequently violated in the real atmosphere when the easterly jet comes under the influence of eastward-propagating disturbances in midlatitude baroclinic westerlies and there are large-scale meridional fluxes of heat and momentum across the boundaries of the jet during such interaction. In fact, the fluxes may vary so much during the period of such interaction that they may lead to large variations in the strength and stability of the jet. Saha and Saha (2002), using twice-daily Reanalysis data, examined the mechanism of this interaction of the jet with midlatitude baroclinic waves of the northern hemisphere in ten cases of formation of African wave disturbances which later in their life cycle, when they emerged over the Atlantic Ocean, formed the nuclei of Atlantic hurricanes. Some aspects of their study and findings are reviewed in the section that follows.
8.7.2 Influence of Midlatitude Forcing The midlatitude forcing arises from the interaction of the quasi-stationary subtropical zonal wave which lies over Northern Africa (which Soliman (1958) called the STF (see Fig. 8.12) and we have called the TCZ in the present text) with
214
8 Monsoon over Africa
eastward-propagating waves in baroclinic westerlies that move eastward across southern Europe and adjoining Mediterranean Sea. The quasi-stationary wave, the existence of which is clearly revealed in the subtropical zonal temperature and pressure anomaly fields at low levels along a mean latitude of say 25◦ N, consists of alternate sectors of high and low temperature/pressure anomalies, the high pressures being generally associated with the cool ocean or land areas under oceanic influence, and low pressures with the heated land. The wave appears to have an average length of about 3000–4000 km and pressure amplitude of 3–5 hPa over the central land section. Thus, structurally, the TCZ over Africa has wave characteristics very similar to those of the traveling baroclinic wave disturbances of midlatitude westerlies. It also appears to be baroclinically as active as the ITCZ, as suggested by Soliman (loc. cit.). Perhaps, it is because of this structural similarity that the TCZ interacts frequently with large-amplitude midlatitude waves. In their study on the origin of African wave disturbances, Saha and Saha (2002) adopted the following procedure to trace the place of their origin: they selected ten named Atlantic hurricanes which appeared to have a history of origin over mainland Africa. There was no rationale behind the selection of these hurricanes except that their predecessor disturbances had come from some part of Africa. They noted the date the predecessor of the selected hurricane had left the west coast of Africa and worked backward in time using all available synoptic and aerological data over a period of 7–8 days prior to its arrival at the coast. In each case, apart from tracing continuity from daily synoptic maps and satellite cloud imagery, they computed the values of eddy kinetic energy and barotropic and baroclinic energy conversions at 0000 and 1200 UTC daily. Additionally, they examined the daily 0000 and 1200 UTC mean sea level pressure maps as well as maps of wind, temperature and geopotential height at 850, 700, 500 and 200 hPa over a wider area extending from equator to 45◦ N and from 20◦ W to 50◦ E in each case in their search for the first signs of formation of a wave disturbance in different latitude belts over the African longitudes. They found that there was a sequence of events which occurred in practically every case examined and identified the following three stages: (i) The first stage is an in-phase interaction of the quasi-stationary wave (STF) with a large-amplitude wave in midlatitude baroclinic westerlies moving eastward across Southern Europe and adjoining Mediterranean Sea, resulting in the formation of an extended trough-ridge system and strengthening of the circulations between the waves over northwestern Africa; (ii) During next stage, the circulations lead to enhanced meridional exchanges of heat and momentum between the midlatitude wave and the STF; and (iii) In the final stage, the meridional flux of heat and momentum from the midlatitude wave to the equatorial trough zone makes the airflow associated with it dynamically unstable and leads to the genesis of a new wave disturbance. It is well-nigh impossible to give here the full details of analysis of every case examined for lack of space. However, in their study, Saha and Saha (loc. cit.) give
8.7
Origin of African Wave Disturbances
215
details of their analysis of the case relating to the origin of the predecessor of a hurricane named ‘GORDON’ which raged over the Atlantic Ocean in the beginning of September 2000. Readers interested in the details of this case may look up the original paper. Here we present only a few salient points which appear to throw light on the problem of origin and evolution of the disturbance. The first indication of the formation of a wave disturbance in the case of Gordon was available from the circulation at 700 hPa at 12 GMT on 31 August, 2000, shown in Fig. 8.16 The circulation map in Fig. 8.16 reveals the presence of the equatorial trough of low pressure over the Nile valley of Sudan oriented in an NE-SW direction with cold air advection from the northeast to its west and warm air advection from the south/southwest to its east. The cold air that converged into the trough diverged from the subtropical anticyclone that was located at the time over eastern Sudan and adjoining Red Sea area and to the east of a midlatitude large-amplitude W’ly wave trough over the Middle East. However, the eastward extent of this equatorwardmoving cold air appears to be limited to about 25 E along the latitude of the equatorial trough at the time. But the situation changed rapidly during the following 24 h, as shown by a series of three synoptic maps at 12 GMT on 01 September 2000, over North Africa, presented in Figs. 8.17, 8.18, and 8.19.
Fig. 8.16 Location of two midlatitude W’ly wave troughs (thick continuous lines) and the equatorial trough of low pressure over North Africa in the streamline field at 700 hPa at 12 GMT on 31 August 2000
216
8 Monsoon over Africa
The mean sea level (msl) pressure (mb) map is at Fig. 8.17:
Fig. 8.17 Mean sea level pressure (mb) field over North Africa at 12 GMT, 01 September 2000, associated with predecessor disturbance of hurricane ‘GORDON’
The msl pressure map shows the approximate location of the equatorial trough of low pressure between about 20 and 10◦ N and the approximate location of the W’ly trough over the central part of the Mediterranean Sea. The large-amplitude W’ly trough over the Middle East appears to have moved away further east. Figure 8.18 shows the distribution of 850 hPa isotherms (thin continuous lines) and wind vectors over North Africa at 12 GMT on 01 September 2000. The distributions show that the strong influx of cold air from the subtropical anticyclone over North Africa to the equatorial trough zone is limited to the north of the equatorial trough only. To the south of the equatorial trough, warm air flows southward in the west and cold air flows northward in the east. Figure 8.19 which presents the distribution of temperature and wind vectors over North Africa at 500 hPa at 12 GMT on 01 September 2000, appears to give some indication of an influx of warm air from the southeast to the east of the 500 hPa trough (indicated by a short thick continuous line), almost directly above the low-level displaced equatorial trough where warm air flows southward on the southwestern side of the trough. The vertical distribution of thermal advection would lead to upward motion to the west and east of the trough with downward motion in the middle.
8.7
Origin of African Wave Disturbances
217
Fig. 8.18 Distribution of 850 hPa isotherms (C) and wind vectors over North Africa at 12 GMT on 01 September 2000. The line of temperature maximum is indicated by a thick continuous line. Note the line of temperature discontinuity on its equatorward side with warm advection to west and cold advection to east
Thus, Figs. 8.17, 8.18, and 8.19 would appear to be in conformity with a mechanism suggested in Chap. 2 (Fig. 2.7), for formation of a wave disturbance in a quasi-stationary trough of low pressure in the tropics by its interaction with a traveling midlatitude W’ly wave trough.
8.7.3 Sudan – The Breeding Ground The procedure described above in the case of hurricane ‘GORDON’ was applied by Saha and Saha (loc. cit.) to a few of the other hurricanes as well. They found that in most cases examined, the predecessor wave disturbances tended to originate over a broad area of Southern Sudan between longitudes 20 and 30◦ E and latitudes 10 and 15◦ N. As may be seen from Fig. 8.20, this broad area consists mainly of dry desert lowlands of Sudan, bounded on its western, southern and eastern sides by an array of mountains, high and low, of Central African States.
218
8 Monsoon over Africa
Fig. 8.19 Distribution of 500 hPa isotherms (C) and wind vectors over North Africa at 12 GMT on 01 September 2000. The line of temperature maximum is indicated by a thick continuous line. Note the line of temperature discontinuity on its equatorward side with warm advection to west and cold advection to east
Fig. 8.20 Topography of Sudan showing the mountains (denoted by hats) in relation to the desert lowlands (dots) and the mean summer equatorial trough (thick continuous line). Heights below 1000 m are indicated by small hats, between 1000 and 2000 m by medium hats, and ≥2000 m by tall hats
8.8
Structure, Development and Movement of the Waves
219
8.7.4 Role of Topography It seems quite likely that the topography of Sudan, shown in Fig. 8.20, contributes greatly to cyclogenesis over the region One may see that the mountains over the region east of 20◦ E are not quite zonally oriented as they, more or less, are to the west; they lie somewhat in the shape of an arc of a circle around the desert lowlands of Sudan. Starting from the high mountains of Tibesti (peak at 3415 m) at about 20◦ N in the west, the arc passes through the mountains of Marra (3088 m) in Darfur area, a chain of mountains over Central African Republic (peaks between 1000 and 2000 m), the mountains of northern Uganda and Kenya (Kinyeti, 3187 m), to the high mountains of Ethiopia (4533 m) in the east. Details of the map including the heights of the peaks wherever stated are taken from the National Geographic Atlas of the World (eighth edn.), Washington, DC.
8.8 Structure, Development and Movement of the Waves It is the author’s belief that the topography of Sudan plays an important role in the process of development of the African wave disturbance by allowing the meridional flux of heat and momentum from the north during interaction of the mean equatorial trough of low pressure with a midlatitude baroclinic wave disturbance. The flux leads to an enhanced adiabatic warming and fall of pressure in the western part of the equatorial trough and cooling and rise of pressure in the eastern part, thereby leading to development of a perturbation low (L) pressure with anticlockwise circulation, as shown schematically in Fig. 8.21. The wave allows moist monsoon winds from the southwest to converge on the mountain ranges located to the southwest of the perturbation center, especially on those of Cameroons, Central African Republic and Uganda and produce heavy rainfall. Such developments will only occur when the ITCZ over Central North Africa
Fig. 8.21 Schematic illustrating low level circulation around the suggested incipient wave disturbance (L) with probable locations of the ITCZ and the TCZ
220
8 Monsoon over Africa
associated with the equatorial trough of low pressure comes under the influence of a midlatitude baroclinic wave disturbance. After formation, the wave disturbance moves westward at a variable speed. In the case of the predecessor of GORDON, the average speed was about 13 m s–1 . The amplitude, phase speed and activity appear to undergo remarkable changes as the disturbance moves westward. The amplitude remains small and it moves slowly with light to moderate rainfall till it reaches a longitude near about 10◦ E. But, once it crosses this longitude, the amplitude increases and it moves faster. Also, the amount of rainfall appears to increase considerably as it moves over the southern part of western Africa. Of course, as pointed out by Carlson (1969b), there could be several reasons for this change in wave structure and activity, such as meridional coupling of the disturbance with the west African ‘heat low’, greater convective instability of the lower troposphere, and increased availability of moisture from the Gulf of Guinea to the disturbance zone.
8.9 Interaction of South African Monsoon with Midlatitude Waves of the Southern Hemisphere Like its counterpart in the northern hemisphere, the monsoon trough of the southern hemisphere periodically comes under the influence of eastward-propagating
Fig. 8.22 Schematic showing interaction of the heat low circulation over Southern Africa with a traveling midlatitude W’ly wave trough. The extended trough is indicated by a thick dashed line. Hatched area shows the likely belt of clouding and precipitation
8.9
Interaction of South African Monsoon
221
midlatitude baroclinic wave disturbances of the southern hemisphere during southern summer (January/February). These are eastward-propagating wave disturbances in baroclinic westerlies of the Southern Atlatic Ocean. Before they reach Africa, they have small amplitude and large phase speed. But on entering Africa, they slow down and interact with the quasi-stationary monsoon low pressure system over Southern Africa, resulting in the formation of an extended trough and an elongated belt of deep convection and clouding and precipitation in the trough zone, which moves eastward with the disturbance over the southern Indian Ocean after its separation from the quasi-stationary trough. An example of an interaction is shown in Fig. 8.22 in which the extended trough is shown by a thick dashed line. The interaction leads to amplification of both the disturbances. The eastward movement of these waves across the southern Indian Ocean at regular intervals during southern summer can be seen in daily satellite cloud imagery.
Chapter 9
Monsoon over South America (Region – VI)
9.1 Introduction Until recently, there had been considerable uncertainty about the presence of a monsoon circulation over South America. Only occasionally a positive note was sounded by a few (e.g., Vulquin, 1971) on the basis of their study of exchange of airmasses between the hemispheres at certain times of the year. Available publications on the climate of South America (e.g., Kendrew, 1953; Taljaard, 1972; Boucher, 1975; Schwerdtfeger, 1976; Pearce and Smith, 1990; Satyamurti et al., 1998) also make no direct mention of a monsoon circulation over South America, though several of them describe climatic conditions which are typical of a monsoon regime. This view is well supported by Fig. 9.1 which presents the distribution of mean monthly rainfall at three southern hemispheric stations (Darwin in Australia, Goias in South America, and Lilongwe in South Africa) all located approximately in the same latitude belt. So, the question raised was: If Australia and South Africa were recognized as monsoon regions, why not South America? One of the likely reasons for not mentioning monsoon in the case of South America earlier, one may imagine, might have been frequent failure of the observed surface wind direction over the continent to show seasonal reversal which was used as the sole criterion for definition of monsoon. Further, it was thought by some (e.g., Ramage, 1971) that the continent was dimensionally too small and narrow to generate a monsoon circulation. However, some recent studies appear to have removed the doubt. Following an original remark by Halley (1686) that the monsoon constitutes a large-scale perturbation in the tradewind circulation, these studies (e.g., Van den Dool and Saha, 1993; Saha et al., 1994, 1998; Zhou and Lau, 1997) have shown that if the background annual mean wind were removed from the observed wind, the remainder would reveal the presence of the monsoon winds. Using the same technique, Zhou and Lau ’s study showed not only the seasonal reversal of the wind at 850 hPa but also a mean meridional circulation with rising motion over the heated continent and sinking motion over the neighboring cold ocean.
K. Saha, Tropical Circulation Systems and Monsoons, C Springer-Verlag Berlin Heidelberg 2010 DOI 10.1007/978-3-642-03373-5_9,
223
224
9
Monsoon over South America
Fig. 9.1 Distribution of mean monthly rainfall at Goias (16◦ S, 50◦ W) in Brazil (South America), Lilongwe (14◦ S, 34◦ E) in Malawi (Southern Africa), and Darwin (12.5◦ S, 131◦ E) in northern territory (Australia), all located within a 5-degree latitude belt in the southern hemisphere (Data from Pearce and Smith, 1990)
9.2 Physical Features and Environment 9.2.1 Physical Dimension of the Continent Situated between latitudes about 12◦ N and 56◦ S and longitudes about 80 and 35◦ W, South America is actually an interhemispheric continent like Africa with landmasses on both sides of the equator. Table 9.1 gives a comparative statement of the Table 9.1 Approximate longitudinal width (km) of landmass in different 5-degree latitude belts of South America, Southern Africa and Australia Longitudinal width (km) Latitude belt(s)
South America
South Africa
Australia
0–5◦ 5–10 10–15 15–20 20–25 25–30 30–35 35–40 40–45 45–50 50–55 55–60
3850 4450 3850 3225 2630 2250 2000 1500 1075 850 725 Ocean
3075 2775 2775 2525 2100 1700 1500 Ocean Ocean Ocean Ocean Ocean
Ocean Ocean 800 2400 3250 3800 3700 1000 300 Ocean Ocean Ocean
9.2
Physical Features and Environment
225
approximate longitudinal width (km) of landmass in different 5-degree latitude belts of South America as against those of South Africa and Australia which are well-known to be monsoonal regimes of the southern hemisphere. It may be seen from Table 9.1 that South America has the largest dimension or landmass coverage of a continent over the tropical belt (0–30◦ S) as compared to the other two southern hemisphere continents, even without a large chunk of it located north of the equator. However, the bulk of the continent of South America is located in the southern hemisphere where it covers not only the tropical belt but also extends deep into the extratropics almost up to the Antarctic. Its shape and dimension is such that its widest part lies between 5 and 10◦ S, and narrows down both northward and southward.
9.2.2 Topography Figure 9.2 shows a map including the smoothed topography of South America. The most striking feature of the topography of South America is the presence of the high Andes Mountains over its western part, running almost north–south close to its boundary with the Pacific Ocean. It has its widest part and the highest peaks rising well over 5000 m over Bolivia and northwestern Argentina. Among the other important orographic features which appear in Fig. 9.2 are a few plateaux or highlands, such as the Venezuelan highlands in the north and the Brazilian plateau in the east. Major lowland areas are the great Amazon basin, southern and northeastern Brazil, Paraguay, Uruguay, and eastern Argentina, along with strips of lowlands of varying width located along the coastal belts.
9.2.3 Oceanic Environment and Its Influence on Climate But for its narrow link with Panama in the northwest, South America is totally surrounded by oceans; the Pacific Ocean to the west and south, the Atlantic Ocean to the east and northeast, and the Caribbean Sea to the north. These oceans play important roles in conditioning the weather and climate of the continent, because of the influence of warm and cold ocean currents and different circulation systems that prevail over the different parts of the oceans. Over both the Pacific and Atlantic Oceans, the tradewinds diverging from the subtropical high pressures blow towards the equatorial trough of low pressure. In the Pacific Ocean, the cold Humboldt or Peruvian current flows northwestward by the side of Peru, causing intense upwelling along the coast. While the main current moves away westward by the side of the Galapagos island, a branch of the current turns northward and then eastward towards the coast of Ecuador and adjoining southern Colombia, where it also meets the warm equatorial countercurrent flowing eastward towards north Ecuador and most of Colombia. Together, the ocean currents exert great influence upon the climates of the coastal states of Chile, Peru, Ecuador and Colombia.
226
9
Monsoon over South America
Fig. 9.2 Mean topographic height (m) over South America at 2.5◦ resolution heights above 500 m are shaded (Courtesy: NCEP/NCAR Reanalysis)
In the Atlantic Ocean, the tradewinds diverging from the subtropical anticyclones of the two hemispheres converge into the equatorial trough of low pressure. South of the equator, the southeasterly tradewinds turn anticlockwise to flow southward along the coast of Brazil and then southeastward over the southwestern Atlantic Ocean. Figures 9.3 and 9.4 shows the distributions of sea surface temperature (SST) and 10 m wind above the sea surface during January and July Figures 9.3 and 9.4 testify how the tradewinds diverging from the oceanic anticyclones carry cold airmass towards the equatorial trough of low pressure to participate in the monsoon circulation over both land and ocean.
9.3
Climatological Features
227
Fig. 9.3 Mean sea surface temperature (SST) and 10 m-wind above ocean surface over oceans around South America during January
9.3 Climatological Features 9.3.1 Air Temperature and Pressure Figures 9.5 and 9.6 shows the distributions of mean January air temperature at 925 and 300 hPa respectively.
228
9
Monsoon over South America
Fig. 9.4 Mean sea surface temperature (SST) and 10 m-wind above ocean surface over oceans around South America during July
Figures 9.5 and 9.6 are of great interest. They show that there is considerable zonal and meridional variation in the distribution of temperature maxima and minima over South America at 925 hPa during January. A temperature maximum is located north of the equator over the Venezuela region. South of the equator, the main temperature maximum appears to be centered over Bolivia and adjoining areas of northern Argentina, Paraguay and southern Brazil, over the latitude belt, 15–25◦ S. From the center, the temperature maximum appears to extend into three directions, one north-north-westward along the eastern slopes and foothills of northern Andes;
9.3
Climatological Features
229
Fig. 9.5 Mean January temperature (◦ C) at 925 hPa over South America and surrounding regions. Thick dashed lines mark the locations of low pressure troughs
the second southward over northern Argentina and adjoining southwestern Brazil; and the third over northeastern Brazil. A southeastward extension of the temperature maximum over Brazil towards Southwest Atlantic Ocean may also be seen. Indeed, the presence of the temperature maxima and minima as stated above suggests the presence of a quasi-stationary wave in temperature and pressure fields. Further discussion regarding the structure and properties of the quasi-stationary waves is deferred to a later section in this chapter. The 300 hPa temperature field (Fig. 9.6) along with the geopotential field (not shown) shows a well-defined ‘warm high’ centered over Central South America, above the low level heat low. A prominent ridge of the warm high appears to extend northeastward and then along the equator.
230
9
Monsoon over South America
Fig. 9.6 Mean January temperature (◦ C) at 300 hPa over South America and surrounding regions. Thick dashed lines mark the locations of low pressure troughs
9.3.2 Atmospheric Circulation – Monsoon Since South America is located across the tradewind belt, its atmospheric circulation is dominated by the tradewinds which often hide the monsoon circulation which appears in it as a perturbation only. So, for viewing the monsoon circulation, the background annual mean tradewind circulation must be removed from the observed circulation, as given by the relationships:
9.3
Climatological Features
231
V(d) = V(o) − V(a) where V is the wind vector, and the bracketed letters o, a, and d denote observed, annual mean, and deviation respectively. So, if there is a monsoon circulation over South America, it would be represented by the field of V(d). Saha and Saha (loc. cit.) used this procedure to compute the V(d) fields from the observed wind fields. An example from their study is presented in Figs. 9.7 and 9.8 for the 850 hPa climatological wind field for January, showing in the observed wind field, and in the deviation wind field.
Fig. 9.7 Climatological wind field at 850 hPa over South America and surrounding areas in January
232
9
Monsoon over South America
Fig. 9.8 The deviation wind field, V(d), at 850 hPa over South America and surrounding areas in January
The difference between the two wind fields of Figs. 9.7 and 9.8 is revealing. In Fig. 9.7, the circulation around the subtropical high pressure over the South Atlantic Ocean penetrates deep into Central South America, camouflaging that due to the heat low over the continent. All that can be seen of the heat low circulation is a feeble trough over Central Brazil between about 10 and 20◦ S. On the other hand, the structure of the monsoon circulation around the ‘heat low’ over Central South America at 850 hPa stands out in the deviation wind field in Fig. 9.8. The tradewinds of the two hemispheres appear to converge into the cyclonic circulation forming the ITCZ on the equatorial side and TCZ on the poleward side of the trough of the heat low.
9.3
Climatological Features
233
Henceforth, the deviation wind fields are used in discussing the structure and properties of the monsoon circulation over South America, unless otherwise stated.
9.3.3 Co-existence of Monsoon and Hadley Circulations In order to demonstrate that the monsoon circulation over South America co-exists with the Hadley circulation, as elsewhere in monsoon regions, Saha and Saha (2004a) used the reanalysis zonally-averaged values of v (the meridional component of the deviation wind) and ω (the vertical p-velocity, suitably scaled), to compute resultant streamlines which are shown in Figs. 9.9 and 9.10 for January and July. The zonal averaging was done between 65 and 45◦ W.
Fig. 9.9 Mean meridional circulations shown by resultant streamlines derived from zonally averaged (between 65 and 45◦ W) values of the v and ω components of the deviation wind over South America during July
234
9
Monsoon over South America
Fig. 9.10 Mean meridional circulations shown by resultant streamlines derived from zonally averaged (between 65 and 45◦ W) values of the v and ω components of the deviation wind over South America during January
The main features of the meridional circulations revealed by Figs. 9.9 and 9.10 are as follows. 9.3.3.1 Summer (January) A deep stream of descending northerly winds representing the lower branch of the Hadley circulation cell of the northern hemisphere appears to cross the equator and converge into the monsoon trough zone over central South America located approximately around 20◦ S (Fig. 9.9). However, while the bulk of the middle and upper troposphere appears to descend steeply north of the equator and rise to the south, the lower troposphere below about 800 hPa appears to execute a remarkable vertical oscillatory movement before it converges into the monsoon trough zone. It rises approaching the equator and sinks after crossing it. The rising branch of the Hadley circulation cell of the northern hemisphere appears to merge with the rising branch of the Monsoon circulation cell of the
9.3
Climatological Features
235
southern hemisphere over the broad latitudinal belt, 20–30◦ S, and then diverge in the upper troposphere to return to the northern hemisphere, which the poleward-moving branch of the Monsoon circulation cell descend further south where it appears to get merged with the descending branch of the meridional circulation associated with the polar trough zone.
9.3.3.2 Winter (July) The circulation over South America in winter appears to be totally reverse of that in summer (Fig. 9.10). The monsoon convection zone is now in the northern hemisphere. Strong descending motion occurs over central South America and diverging southerly winds from this area of subsidence appear to cross the equator at low levels and converge into the monsoon trough zone between about 10 and 20◦ N. Thus, the common rising branch of the Hadley circulations of the two hemispheres is located north of the equator and the winter monsoon over South America (south of the equator) is characterized by wholesale subsidence of air.
9.3.4 Rainfall over South America Figures 9.11 and 9.12 shows the distribution of climatological rainfall over South America and neighboring areas during January and July respectively. In summer, heavy rain with concentrations between 9 and 11 mm day–1 falls over the large catchment areas of River Amazon in Brazil. Comparable heavy rain also appears to fall over equatorial eastern Atlantic Ocean a few degrees north of the equator. The Southwest Atlantic Convergence Zone is another area marked by heavy rain. In the Pacific Ocean, the rain belt appears north of the equator along the ITCZ. By contrast, little rain falls over extreme NE Brazil and Venezuela as well as in the Patagonia region of Argentina where heat sources are located. The subtropical oceans appear to be mostly dry. In winter (Fig. 9.12), the whole rainbelt appears to have shifted northwestward with heavy rain mostly concentrated in the northern hemisphere. In South America, although the northern part of the Amazon River valley continues to receive rainfall, the heaviest falls occur over areas further north over Venezuela, Colombia and adjoining Central America. By contrast, northeast Brazil appears to be almost dry, except a narrow coastal belt south of Cape Roque where rainfall upto 4 mm day–1 may be expected. The oceanic area off the east coast of Southern Brazil also experiences heavy rainfall. The ITCZ over the eastern Pacific Ocean, especially that in the vicinity of Central America, appears to be very active with heavy rain exceeding 12 mm day–1 occurring over a large area. The belt of heavy rain over the equatorial eastern Atlantic appears to be located a few degrees north of the equator.
236
9
Monsoon over South America
Fig. 9.11 Climatological (1979–2003) mean rainfall (mm day–1 ) over South America and surrounding areas during January (Xie and Arkin, 1996)
9.4 Quasi-stationary Waves and Their Associated Weather There are several aspects of the summer monsoon weather and climate of South America which appear to be related to two quasi-stationary waves which are forced by land-sea thermal contrast and/or orography along the northern and the southern
9.4
Quasi-stationary Waves and Their Associated Weather
237
Fig. 9.12 Climatological (1979–2003) mean rainfall (mm day–1 ) over South America and surrounding areas during July (Xie and Arkin, 1996)
boundaries of the heat low over the continent. We may call these stationary waves as follows: (a) The monsoon stationary wave with the ITCZ forming a part of it along the northern boundary, and (b) The subtropical stationary wave with the TCZ forming a part of it along the southern boundary.
238
9
Monsoon over South America
9.4.1 Weather Phenomena Related to the Northern Boundary (a) A Low-level Jet (LLJ) near the foothills of the northern Andes, and (b) Drought over NE Brazil
9.4.1.1 Low Level Jet (LLJ) This jet often speeding to 20/25 m s–1 in the observed wind field is a long narrow band of northerly wind that blows along the western boundary of Brazil within the shadow of the northern Andes. It forms when under influence of the heat lows on both sides of the equator in South America, the NE tradewinds diverging from the subtropical high pressure belt of the northern hemisphere cross the equator and enter deep into South America. A branch of the tradewinds which is diverted westward to blow around a heat low over the landstrip (Venezuela and the Guyanas) north of the equator also crosses the equator separately. South of the equator, the two branches converge into each other to form the mainstream tradewinds which blow southward. However, the converging currents after traversing long distance eventually diverge into two branches under the influence of heat lows that lie to west and east. The branch moving westward converges into the cyclonic circulation around the elongated heat low near the foothills of the Andes but that moving eastward has first to turn anticyclonically around the ridge of a cold high pressure formed by the tradewinds of the northern hemisphere in order to converge into the cyclonic circulation around the summer heat low over eastern Brazil. This wedge of high pressure plays an important role in sustaining the Low Level Jet (LLJ) over South America. The differential heating between this cold wedge and the heat low to the west appears to accelerate the airflow between 10 and 15◦ S and decelerate it south of about 15◦ S. In satellite cloud imagery, the location of the LLJ over South America is marked by an elongated band of well-developed clouds. Heavy precipitation is known to occur from these clouds over the catchment areas of R. Amazon. The satellite cloud imagery also shows a prominent narrow cloud band along the foothills of the Andes which, according to Paegle and Nogues-Paegle (1997), may be largely due to the frictional effect of the mountainside on the aircurrents converging at it. In several ways, the LLJ over South America may be looked upon as a crossequatorial jet ushering in the summer monsoon current to South America, as in many other parts of the globe. It may be compared with the Somali jet in the Indian Ocean, or the Indonesian jet in the Asia–Australia region (see Fig. 4.9). In all cases, it is the differential heating between a heat source (usually land) and a heat sink (usually ocean) that appears to drive the jets. In the cases of Somali jet, West Australian jet, Benguela jet, and the Peruvian jet, coastal upwelling further strengthens the jets. Figure 9.13 illustrates the formation of the Low Level Jet (LLJ) (thick continuous line) over South America. It also shows the locations of the monsoon trough (thick double-dashed), the ITCZ and the TCZ over South America in relation to the LLJ.
9.4
Quasi-stationary Waves and Their Associated Weather
239
Fig. 9.13 Illustrating the formation of the Low Level Jet (LLJ) (indicated by thick line with arrow) over South America. L denotes low pressure, H – high; dotted areas mark the approximate locations of the ITCZ and TCZ. Double-dashed lines indicate troughs, single-dashed ridges
9.4.1.2 Zonal Anomaly of Rainfall over NE Brazil In a pioneering study on rainfall over NE Brazil, Namias (1972) had noted that an increased cyclonic activity over the extratropical belt of the North Atlantic ocean off the coast of North America had the effect of intensifying the tropical Hadley circulation and hence increased rainfall over NE Brazil where the low level moisture-bearing NE tradewinds converge. Riehl (1977) who studied rainfall over Venezuela also found the same effect and proposed a similar mechanism linking the strength of the northern hemisphere Hadley circulation with the variation of rainfall over South America. Hastenrath and Heller (1977) suggested that a strengthening of the subtropical high pressure of the North Atlantic Ocean would displace the equatorial trough of low pressure southward and account for increased rainfall activity over northeast Brazil. A few studies (e.g., Maura and Shukla, 1981; Nobre and Shukla, 1996) have sought to relate the variability of monsoon rainfall over NE Brazil to anomalies of sea surface temperature (SST) over the equatorial Atlantic ocean. According to these studies, the development of a warm SST anomaly to the north of the equator and a cold SST anomaly to the south would have the effect of shifting the equatorial trough of low pressure and its associated rainbelt northward, thereby reducing rainfall over South America. While the SST factor remains a distinct possibility, the main cause may lie nearer home in the circulation field associated with the trough of the ‘heat low’ over NE Brazil. Rainfall maps presented in Figs. 9.11 and 9.12 testify that during both summer and winter, there is much less rainfall over northeast Brazil (east of about 40◦ W)
240
9
Monsoon over South America
than areas to the west. For example, in January, the mean rainfall over the Amazon delta near Belem in the western part of NE Brazil is about 320 mm, whereas that over the eastern part between Fortaleza (Ceara) and cape Roque hardly amounts to 60 mm (see Fig. 9.11). The disparity stands out in winter. What is the cause of this great disparity in rainfall between the two parts? In this context, it is important to recognize the real character of the monsoon circulation in relation to the heat low circulation over NE Brazil. The heat low as a heat source is flanked by two heat sinks, one over the North Atlantic Ocean and the other over the Southwest Atlantic Ocean. In this structure of the monsoon circulation, heavy rain falls along the ITCZ to the west of the trough and only light rain along the TCZ to the east. The trough zone itself appears as a rainfall deficiency area. In Fig. 9.13, the trough of the monsoon heat low is indicated by a thick double-dashed line, while the ITCZ and the TCZ are denoted by thick continuous lines. Thus, the evidence is overwhelming that the perennial drought conditions over NE Brazil are created largely by the monsoon circulation system itself. Of course, periodical variations do occur due to external influences such as, anomalies of sea surface temperature, movement of wave disturbances, and shift in the locations of the subtropical high pressure belts of the two hemispheres.
9.4.2 Weather Phenomena Associated with the Southern Boundary 9.4.2.1 Southwest Atlantic Convergence Zone (SACZ) A trough of the ‘heat low’ over SE Brazil at low levels extends southeastward into southwestern Atlantic Ocean. Further southeastward from the Brazil coast over the ocean, this trough of the ‘heat low’ appears to form an extended trough with a quasistationary trough in the midlatitude westerlies of the southwestern Atlantic ocean where the poleward – flowing warm Brazil coastal current forming the western boundary current of the South Atlantic subtropical anticyclonic ocean gyre meets the equatorward-flowing cold Malvinas current. This is the trough of the so-called Southwest Atlantic Convergence Zone (SACZ), sandwiched between two oceanic high pressures of the South Atlantic Ocean. In the SACZ, there is warm air convergence at low levels with divergence above to the east of the trough axis, while there is divergence at low levels with convergence above to the west. This leads to a secondary zonal vertical circulation around the trough axis with penetrative convection to the east and strong subsidence to the west. According to Satyamurti et al. (1998), the SACZ is a strongly convergent zone with heavy rainfall. But this convergence and rainfall occurs distinctly to the east of the trough axis. Its presence as an elongated cloud band, broad and bright, in conjunction with that along the east coast of Brazil extending deep into the South Atlantic in a NNW–SSE orientation is revealed prominently in satellite cloud imagery. The SACZ frequently interacts with eastward-propagating wave disturbances in midlatitude baroclinic westerlies. During such interaction, an extended trough of low pressure may form between the two troughs when they get into the same phase.
9.5
Tropical Disturbances over South America
241
However, the interaction may not last long, since the extended trough gets ruptured when the trough of the midlatitude disturbance moves away eastward. 9.4.2.2 Semi-arid Conditions over Patagonia Since the warmest part of the ‘heat low’ over central South America tends to lie over the southwestern part of Brazil and adjoining region of Bolivia and northern Argentina, a trough of low pressure extends southward from the center of the heat low across Argentina. Further, Patagonia lies in the rain shadow of the Southern Andes and midlatitude disturbances passing over the region produce little rain. So the region suffers from lack of rain and persistent drought conditions.
9.5 Tropical Disturbances over South America 9.5.1 Types of Disturbances Since the pioneering studies of Ramos (1975), Virji (1981), Kousky and Gan (1981) and others, it has been known that quite often during southern summer, an uppertropospheric cyclonic vortex forms over the Southwest Atlantic ocean, close to the coast of northeastern Brazil, which moves inland (westward) carrying with it a frontlike zone of cloudy and rainy weather to the northeastern part of Brazil. Kousky and Gan (1981) in particular carried out a detailed study of this type of disturbances and threw light on their structure and properties. A recent study (Saha and Saha, 2004b), using observed winds, identified two types of tropical disturbances which may form in the circulation around the heat low over South America. These are: (1) Monsoon lows and depressions over the continent; and (2) Upper-tropospheric cyclonic vortices over nearby Atlantic Ocean off the coast of northeastern Brazil During January of a period of 4 years, 1999–2002, Saha and Saha (loc. cit.) detected as many as nine synoptic-scale tropical disturbances over South America, the tracks of which are shown in Fig. 9.14. Of the disturbances observed, four were monsoon lows and depressions over the southeastern part of Brazil, and five upper-tropospheric cyclonic vortices over the oceanic area close to the coast of NE Brazil. Saha and Saha (loc. cit.) noted several common features amongst the disturbances of each type. Since it is not practicable to give full details of each of these disturbances in the present text, attempt is made to give here some details in respect of one typical monsoon low/depression and one upper-tropospheric cyclonic vortex, as examples only.
242
9
Monsoon over South America
Fig. 9.14 Tracks of monsoon lows and depressions (crosses) and upper-tropospheric cyclonic vortices (open circles). Dates are given against every location, with the year (last two digits) at first location only
9.5.2 Monsoon Lows and Depressions In the Southern hemisphere, eastward-propagating extratropical wave disturbances of the Southern Pacific Ocean, after crossing the Southern Andes, often interact with the heat low circulation over South America. However, not all of them produce any significant impact. Only those with large amplitudes and extending equatorward to at least 20◦ S, or closer, are found to interact with the quasi-stationary wave. Generally, the traveling wave interacts with the quasi-stationary trough of low pressure along the southern boundary of the heat low over the continent, forming an extended trough which moves eastward. The formation of the extended trough leads to an amplification of the waves. The southerly winds to the west of the extended trough rush equatorward advecting cold polar air to lower latitudes, while the northerly winds to the east move in the opposite direction advecting warm tropical air to higher latitudes. Occasionally during such interaction, a monsoon disturbance may form in the extended trough as a cut-off low or depression. Indeed, Saha and Saha (loc. cit.) detected quite a few monsoon lows and depressions that formed in this way. After formation, the cut-off low moves westward/northwestward. In what follows, we give some details of a cut-off low of this kind which formed over Central South America on 7 January 2002.
9.5
Tropical Disturbances over South America
243
9.5.2.1 Case Study: Monsoon Low/Depression of 7–11 January 2002 This was a depression which formed in a trough of low pressure in the quasistationary subtropical wave when an eastward-propagating midlatitude wave in the baroclinic westerlies interacted with it, forming an extended trough of low pressure along around longitude 50◦ W and latitude 20◦ S at 12 GMT 7 January 2002. It first formed as a cut-off low, and later developed into a depression and moved northwestward. The NCEP Global Data Analysis System (GDAS) throws considerable light on the genesis of this depression. The fields of geopotential, temperature, wind and specific humidity relating to the disturbance at 12 GMT on 7 January 2002 at 925 and 300 hPa are shown in Figs. 9.15 and 9.16 respectively. Further evolution and movement of the depression ‘D’ during the period 8 through 11 January 2002 is shown in Fig. 9.17(a–d).
9.5.3 Upper-Tropospheric Cyclonic Vortices These are synoptic-scale cyclonic disturbances which form in the upper troposphere as cut-off ‘cold lows’ over the oceanic area close to the northeastern coast of Brazil during southern summer and move generally westward. According to Kousky and Gan (1981), the vortices form as the downstream effect of divergence from an intensified upper-tropospheric warm ridge of high pressure over South America as a result of its interaction with an eastward-propagating baroclinic wave disturbance in the midlatitude westerlies of the southern hemisphere. While the eastward-propagating wave disturbances in the midlatitude baroclinic westerlies of the southern hemisphere no doubt interact with the summertime quasistationary subtropical wave associated with the ‘heat low’ over South America, the interaction does not always lead to the formation of an upper-tropospheric cyclonic vortex. More often than not, it is an eastward-propagating wave disturbance in the midlatitude baroclinic westerlies of the northern hemisphere which interacts with the quasi-stationary monsoon wave of the southern hemisphere that leads to the genesis of an upper-tropospheric cyclonic vortex off the coast of NE Brazil. It is well-known that during northern winter, the midlatitude baroclinic westerlies blow over latitudes close to the equator. It is proposed to demonstrate the validity of the mechanism suggested in the preceding para through a case study. The vortex selected for the purpose appeared first at 300 hPa near the coast of NE Brazil on 1 January 2001 and moved inland in a mean westsouthwestward track and could be followed over the following 6 days (the track is shown in Fig. 9.14). 9.5.3.1 Case Study: Upper-Tropospheric Cyclonic Vortex of 1–6 January 2001 The synoptic situation relating to this disturbance is presented in Fig. 9.18 in which (a) shows the distribution of msl pressure, (b) flow patterns at 300 hPa, and (c) the zonal-vertical circulation through the center of the vortex (V).
244
9
Monsoon over South America
Fig. 9.15 Fields of geopotential (gpm), temperature (◦ C), wind (m s–1 ), and specific humidity (g/kg) at 925 hPa at 12Z, 07 January 2002. D marks the location of the depression
The maps in Fig. 9.18(a–c) would appear to throw light on several aspects of the formation, structure and movement of the upper-tropospheric cyclonic vortex. The mean sea level pressure (MSLP) distribution in (a) reveals the presence of a well-marked W’ly wave trough along about longitude 50◦ W in the subtropical high pressure belt of the northern hemisphere. The equatorial belt on both sides of the
9.5
Tropical Disturbances over South America
245
Fig. 9.16 Fields of geopotential (gpm), temperature (◦ C), wind (m s–1 ), and specific humidity (g/kg) at 300 hPa at 12Z, 07 January 2002. D marks the location of the depression
equator east of this longitude appears to be in a low pressure zone. In the southern hemisphere, a deep trough of low pressure flanked by two intense high pressure cells appears to extend southeastward from the low pressure area over Southeastern Brazil into the Southwest Atlantic Ocean. This is the extended trough between the quasi-stationary low pressure over Southeastern Brazil and a deep trough of low pressure of a traveling southern hemisphere midlatitude westerly wave.
246
9
Monsoon over South America
Fig. 9.17 The wind field at 925 hPa over South America and neighboring areas at 12 GMT daily, 8 through 11 January 2002. D marks the depression center
The 300 hPa flow pattern in Fig. 9.18(b) shows the interaction between the circulation fields of the quasi-stationary wave of South America with a traveling baroclinic wave disturbance of the northern hemisphere. It is easy to see that it is this interaction between the two hemispheric waves which appears to have contributed to the generation of the upper-tropospheric cold cyclonic vortex (V) by convergence into it of air diverging from two warm highs, one to its east and the other to the west. After formation, the cut-off vortex moved westward. It is interesting to see that while the above-mentioned developments took place leading to the formation of the vortex, a baroclinic wave disturbance in midlatitude
9.5
Tropical Disturbances over South America
(a)
247
(b)
(c)
Fig. 9.18 (a) MSLP (hPa), and (b) flow patterns at 300 hPa over South America and neighboring oceans, and (c) zonal-vertical circulation through the center of the vortex (V) at 12 GMT, 1 January 2001. L denotes low pressure, H – high pressure
westerlies of the southern hemisphere was moving across South Atlantic, but it just moved away eastward without interacting with the vortex. Figure 9.18(c) illustrates the structure of the vortex in the upper troposphere as a cold convergent area between two divergent warm highs on either side, one being the quasi-stationary Bolivia-Brazil high and the other the equatorial high. A similar structure was earlier suggested by Kousky and Gan (loc. cit.). According to current view, as cold air descends inside the vortex center, there is adiabatic warming of air in the middle of the troposphere with divergence of cold air below. Where the diverging cold air converges into the warm air to the east and the west, there is strong penetrative convection, clouding and precipitation. In the following section, we discuss the formation of an upper-tropospheric cyclonic vortex which formed over the Southwest Atlantic Ocean as a result of interaction of the quasi-stationary wave of South America with eastwardpropagating baroclinic wave disturbances of both the hemispheres, and developed later into a tropical cyclone.
248
9
Monsoon over South America
9.6 A Tropical Cyclone over the South Atlantic Ocean 9.6.1 Introduction The formation of a tropical cyclone over the South Atlantic Ocean in March 2004 that hit Brazil on 28 March was an event that took the meteorological community all over the world by surprise, since in the past there had been no evidence of a tropical cyclone over this ocean and there had been a general belief that no tropical cyclone could ever form over this ocean, since it was too cold to breed and sustain a tropical cyclone. The belief was not entirely unfounded, for Gray (1968) had shown (see Fig. 2.8) that during a 20-year period no tropical cyclone had ever been reported from the South Atlantic Ocean. Unfortunately, there were no conventional meteorological data in the form of surface and upper-air temperature, height, and wind observations available within a few 100 km of where the March 2004 cyclone formed to enable a thorough diagnostic analysis of the situation. The only data available were those from a few coastal stations of Southeastern Brazil (nearly a thousand kilometers away from the site of the first detection of the storm), and daily satellite cloud imagery, and NCEP daily data analyses. Fortunately, we had daily Geostationary Operational Environmental Satellite (GOES)-12 reports of cloud formations and mid-upper level winds derived from cloud motion vectors to aid analysis and interpretation. Thus, the Brazil cyclone posed a real challenge to meteorologists. There are several questions regarding the origin, structure, development and movement of this cyclone that need to be addressed. Little definite can be said about them till we have the ground truth from a viable synoptic or subsynoptic-scale surface and upper-air data network over the ocean which are currently unavailable and may never be available in the future. So, anything that may be said now about this cyclone must be of a speculative and tentative nature and subject to future review. However, a preliminary analysis and study of the available data, especially SST, cloud formations and satellite-derived upper-air temperature and wind data and analyses, as available through the courtesy of NCEP/NCAR and the University of Wisconsin Satellite Facility, appears to provide a clue to a possible physical mechanism that might have been involved in the genesis of this cyclone vortex. In what follows, we state the result of our analysis of this cyclone. Unfortunately, the exact date and place of its first formation over the ocean is not known, but it was first sighted by satellie on 23 March 2004 as a westward-moving depression at a location near 29.5◦ S, 39.5◦ W over the Southwest Atlantic. Its subsequent track and intensity, as estimated by the University of Wisconsin Advanced Microwave Sensing Unit (AMSU), are given in Fig. 9.19. A satellite view of the disturbance at 18 UTC, 25 March 2004, when it developed into a cyclone with eye at the center, with the cloud field and distribution of mid upper-level winds associated with it, is shown in Fig. 9.20
9.6
A Tropical Cyclone over the South Atlantic Ocean
249
Fig. 9.19 Track and intensity of the Brazil tropical cyclone, showing the daily location of its center (with minimum pressure in bracket) during the period, 23 through 28 March 2004. Open circle signifies low/depression, spiral cyclone with an ‘eye’
Fig. 9.20 A satellite viewof the Brazil cyclone, showing the distribution of clouds (white areas) and GOES-12 mid-upper level winds in colours (blue between 100 and 250 mb, yellow between 251 and 350 mb, and green between 351 and 500 mb) at 18 UTC on 25 March 2004
250
9
Monsoon over South America
9.6.2 Formation of the Initial Vortex – Interaction with W’ly Waves The first step in the investigation was to find out if there was any involvement of the quasi-stationary wave over the continent with any traveling waves in the formation of the initial vortex. For an answer, we noted the daily locations of the W’ly troughs of the two hemispheres at 300 hPa during a period of 8 days prior to 23 March 2004 when it was first detected. These locations are shown in Fig. 9.21 An examination of the daily locations of the troughs at 300 hPa reveals that on or around 21 March, the W’ly waves of the two hemispheres got into a phase which allowed cross-equatorial movements (with change of sign of pressure systems across the equator) were facilitated and which led to the coupling of the circulation system of the quasi-stationary wave over South America and neighboring Atlantic Ocean with those of the traveling waves of the two hemispheres. The two-way interaction led to the formation of an extended trough of low pressure (indicated by a thick double-dashed line) in the southern hemisphere and it is speculated that the nucleus of the cyclone vortex formed in the extended trough so formed during the period of coupling as a cut-off low (indicated by letter V) at a location near the poleward boundary of the tropics (but still in the tropics) (see Fig. 9.22). The important point to note here is that the upper-tropospheric cyclonic vortex (V) in this case formed not near the northeastern coast of Brazil, but at a much
Fig. 9.21 Locations of the extratropical wave troughs of the southern hemisphere (double-dashed line) and the northern hemisphere (dotted) in the flow field at 300 hPa over South America and neighboring ocean area at 12 UTC daily during the period, 16 through 23 March 2004. Dates are entered at the end of each line. The location of the vortex first sighted by satellite on 23 March is indicated by a circle
9.6
A Tropical Cyclone over the South Atlantic Ocean
251
Fig. 9.22 Schematic illustrating the suggested mechanism of formation of the initial upper-air cyclonic vortex (V) in the extended trough (thick dashed line). H – high, L – low thin lines are streamlines with arrow showing direction of air motion
higher latitude near the poleward boundary of the tropics where warm air diverging from the two high pressure areas, one over South America and the other over mid-Atlantic Ocean converged to form the vortex in the extended trough formed with the southern-hemispheric W’ly wave After formation of the vortex, the midlatitude W’ly waves moved away eastward, leaving the vortex to move westward as a tropical depression.
9.6.3 Structure, Movement and Development of the Vortex The structure of an upper-tropospheric cyclonic vortex over the Southwest Atlantic Ocean was discussed earlier in Sect. 9.5 (see Fig. 9.18c). According to it, the vortex is characterized by cold air convergence at the top, subsidence at the center, and divergence at the bottom. The vortex is flanked by two circulation systems which have warm air divergence above low level convergence. It is believed that the vortex retained this structure while it moved westward as a depression. However, there must have been a major structural change in the vortex on or after 25 March when it turned into a tropical cyclone with an eye at the center. The process appears to have been aided by the distinctly warmer waters of the Brazil
252
9
Monsoon over South America
Fig. 9.23 SST (C) and wind (m s–1 ) at 10 m above surface at 00Z 24 March 2004. The center of the vortex is indicated by the letter ‘V’
coastal current, as would be evident from Fig. 9.23 which shows the distribution of SST and wind at 10 m-above the ocean surface at 00Z 24 March 2004, a few hours prior to the development of the vortex. It shows strong warm air advection from the north where the temperatures are as high as 27–28◦ C in the eastern side of the vortex and cold air advection from the south where the temperatures are as low as 22–23◦ C in the western side. It is likely that continued subsidence warming of the vortex and sensible and latent heat influx from the ocean surface might have led to its transition from a coldcore depression to a warm-core tropical cyclone. Undoubtedly, this crucial structural transition must have taken quite a few days to materialize. According to NCEP analysis of the thermal field at 12 UTC on 27 March when the cyclone was within about 300 km of the coast, it had developed an intense warm core at 300 and 200 hPa. The warm-core structure continued till the cyclone weakened near the Southeastern coast of Brazil and dissipated after crossing the coast on 28 March.
Part III
Extratropical Monsoons
Chapter 10
Monsoon over Central America and Adjoining Southwestern North America (Region – VII)
10.1 Introduction Although Central America is located over the tropical belt of the northern hemisphere, its monsoon extends to extratropical latitudes. It forms an important link between the two American continents (see Fig. 10.1). Central America is widest in the north and gradually narrows down to the State of Panama in the South. Mexico is the largest State in the north which occupies nearly half of the region and has to its south the States of Guatemala, Honduras, Belize, El Salvador, Nicaragua, Costa Rica and Panama. The region is oriented in a NW–SE direction. Its oceanic environment consists of the Pacific Ocean to the west and the Atlantic Ocean, the Gulf of Mexico, and the Caribbean Sea to the east. Two ocean currents importantly affect the climate of Central America. These are the equatorwardflowing cold California Current of the Eastern Pacific Ocean which flows by the side of California and Baja California, and the poleward-flowing Gulf Stream of the Western Atlantic Ocean which transports warm water by the side of the East Coast of North America. The topography of the region plays an important role in determining the circulation and climate of the region. The region has off its northeastern boundary the great Rocky Mountains of North America, Nearly half of the State of Mexico is an elevated plateau of average height of about 1.5 km a.s.l., which has on its sides two prominent northwest–southeast oriented mountain ranges, viz., the Sierra Madre Oriental in the east which rises from the Plains of the west coast of the Gulf of Mexico, and the Sierra Madre Occidental which rises from the Plains of the east coast of the Gulf of California and the Pacific Ocean, in successive steps to the peak of the mountain ranges. The plateau level rises southward from the US–Mexico border where it is at a little less than the mean height to attain a maximum altitude of about 2.2 km at the Mesa de Anahuac. At the southern edge of the plateau lies a belt of volcanic mountains known as Sierra Volcanica Transversal that extends west to east and rises to great heights at places before dropping precipitously to very low levels. There are also several other mountain ranges to the south of Mexico, some low mountains on either side of the Panama
K. Saha, Tropical Circulation Systems and Monsoons, C Springer-Verlag Berlin Heidelberg 2010 DOI 10.1007/978-3-642-03373-5_10,
255
256
10
Monsoon over Central America and Adjoining Southwestern North America
Fig. 10.1 Physical map of Central America and its oceanic environment
Canal, till the topography rises to the high Andes mountains of Colombia in South America.
10.2 Heat Sources and Sinks and Their Seasonal Movement We have seen in Chap. 9 that during northern winter the equatorial trough of low pressure is located over Central South America. So, it follows that the equatorial trough has to travel all the way from South America to Central America, a distance spanning more than 45◦ of latitude across the narrow Panama Strait in the run-up to summer. This movement takes place gradually with the help of a series of heat lows that develop over the land to the north while the heat low in the south winds up with change of season. The movement generally takes place along the land route via the Panama Canal Zone. It is the gradient of pressure tendency between the two sides that appears to force the movement. At the final stage, the moving heat low which brings up the monsoon circulation to Central America merges with the summer heat low over northwestern Mexico. For this reason, some meteorologists (e.g., Mosino Aleman and Garcia, 1974; Douglas et al., 1993) have called it the ‘Mexican monsoon’. Other terms such as ‘Arizona monsoon’ have also been used. For the Central American monsoon, the Eastern Pacific Ocean serves as the main heat sink from which the tradewinds converge into the heat low over the region,
10.3
The Climate of Central America and Adjoining North America
257
though, occasionally, the tradewinds from the Atlantic Ocean also find their way to the region.
10.3 The Climate of Central America and Adjoining North America In the past, the climate of Mexico and adjoining North America has been studied by several workers (e.g., Bryson and Hare, 1974; Portig, 1976; Mosino Aleman and Garcia, 1974; Douglas et al., 1993). These studies have considerably improved our knowledge and understanding of the monsoons over Central America and adjoining ocean areas. In the present section, we discuss the climate of the region on the basis of a series of surface and upper air climatological maps from NCEP Reanalysis for both Central and North America, extending from 10 to 70◦ N. Besides surface maps, we use maps at 850 and 300 hPa for upper-air climatology.
10.3.1 Surface Temperatures and Winds NCEP Reanalysis presents the distribution of the surface temperature along with wind at 10 m above surface over both land and ocean over Central America and adjoining North America and surrounding oceans in Fig. 10.2, for January and for July. In general, the January surface temperatures (SST) and winds over the land and ocean would appear to show the following features: The land surface temperatures are generally low compared to the SSTs at the same latitude and appear to decrease continuously with latitude. However, topographic effects on the distribution of surface temperatures are quite apparent. Much lower temperatures appear over the Plateau surfaces and Mountain tops. Relatively warmer temperatures appear over the Great Basin area of North America. Over Central America, the cold land surface is flanked by warm oceans on both sides of the land. In January, the winds that affect Central America appear to originate in two anticyclonic circulations, one centered near the southern coast of the Gulf States to the west of the Florida Peninsula and the other over the Subtropical belt over the Atlantic Ocean off the east coast of Florida. The NE-ly winds diverging from these anticyclones appear to cross Southern Mexico and converge into the ITCZ further south and also into the heat low over South America. In the Pacific Ocean, the warmest temperatures appear south of about 15◦ N with values of SSTs higher than 27.0◦ C at places. The temperatures decrease steadily poleward, reaching a value of about 6◦ C near Alaska. The dominant feature of the wind field over the Pacific Ocean is a well-defined anticyclonic circulation over the subtropical belt centered at about 30◦ N, 130◦ W. Winds diverge from this center both equatorward and poleward. The equatorward-diverging winds blowing by the side of California converge into the Intertropical Convergence Zone (ITCZ) near the
258
10
Monsoon over Central America and Adjoining Southwestern North America
Fig. 10.2 Climatological (1979–1996) surface temperature (◦ C) and winds (m s–1 ) at 10 m above surface over the region of Central and North America and surrounding oceans during January
equator. They transport cold, humid air equatorward. In the Gulf of Mexico and the Caribbean Sea also, the warmest temperatures lie south of about 15◦ N, with values decreasing northward. In July (see Fig. 10.3), the seasonal changes in temperature and wind over both land and ocean surfaces are quite apparent. Temperatures have risen all over, but the rise over the land has been so high as to reverse the temperature gradient between the land and the ocean. Two pronounced temperature maxima appear over the land, one over the Great Basin area of southwestern North America and adjoining Sonoran desert of northern Mexico, and the other over the Great Plains area of North America with the thermal ridge running north-north-westward. These temperature maxima are separated in the south by low temperatures on the Mexican Plateau and in the north by low temperatures on the Rocky Mountains. The anticyclonic circulation centers over both the Pacific and the Atlantic Oceans have shifted northward, that over the Pacific Ocean to about 40◦ N and that over the Atlantic Ocean to about 35◦ N. The winds diverging from these anticyclones converge into the heat low circulations over the land and the ITCZ over the oceans.
10.3
The Climate of Central America and Adjoining North America
259
Fig. 10.3 Climatological (1979–1996) surface temperature (◦ C) and winds (m s–1 ) at 10 m above surface over the region of Central and North America and surrounding oceans during July
10.3.2 Upper Air Temperatures Figures 10.4 and 10.5 show the temperatures over Central and North America at 850 and 300 hPa during January and July. While interpreting the distribution of temperature, or any other meteorological variable, at 850 hPa, over the region, it may be borne in mind that this pressure surface is close to the average height of the Mexican Plateau and the Western Plateau of North America. In January (Fig. 10.4), the warmest temperatures appear to be located over the southern parts of Central America and the northern parts of South America and adjoining ocean areas at both 850 and 300 hPa. At 850 hPa, a thermal ridge appears to run from the extreme southern part of Mexico to the northern part of South America. Temperatures appear to decrease generally with latitude over both land and ocean, but the gradient is steeper over the land than over the ocean, especially over the extratropical latitudes over the land. Lowest temperatures are to be found over Northern Canada and the Frozen Arctic Ocean at both the pressure surfaces in January.
260
10
Monsoon over Central America and Adjoining Southwestern North America
Fig. 10.4 January air temperatures at 850 and 300 hPa over Central America and adjoining North America
The July temperature field (Fig. 10.5) shows that at 850 hPa, the warmest temperatures are located over the northwestern part of Central America and adjoining Western Plateau of North America. A well-marked thermal ridge from this temperature maximum appears to extend northwestward towards Western Canada and Alaska. As at surface, the lowest temperatures appear over the Arctic Ocean. In general, the temperatures at the same latitude are now much higher over the land than over the ocean. The temperature field at 300 hPa shows almost the same features as that at the lower pressure surface.
10.3.3 Upper Air Height (gpm) The isobaric height fields at 850 and 300 hPa for January and July, shown in Figs. 10.6 and 10.7 respectively conform well to the corresponding temperature
10.3
The Climate of Central America and Adjoining North America
261
Fig. 10.5 July air temperatures at 850 and 300 hPa over Central America and adjoining North America
fields shown in Figs. 10.4 and 10.5, with height maxima appearing over the warmest areas and minima over coldest areas. In January (Fig. 10.6), height maxima at 850 hPa appear over the subtropical oceans on both sides of Central America with lower heights over the cold land. Heights fall uniformly with latitude towards the North Pole over both land and ocean. A well-marked height ridge appears to run along the Rocky Mountains area towards Alaska at both the pressure surfaces, especially during winter. The July height field (Fig. 10.7) shows a well-marked low over southwestern United States and adjoining Mexico between two powerful highs over the subtropical oceans at 850 hPa, and a reversal of the height field at 300 hPa.
262
10
Monsoon over Central America and Adjoining Southwestern North America
Fig. 10.6 January height field (gpm) at 850 and 300 hPa over Central America and adjoining North America
10.3.4 Upper Air Wind Field and Circulation The wind fields at 850 and 300 hPa over the region during January and July are shown in Figs. 10.8 and 10.9 respectively. The main features of the wind fields are briefly the following: In January, the wind field at 850 hPa shows the prominence of the subtropical anticyclonic circulations over the oceans and an extended trough of weak cyclonic circulation over Central America. The tradewinds diverging from the anticyclones after passing over the warm tropical ocean appear to converge into the near-equatorial ITCZ over the eastern Pacific Ocean.
10.3
The Climate of Central America and Adjoining North America
263
Fig. 10.7 July height field (gpm) at 850 and 300 hPa over Central America and adjoining North America
At 300 hPa, westerlies dominate the flow field over Central America. There appears to be a general tendency for these upper winds to converge on entering land and diverge after emerging over the ocean. In July, there is a well-marked cyclonic circulation at 850 hPa over the northwestern part of Mexico to which tradewinds diverging from the subtropical anticyclones appear to converge from both west and east. The cross-equatorial flow from the Southeastern Pacific Ocean also appears to converge into the ITCZ. At 300 hPa, a well-marked anticyclonic circulation with its axis running along about 27.5◦ N over Central America appears to be centered over the low-level cyclonic circulation over the northwestern part of Mexico. Strong easterlies appear to the south of the axis, with westerlies to the north. Strong westerlies prevail over the oceans on both sides of Central America. A well-marked trough appears in the
264
10
Monsoon over Central America and Adjoining Southwestern North America
Fig. 10.8 January wind fields and circulation at 925 and 300 hPa over Central America and surrounding areas
westerlies over Central Atlantic Ocean. Thus, tropical easterlies prevail in the upper troposphere over Central America during the summer monsoon season.
10.4 Rainfall over Central America and Adjoining Areas Figures 10.10 and 10.11 shows the distribution of mean monthly rainfall (precipitation rate, mm day–1 ) over the region: January and July (after Xie and Arkin, 1996).
10.4
Rainfall over Central America and Adjoining Areas
265
Fig. 10.9 July wind fields and circulation at 925 and 300 hPa over Central America and surrounding areas
In January, it is mostly the eastern parts of Central America, especially the Yucatan Peninsula of Mexico and the States further south that appears to get some good amount of rain. The rest of Central America is practically dry. In the Pacific Ocean, the subtropical high pressure zone appears to have little rain, but the areas to its north and south appear to be quite rainy. A rainfall maximum appears over the ITCZ. A belt of heavy rain appears along the West Coast extending from California to British Columbia and Alaska on the windward slopes of the Great Rocky Mountain
266
10
Monsoon over Central America and Adjoining Southwestern North America
Fig. 10.10 Climatological (1971–2000) rainfall (mm day–1 ) over Central and North America and adjoining oceans during January (after Xie and Arkin, 1996)
Ranges. In North America, the Western Plateau region registers little rain. The main rainbelt of the Continent lies in the east, with a maximum centered over the Atlantic Ocean, some distance off-shore. Another concentrated area of rain appears over the Gulf Coast States. Most of Canada and Polar regions appear to have little rain. In July, the narrow rainbelt along the southeastern States of Central America appears to have expanded to cover practically the whole region. However, the main rainbelt still lies in the south, especially in the ITCZ which appears to have moved further north from near equator to about 15–20◦ N. In the area of Mexican monsoon the rainbelt appears to have extended as far north as the Sonoran-Arizona desert. However, the intensity of rain appears to decrease steadily from south to north. The dry zone of the subtropical high pressure belt over the Pacific Ocean appears to have expanded in July to cover a much wider area of the ocean as well as some parts of Western United States. The January rainbelt along the western coast of North America has receded northward and now covers only the extreme northwestern parts of British Columbia and Alaska. The major rainbelt centered off the east coast of North America appears to have moved northwestward to cover an extensive area over the land. This rainbelt appears
10.4
Rainfall over Central America and Adjoining Areas
267
Fig. 10.11 Climatological (1971–2000) rainfall (mm day–1 ) over Central and North America and adjoining oceans during July (after Xie and Arkin, 1996)
to extend as far west as the Western Plateau and the Rocky Mountains. The rainbelt also appears to extend northward to Canada and the Polar regions.
10.4.1 Annual Rainfall over Mexico Garcia (1965) presented the mean monthly rainfall at a number of stations in Mexico, from which Mosino Aleman and Garcia (loc. cit.) worked out the isohyets of annual rainfall distribution over Mexico. Figure 10.12, adapted from their work, shows the distribution of the annual rainfall over Mexico. The isohyets in Fig. 10.12 appear to suggest the great influence of the Mexican plateau and the mountain ranges of Mexico and other Central American States on the distribution of annual rainfall over Mexico. According to Mosino Aleman and Garcia (loc. cit.), the distribution displays a ‘U’ shape with large amounts falling along the outer seaward-sloping surfaces of the two northwest-southeast oriented mountain ranges, viz., the Sierra Madre Occidental in the west and south and the Sierra Madre Oriental in the east, and much reduced rainfall on the elevated Mexican Plateau in between. Over most of the southern parts of Mexico and other Central
268
10
Monsoon over Central America and Adjoining Southwestern North America
Fig. 10.12 Isohyets of mean annual precipitation (cm) over Mexico (adapted from Mosino Aleman and Garcia, 1974)
American states, more than 90% of the annual rain falls during summer, the percentage decreasing rapidly towards the north of Mexico. It is likely that a large part of the heavy rain that falls along the coastal belts in the south is due to either Pacific cyclones or Atlantic hurricanes depending upon the track of these disturbances.
10.4.2 Source of Moisture for Rainfall over the Arizona-Sonoran Desert Until recently, a question that plagued the minds of many was regarding the source of moisture for summertime precipitation over the Sonora-Arizona desert region. Was the source the Gulf of Mexico, as some believed, or was it the Gulf of California and the Pacific Ocean which had long been suspected? Some early circulation studies (e.g., Bryson, 1957) based on harmonic analysis of rainfall over the southwestern United States and adjoining northwestern Mexico had concluded that bulk of the moisture for summertime rainfall over the desert regions of Sonora-Arizona was drawn from the Gulf of Mexico. But this conclusion was challenged by Sands (1959) who believed that the root of the moisture tongue may lie over the eastern Pacific ocean where the southeasterly tradewinds of the southern hemisphere after crossing the equator turn into a S/SWly aircurrent and converge into the northeasterly trades of the northern hemisphere and rise in convection before being advected by the monsoon aircurrent northward to the desert region. The exact mechanism of the advection could not be worked out at the time for lack of observational support, especially from ocean areas.
10.4
Rainfall over Central America and Adjoining Areas
269
The question, however, was reopened by some recent studies (e.g., Hales, 1972, 1974; Brenner, 1974; Badan-Dangon et al., 1991; Douglas et al., 1993) which favored a Pacific Ocean origin of the moist aircurrent. The findings of these studies appear to be supported by a recent study by Higgins et al. (1997) who showed that a large part of the SE/S-ly moisture-bearing aircurrent from the Gulf of Mexico which enters the Gulf coast near Texas and New Orleans in summer is diverted into a low-level jet over the Great Plains of North America, leaving little to enter into the ‘heat low’ circulation centered over the Sonora-Arizona region. A possible mechanism for supply of moisture from the Gulf of California suggested here is an interaction between the circulations around the ITCZ over the equatorial eastern Pacific Ocean and the quasi-stationary ‘heat low’ over the Sonora-Arizona region. It is well-known that with advance of summer, the ITCZ moves northward of its winter location near the equator. The movement allows the cold Peruvian aircurrent of the southern hemisphere to cross the equator and converge into the ITCZ. The tradewinds from the subtropical anticyclone of the eastern Pacific Ocean also carry cold California current to converge into the ITCZ. Thus, during interaction, the formation of an extended trough between the ITCZ and the quasi-stationary heat low over the Sonora-Arizona desert in the lower troposphere, as shown schematically in Fig. 10.13, enables the moist southerly flow from the Gulf of California to enter
Fig. 10.13 Schematic showing interaction between circulations around the ‘heat low’ (HL) over the Sonora-Arizona region and the ITCZ (thick dashed line) over the eastern Pacific Ocean and formation of an extended trough in the lower troposphere during the northern summer. Thin (thick) arrow shows the direction of airflow (moisture tongue). Areas of low-level moisture convergence and rainfall are hatched
270
10
Monsoon over Central America and Adjoining Southwestern North America
the desert region. Moist convection over the region is favored by the presence of a diverging anticyclonic circulation in the upper troposphere (see Fig. 10.7).
10.5 Some Characteristic Features of Weather over Central America Central America and adjoining ocean areas on both sides of the landmass especially the coastal belts are exposed to some special types of circulation and weather in different seasons. The following are some of them:
10.5.1 Weather Associated with W’ly Waves Throughout the year, waves in midlatitude westerlies affect Central America, as they move eastward across the northern part of the region. Their frequency appears to be much higher during the winter than summer. Their interaction with the quasistationary monsoon systems produces diverse effects. During their movement, surface pressure alternately rises and falls, and in their wake they bring extremely cold and dry weather over some areas. However, it is also during the passage of these disturbances that light rain or snowfall occurs over the mountainous regions, especially over the elevated Mexican plateau, during winter.
10.5.2 Weather Associated with ‘Northers’ During the passage of a westerly wave disturbance across the subtropical belt of North America, an intense high pressure system often develops over Mexico behind a wave trough located over the southern United States of America, as shown schematically in Fig. 10.14. As part of the anticyclonic circulation around such a high pressure system, extremely cold, dry winds blow over the western half of the Gulf of Mexico from a northerly direction. These winds are popularly called ‘Northers’. As the northers move southward, they pick up moisture from the underlying ocean surface before they meet the northeasterly tradewinds which diverge from the subtropical high pressure belt of the Atlantic Ocean. Together, the convergent winds sense the presence of the mountain and turn westward to cause some precipitation along the slopes of the Sierra Madre Oriental Mountain. However, unable to cross the high mountain, the flow turns southward till it finds a mountain gap in the Tehuantepec area where it crosses the land to emerge into the eastern Pacific Ocean as a cold dry wind. During the outbreak of the Northers, convergence in the
10.5
Some Characteristic Features of Weather over Central America
271
Fig. 10.14 Schematic illustrating the flow of northerly winds (Northers) over western Gulf of Mexico from a high pressure system located over Mexico and southwestern United States. The thick line with arrow shows a typical northerly streamline
northeasterly tradewinds over the northwestern Caribbean Sea produces heavy rain on the windward sides of Yucatan peninsula and Nicaragua. The onset of the Northers is heralded by a rise of surface pressure over most parts of Mexico.
10.5.3 Land and Sea Breezes on the Pacific Coast of Mexico During the early part of summer, i.e., the premonsoon season, strong land and sea breezes are experienced on the Pacific side of Mexico under the sheltering action of the Sierra Madre Occidental mountain range. The sea breeze occurring during the day is aided by anabatic winds moving up the mountain slopes to cause convective clouds (anabatic winds are winds which blow upward along mountain slopes in daytime when the mountainside gets warm). Similarly, the land breeze during the night is aided by katabatic flow (katabatic winds are opposite of anabatic winds, i.e., they are downslope winds when the mountainside gets cold) from the mountains to cause convective clouds offshore over the nearby sea. Thus strengthened by the mountain winds, land and sea breezes play an important role in the low-level atmospheric circulation over Mexico. The people of Mexico living in the coastal belt welcome the sea breeze during summer, because its cooling effect gives them relief from the sweltering heat of the day. However, land and sea breezes fail to develop on days when the synoptic-scale background winds over the region are too strong and mask the local winds.
272
10
Monsoon over Central America and Adjoining Southwestern North America
10.5.4 Temporales of the Caribbean Sea and the Gulf of Mexico During summer, easterly waves travelling in the tradewinds over the Caribbean Sea often move up to the eastern coast of Central America. The structure and properties of these waves over the tropical North Atlantic Ocean were discussed in detail by Riehl (1945). On approaching the landmass of Central America, an E’ly wave comes under the influence of the summertime ‘heat lows’ as well as the blocking effect of the coastal mountain ranges. In order to negotiate these land surface features, a part of the flow around the wave trough first turns anticyclonically and then cyclonically so as to converge into the ‘heat low’, or converge and rise at the mountainside. In either case, the moisture convergence produces a lot of precipitation along the east coast of Mexico and the Yucatan and Nicaragua peninsulas. Locally, such episodes of bad weather brought in by E’ly waves are known as ‘temporales’ in Central America. Figure 10.15 illustrates the formation of a ‘temporale’ in an E’ly wave trough (double-dashed line) approaching the Yucatan peninsula. In a few cases, while flowing around the ‘heat low’ over the Yucatan peninsula, a part of the easterlies which turn southward may find a mountain pass, such as that between the Bahia de Campeche and the Gulf of Tehuantepec, to escape to the Pacific Ocean.
Fig. 10.15 Illustrating the formation of a ‘temporale’ in an E’ly wave trough (double-dashed line) approaching Yucatan peninsula
10.5.5 Hurricanes and Tropical Storms These are unusually strong synoptic-scale cyclonic circulations of the tropical North Atlantic Ocean which affect Central America every summer, mainly from June through November. Most of them develop from easterly waves which originate over
10.5
Some Characteristic Features of Weather over Central America
273
the African continent and travel westward in the deep easterly tradewinds which blow between the ITCZ and the ridge of the subtropical high pressure over the Atlantic. However, on entering the western Caribbean Sea and the Gulf of Mexico where the ocean surface is warm, most of them develop rapidly into tropical storms and hurricanes. The mechanism of formation of these disturbances and their structure, development and movement were discussed in detail in Chap. 3. The oceanic area around the Yucatan peninsula where a ‘heat low’ pre-exists is particularly prone to the development of these storms and hurricanes. Westward-propagating E’ly waves, depressions, or storms, developing into hurricanes over the Gulf of Mexico usually recurve northwestward and then northward to enter the southern coast of the United States but a few may continue to move westward to strike the eastern seaboard of Mexico. Few of them penetrate into the interior of the Mexican plateau. So, the damage, destruction and loss of life due to these storms and hurricanes are usually confined to the coastal belt only. However, a few of them which follow a more southerly track, say a latitude between 15 and 20◦ N, may cross over to the Pacific Ocean side and redevelop. After development, they usually move northwestward to affect the west coast of Mexico but a few may continue on their westward track. However, few of them survive beyond a few 100 km over a cold ocean.
Chapter 11
Extratropical Monsoon over North America
11.1 Introduction In the following pages, we intend to show using materials from several earlier studies (e.g. Bryson and Hare, 1974; Tang and Reiter, 1984; Higgins et al., 1997) as well as NCEP Reanalysis and other supplemental data that extratropical monsoon over North America occurs in two different longitudinal segments, one over the central part on the leeside of the Rocky Mountains Plateau, which Tang and Reiter (1984) termed a Plateau Monsoon, and the other over the East Coast region which we decide to call the East Coast Monsoon. In the course of our analysis of monsoons over North America, we try to show that the structure and properties of the monsoon circulation over the extratropical belt of this continent, may be little different from those of monsoons occurring over other regions of the globe. Here also, seasonal reversals occur in circulation and airmass properties and monsoon circulation shows a wave structure with alternate zones of moisture convergence and divergence. One of these moisture convergence zones appears to be the eastern slopes of the Rocky Mountains and adjoining Great Plains area, and the other in Eastern North America across the eastern seaboard of the continent during northern summer.
11.2 Climatological Background of North American Monsoon Since a discussion of the climatology of North America and surrounding oceans may facilitate an understanding of the monsoon over the continent, we begin this chapter by drawing attention to some of the important physical and dynamical properties of the land and the surrounding oceans which were introduced in Chap. 10, in connection with monsoon over Central America and adjoining Southwestern North America.
K. Saha, Tropical Circulation Systems and Monsoons, C Springer-Verlag Berlin Heidelberg 2010 DOI 10.1007/978-3-642-03373-5_11,
275
276
11
Extratropical Monsoon over North America
11.2.1 Physical Features of the Land The Continent of North America including the United States, Canada and Alaska has a large landmass, extending latitudinally from near the subtropical belt to almost the Arctic Circle and longitudinally from about the dateline to about 50W. The geological wonder, the ‘Continental Divide’, separates the western and the eastern parts of the continent. The continent has its main mountain ranges and high ground in its western part, extensive Plains and Lowlands over the central part, and then a low mountain range almost paralleling the coast in its eastern part. Figure 11.1 presents a vertical cross-section along 40◦ N, which passes through almost the middle of the United States, showing the approximate topographic heights above mean sea level of the main mountain systems and valleys and lowlands along this latitude. The vertical profile in Fig. 11.1 shows the structure of the Western Plateau consisting of several north-south oriented mountain ranges and valleys from the narrow Coastal Mountain Range of the west coast to the Great Rocky Mountains, spanning about 30◦ of longitude, with an average height of about 2 km a.s.l. East of Continental Divide along about 105◦ W, the Plateau descends to the level of the Great Plains of the Mid-West. The Great Plains descends to its lowest level along the Mississippi River valley and then rises gently over the Central Lowlands before rising steeply to the Appalachian Mountain Ranges and descending to the Coastal Lowlands of Eastern North America.
Fig. 11.1 A vertical section along 40◦ N showing the approximate topographic heights (a.s.l.) of the Western Plateau and other landforms across the continent
11.2
Climatological Background of North American Monsoon
277
The importance of the above-mentioned Mountain systems of North America in the context of water resources of the continent can hardly be over-emphasized. Rain and snowfall at these mountains feed the Great Mississippi and Missouri Rivers and their numerous tributaries. Similar rain and snowfall along the East Coast Mountains feed several river systems, large and small, in that region.
11.2.2 Semi-permanent High and Low Pressure Systems over Oceans The oceans around North America exercise great thermal and dynamical influence upon the climate of the continent and the seasonal monsoon circulations. Maximum influence comes from the semi-permanent high pressure cells over the subtropical belts of the Pacific and Atlantic Oceans. The polar high pressure cell of the Arctic Ocean has also a great influence. Semi-permanent low pressures over the Aleutian Islands area of the North Pacific and the Icelandic area of the North Atlantic Ocean also influence monsoon circulation over the North American continent. 11.2.2.1 The Subtropical High Pressure System of the Pacific Ocean Airstreams diverging from the subtropical high pressure cell of the Pacific Ocean move equatorward as well as poleward. Those moving equatorward carry cold air to converge into the circulation around the warm ITCZ Those blowing poleward carry warm, moist air of the tropical ocean to converge into the circulation around the Aleutian Islands low pressure system. During winter, the diverging winds also move across the continent to converge into the circulation around the low pressure system over the Icelandic area. They also find their way to the coastal mountain ranges where they may converge to yield heavy precipitation. 11.2.2.2 The Subtropical High Pressure System over the Atlantic Ocean Like its counterpart over the Pacific Ocean, the subtropical high pressure cell over the Atlantic Ocean play the dual role of feeding cold air to the tropical convergence zone and warm air to the Icelandic area low pressure zone. The diverging winds also find their way to the North American continent during the northern summer when a heat source advances northward across the Western Plateau. A branch of the diverging currents also move northward across the eastern seaboard to converge into a low pressure area over the land. In all the areas, on convergence, they rise in convection to produce clouding and precipitation. 11.2.2.3 The Arctic Ocean High Pressure System A semi-permanent high pressure with anticyclonic circulation resides over the frozen Arctic Ocean. Icy-cold winds diverging from this cold source find their
278
11
Extratropical Monsoon over North America
way to the polar front zone where they converge into the warmer airstreams of the lower latitudes. 11.2.2.4 Semi-permanent ‘Aleutian Low’ and ‘Icelandic Low’ Two semi-permanent low pressure systems over oceans influence the atmospheric circulation over North America. These are the ‘Aleutian Low’ of the North Pacific Ocean and the ‘Icelandic Low’ of the North Atlantic Ocean.
11.3 The Seasonal Movement of Heat Sources and Sinks During northern winter, the heat source for the North American monsoon is located over South America. With decline of winter, it moves northward. We do not know for certainty how this heat source moves, or, by what route. Most likely, as in other regions, it follows a land route via the narrow Panama States and then from near Yucatan Peninsula along the Mexican Plateau to the southwestern part of North America. From there, it moves further northward along the Western Plateau of North America to as far as the US–Canada border. A similar move occurs of an oceanic heat source which lies over the western Atlantic Ocean off the east coast of
Fig. 11.2 Approximate northern boundary of the heat source in different months, January–July, for monsoon over North America. Western Plateau monsoon (full line), East Coast monsoon (dashed)
11.4
Seasonal Circulations – Monsoons
279
North America during the winter. With increased warming of the land surface and differential heating between land and ocean, the oceanic heat source moves northnorth-westward across the east coast to arrive at the southeastern part of Canada by July. The subtropical high pressures over the oceans also respond to the change of season from winter to summer. Simultaneously with the seasonal movement of the heat sources, they also migrate northward from their winter location along 30◦ N to about 40◦ N over the Pacific Ocean and about 35◦ N over the Atlantic Ocean. With decline of summer, the heat sources and sinks in both the sectors return to their winter locations, following almost the same route as during advance, but in reverse direction. In Fig. 11.2 we show the approximate northern boundary of the heat source in different months, from January to July, in the two sectors.
11.4 Seasonal Circulations – Monsoons Since monsoon is basically a circulation between a heat source and a heat sink, a circulation initiated by a heat source on a mountain plateau has been called a plateau monsoon. Tang and Reiter (1984) identify two regions of the globe where such plateau monsoons are to be found, one over the Tibetan Plateau of Asia and the other over the Western Plateau of North America. A monsoon circulation is also to be found in the eastern part of North America where the heat source is located over the plains of North America and the heat sink over the adjoining Atlantic Ocean. The differential heating between land and ocean drives this monsoon across the eastern seaboard of the continent. To study the circulation over North America, Bryson and Hare (loc. cit.) constructed resultant surface streamlines over North America and adjoining ocean areas with data derived from various sources including NAVAIR (1966) for different months of the year. Tang and Reiter (loc. cit.) also presented maps of low level circulations over North America in different months in connection with their study on Plateau monsoon. The broad features of the seasonal circulations discussed below are largely based on their maps. There are often large variations in circulations within a season, especially during the transition months represented by April and October.
11.4.1 The Winter Monsoon (December–February) During the northern winter months, December to February, an intense anticyclonic circulation develops around a cold high pressure over the Western Plateau of North America. An anticyclonic circulation also develops around a high pressure area over the southeastern States along the Gulf Coast. A deep trough of low pressure with cyclonic circulation lies in between these two high pressures. An anticyclonic
280
11
Extratropical Monsoon over North America
circulation also appears in the east with its ridge along the Appalachian Mountains. Between the subtropical anticyclone over the Atlantic Ocean and the mountain ridge, a trough of low pressure with cyclonic circulation lies off the east coast of North America. The Pacific westerlies are unusually strong and active in this season. Diverging from the subtropical anticyclone they converge at the sides of the coastal and inland mountains, producing rain and snowfall on the windward sides. After moving around the high pressure ridge on the Rocky Mountains, they descend steeply on the leeside where subsidence warming produces a well-marked trough of low pressure. The formation of this leeside trough over the Great Plains prepares the ground for a strong, warm dry wind, known as the Chinook wind, to blow from the north by the mountainside. However, the actual direction of the wind depends upon the location of the leeside trough and topography. A similarly strong, hot and dry mountain downslope wind, known as a Santa Ana wind, blowing from a E/NE’ly direction, is also encountered in the mountain pass and river valley of Santa Ana, California, on the southern side of the Rocky Mountains block. Blowing over the desert surface, it carries a lot of dust. Its extreme heat and dryness often causes forest fires. The Meteorological Glossary (2000) of the American Meteorological Society describes a Chinook wind as follows: On the eastern slopes of the Rocky Mountains the Chinook generally blows from the west or southwest, although the direction may be modified by topography. Often the chinook begins to blow at the surface as an arctic front retreats to the east, producing dramatic temperature rises. Jumps of 10–20◦ C can occur in 15 minutes, and at Havre, Montana, a jump from –12 to +5◦ C in 3 minutes was recorded. Occasionally the arctic front is nearly stationary and oscillates back and forth over an observing station, causing the temperature to fluctuate wildly as the station comes alternately under the influence of warm and cold air. As in the case of any foehn, Chinook winds are often strong and gusty. They can be accompanied by mountain waves, and they can occur in the form of damaging downslope windstorms. The air in the Chinook originates in the midtroposphere above the ridgetops, and its warmth and dryness result from subsidence. When moisture is present, a variety of mountain wave clouds and lee wave clouds can form, such as the Chinook arch of the Canadian Rocky Mountains west of Calgary, Alberta. The Chinook brings relief from the cold of winter, but its most important effect is to melt or sublimate snow: A foot of snow may disappear in a few hours. . .
The Chinook wind constitutes an important component of winter monsoon over North America. It is the winter counterpart of the S’ly Low Level Jet (LLJ) of the summer monsoon, to be dealt with shortly. At 300 hPa (see Fig. 10.8), the W’ly airstreams diverge over the Pacific Ocean. However, on entering North America they converge over the eastern parts of the continent before diverging again before leaving the eastern shores. According to Bryson and Hare (1974), the W’ly airstream that enters the continent reaches a speed maximum of 36 m s–1 at 300 hPa over the eastern part of the United States during January. The above-mentioned features of the mean surface circulation over North America during winter are shown in Fig. 11.3.
11.4
Seasonal Circulations – Monsoons
281
Fig. 11.3 Streamlines showing winter monsoon surface circulation over North America. Thick arrow denotes Chinook wind; dashed lines are convergence lines
11.4.2 The Spring Transition Season (March–May) The winter monsoon over the Western Plateau region undergoes rapid changes during the spring transition season, March–May. With increased warming of the land surface, the high pressure ridge on the mountains weakens, leading to a weakening and northward shift of the trough of low pressure on the leeside of the Rockies. The Chinook events are rare. The southern anticyclone also weakens and ultimately becomes non-effective. These changes allow the warm, moist tradewinds from the subtropical anticyclone of the Atlantic Ocean to extend its sphere of influence westward and northward. The Pacific airstream loses its influence and becomes marginalized, mostly to the western side of the Rocky Mountains, allowing the arctic airstream to make deep inroads southward to meet the tradewinds directly over the United States. These features are depicted in Fig. 11.4, which shows the mean surface streamlines over North America during the spring transition season. At 300 hPa, the W’ly wind over the northern Pacific Ocean strengthens, while that over the United States weakens.
11.4.3 The Summer Monsoon (June–August) A comparison of the mean thermal and pressure fields over the continent during January and July, presented earlier in Figs. 10.4, 10.5, 10.6, and 10.7, leaves little
282
11
Extratropical Monsoon over North America
Fig. 11.4 Streamlines showing surface circulation over North America during the spring transition season
doubt about the seasonal reversals that occur in these fields. They are also suggested by a comparison of the circulation field presented in Fig. 11.3, with that in Fig. 11.5. The most conspicuous feature of the summer circulation over the Western Plateau is the replacement of the winter high pressure ridge over the Plateau by a low pressure trough and the low pressure trough on the leeside of the plateau on its eastern side by a high pressure ridge. In fact, the high pressure ridge on the eastern side of the plateau becomes a part of the subtropical high pressure belt of the Atlantic Ocean, which extends its influence westward in summer. The pressure gradient across the eastern slopes of the plateau drives a strong S’ly wind which often reaches jet speed and has come to be known as the Great Plains Low Level Jet (LLJ). The LLJ has been studied extensively by several workers (e.g., Means, 1952; Blackadar, 1957; Hoecker, 1963; Lettau, 1967; Bonner, 1968; Bonner and Paegle, 1970; Tang and Reiter, 1984; Helfand and Schubert, 1995; Higgins et al., 1997). These studies have improved our knowledge and understanding of this low level jet and shown how it forms an important component of the Plateau monsoon system. Their findings throw light on the following.
11.4
Seasonal Circulations – Monsoons
283
Fig. 11.5 Mean surface circulation over North America during summer (JJA). Thick long arrow along the eastern slopes of the Rocky Mountains shows the Low Level Jet (LLJ). Dashed lines indicate moisture convergence zones
11.4.3.1 Structure of the Monsoon Boundary Layer Tang and Reiter (1984) in their analysis of 12 GMT data during July found the thickness of the layer in which the resultant wind is southerly (i.e., from a direction between 135 and 225◦ ) and the height above ground at which a windspeed maximum is encountered. The details of their analysis are presented in Fig. 11.6 (their Fig. 9). It may be seen from Fig. 11.6 that a strong S’ly airflow which emanates from the region of the US–Mexico border proceeds northward as far as the Dakotas. The average height of this low-level wind maximum above ground appears to be about 600–700 m and remains uniform at this value, though the ground slopes gently upwards from the Gulf Coast towards the northwest. Though the wind in the boundary layer is basically driven by the summertime large-scale pressure gradient between the low pressure over the Plateau and the high pressure over the Great Plains, the height at which the LLJ is encountered and its intensity appears to undergo a remarkable diurnal variation. About this variation, Tang and Reiter (loc. cit.) write: The LLJ has been described as a boundary layer phenomenon which maximizes at inversion height at night. It is thought to be a consequence of the diurnal coupling and decoupling of this layer with the ground by quasi-periodic intensity changes in the frictional momentum transport and of similar, quasi-periodic changes in the vertical, geostrophic wind profile.
284
11
Extratropical Monsoon over North America
Fig. 11.6 Thickness of the layer (km, solid lines) in which during July the resultant wind has a direction between 135 and 225◦ (‘south wind’) and maximum mean speed (m s–1 , dashed lines) of ‘south wind’ in the lowest 3 km above mean sea level. The numbers plotted next to stations indicate the height above ground (km) at which the wind maximum is encountered. Analysis pertains to 1200 GMT data
This means that the height and intensity of the LLJ fluctuates between day and night, going up and becoming weaker during daytime and lowering and becoming strong during nighttime. The above description of the monsoon boundary layer and its diurnal variations has been well supported by subsequent studies (e.g., Higgins et al., 1997). 11.4.3.2 Moisture Budget and Precipitation Higgins et al. (loc. cit.) used observations as well as recently available assimilated datasets produced by the National Centers for Environmental Prediction (NCEP)/National Center for Atmospheric Research (NCAR) as well as the National Aeronautics and Space Administration/Data Assimilation Office (NASA/DAO) to examine the influence of the LLJ on summertime precipitation and moisture transport over the Central United States, and reported the following results: (i) There occurs a well-defined nocturnal maximum in the diurnal cycle of precipitation over the Great Plains region during Spring and summer months;
11.4
Seasonal Circulations – Monsoons
285
(ii) During summer, an excess of 25% more precipitation falls during the night than during the day over a large area of the Great Plains, with a commensurate decrease in the percentage amount of nocturnal precipitation along the Gulf Coast; (iii) The nighttime maximum in precipitation along the Gulf Coast slowly shifts northward from the Lower Mississippi valley to the Upper Midwest during the late spring and summer months and then back again during the fall; (iv) Composites of observed nighttime precipitation during LLJ events show a fundamentally different pattern in the distribution of precipitation compared to non-jet events; (v) LLJ-related precipitation is found to be associated with maximum moisture transport to the precipitation zone. It decreases when the area of maximum moisture transport recedes towards the south and east; (vi) The impact of the LLJ on the overall moisture budget and precipitation is considerable with low level inflow from the Gulf of Mexico increasing by 45%, on average, over nocturnal mean values.
11.4.3.3 Seasonality in the Distribution of Orographic Precipitation The winter and summer monsoons over the Western Plateau make great impact on the distribution of mean monthly precipitation at stations located on either side of a mountain range. Tang and Reiter (loc. cit.) cite the case of two stations, Crested Butte (38◦ 52 N, 106◦ 58 W, 2699 m MSL) and Buena Vista (38◦ 51 N, 106◦ 08 W, 2425 m MSL), located almost at the same latitude less than 100 km apart, but on opposite sides of the Sawatch Mountain Range of Colorado, which shows a different seasonal pattern of precipitation (Fig. 11.7). Crested Butte situated on the western side of the Mountain Range has a winter maximum with a lot of snowfall which qualifies it as a ski resort, while Buena Vista situated on the eastern side has a
Fig. 11.7 Monthly precipitation (percent of annual) at Crested Butte (west) and Buena Vista (east), Colorado (reproduced from Tang and Reiter, 1984; their Fig. 13)
286
11
Extratropical Monsoon over North America
summer maximum in rain which falls mostly during July–August, as a consequence of upslope motion of moist air of the LLJ forced by the heat low over the Plateau. On the other hand, at Crested Butte, snowfall is caused by the upslope motion of moist Pacific air as it seeks to cross the Mountain.
11.4.4 The Autumn Transition Season (September–November) After attaining peak intensity during July, the summer monsoon over North America starts withdrawing southward, making room for the winter monsoon to take over. However, the movement at first is gradual and it is not until September that the Southern anticyclone re-forms over southeastern United States, a change which allows the W’ly Pacific airstream to cross the Rocky Mountain Ranges and move eastward and the N’ly Arctic airstream to make greater inroads southward. Approximate boundaries separating these airstreams are marked by bold dashed lines in Fig. 11.8 which shows the mean surface streamlines for the transition season, September–November.
Fig. 11.8 Mean surface streamlines during the transition season, September–November (after Bryson and Hare, 1974; NAVAIR, 1966)
11.6
Some Characteristic Features of East Coast Monsoon
287
11.5 Interaction of Monsoons with W’ly Wave Disturbances Seasonal monsoons we have discussed in preceding paragraghs are seldom free from interference by W’ly waves which move across the continent with their frontal systems in quick succession almost throughout the year. During their passage they interact with both the Plateau Monsoon as well as the East Coast Monsoon. The interaction leads to large-scale redistribution of the meteorological fields and reorganization of the circulation systems with periodic weakening and strengthening of observed weather systems, as the W’ly wave gets in and out of phase with the monsoon wave. For example, during winter, when the monsoon wave and the W’ly wave are in phase, we may have a deep trough of low pressure positioned over the Great Plains of North America, a strong N/NW’ly jetstream converging into the trough from the northwest and a still stronger S/SW’ly jetstream diverging from it to the east in the upper air. On such occasions, icy cold arctic air may penetrate deep into the southwestern part of Central United States to the west of a cold front, producing stormy conditions and a severe cold wave across the continent as the W’ly wave moves eastward. The interaction process appears to be reversed during the summer. An in-phase interaction then causes a deep penetration of tropical airstream northward to as far as southwestern Canada, thereby restricting the influence of the Pacific airstream to the west of the Mountains and allowing tropical airstream to meet arctic airstream directly along a boundary which lies along the southern parts of Canada. On such occasions, a heat wave may sweep across the United States as the W’ly wave moves eastward. The summer jetstreams are usually much weaker over North America than winter jetstreams.
11.6 Some Characteristic Features of East Coast Monsoon 11.6.1 Seasonal Variations and Reversals Figure 11.9 shows the distribution of seasonal variations (deviation from the annual mean) of (a) surface pressure P, and (b) rainfall R, along 40◦ N across the eastern seaboard of North America during winter (DJF, full line) and summer (JJA, dashed line). The zonal distribution of seasonal pressure variation shows a very large increase in surface pressure over the land and corresponding decrease over the ocean during winter. The field reverses during the summer with a large fall of pressure over land and rise over the ocean. The characteristic feature is also reflected in the field of rainfall which shows a rainfall maximum over the ocean and almost dry conditions over the land in winter. In summer, the rainfall maximum is still over the ocean, but its intensity is less, while there is considerable increase of rainfall over the land. The seasonal reversals of the distributions of pressure and rainfall between land and ocean across the eastern seaboard of North America stand out in Fig. 11.9.
288
11
Extratropical Monsoon over North America
Fig. 11.9 Distribution of seasonal variation (deviation from annual mean) of (a) mean sea level pressure P (mb), and (b) Rainfall R (mm day–1 ) during winter (DJF, full line) and summer (JJA, dashed line) along 40◦ E across the eastern seaboard of North America. Arrow at baseline shows location of coast
11.6.2 Monsoonal Characteristics of the East Coast Region If differential heating between a land and an ocean and its seasonal reversal are what govern a monsoon circulation and distribution of monsoon rains, then the extratropical belt of Eastern North America and its adjoining North Atlantic Ocean can rightfully claim to be a monsoonal region, as shown by its climatology in the preceding section. In Table 11.1, we summarize the monsoonal characteristics of the region. Table 11.1 Monsoonal characteristics of Eastern North America and adjoining Atlantic Ocean Season
Variable
Eastern North America
Atlantic Ocean
DJF (JJA)
Temperature Pressure Rainfall (anom)
Cold (warm) High (low) – (+)
Warm (cold) Low (high) + (–)
It is often difficult to see the monsoonal characteristics in the circulation field over North America on account of interference from baroclinic wave disturbances which travel in midlatitude westerlies.
11.7 Role of the Appalachian Mountain Range – Leeside Cyclogenesis – Northeast Storms It is often difficult to assess the full impact of the Appalachian Mountain Range, the height of which varies from about 1 to 2 km a.s.l., upon the seasonal monsoon and the approaching baroclinic waves in the westerlies. Here, both thermal
11.8
Interaction of Monsoon with Storms and Hurricanes
289
and orographic effects come into play. The airflow negotiating the mountain range is required to conserve potential vorticity in the sheared flow on both sides of the mountain range. In order to fulfill this requirement, the airflow develops strong cyclonic vorticity on the windward side of the mountain range, anticyclonic vorticity on the climb, and cyclonic vorticity during descent on the leeside. On account of these changes in vorticity, a strong cyclonic circulation often tends to develop on the windward side which draws warm, humid air from the Gulf of Mexico and Southeastern States to produce heavy rain and stormy weather on the windward side of the Mountain Range. However, loss of cyclonic vorticity while negotiating the mountain often makes a difference to the type and intensity of weather on the leeside. Most often it regains cyclonic vorticity as it rolls out over the adjoining ocean and with added heat and moisture drawn from the underlying ocean surface develop into storms which have come to be known as ‘Northeast Storms’. According to Meteorological Glossary (Second Edition) of American Meteorological Society (2000), ‘a Northeast Storm is a cyclonic storm of the east coast of North America, so called because the winds over the coastal area are from the northeast. They may occur at any time of year but are most frequent and most violent between September and April. Northeast storms usually develop in lower-middle latitudes (30–40◦ N) within 100 miles east or west of the coastline. They progress generally northward to northeastward and typically attain maximum intensity near New England and the Maritime Provinces. They nearly always bring precipitation, winds of gale force, rough seas, and, occasionally, coastal flooding to the affected regions’.
11.8 Interaction of Monsoon with Storms and Hurricanes The W’ly waves during their eastward movement across Eastern North America and adjoining Western Atlantic Ocean occasionally encounter poleward-moving hurricanes coming up from the southern part of the North Atlantic Ocean. What happens at such encounter? It is difficult to predict the outcome in any given case without detailed observations over the ocean. However, available observations appear to suggest that most often they cooperate with each other and move together, retaining their individual identity and circulations, though with varying intensity, over long distances downstream. They usually follow the route of the Gulf Stream to a considerable distance from the coast. The airmass transformation and cyclogenesis in this case favor development and movement. The case is similar to that which occurs over the Western Pacific Ocean off the coast of Northeastern Asia. These were well demonstrated by Sanders and Gyakum’s study of cyclogenesis over the two oceans, the results of which were presented in Fig. 5.7. Another interesting situation arises when a northward-moving hurricane from the Gulf of Mexico after crossing the Coast encounters a W’ly wave disturbance over the eastern part of North America during the peak summer monsoon season. Though the hurricane might have lost some of its original vigour after landfall and
290
11
Extratropical Monsoon over North America
been reduced to the status of a storm or low pressure system, the encounter appears to reorganize the tremendous amount of moisture carried by the storm circulation over a region which may already be holding a lot of moisture to produce widespread heavy rain and extensive floods and ground dislocations for prolonged periods. On such occasions also, one may observe heavier rainfall and more extensive inundations and dislocations on the windward side of the Appalachian Mountain Range than on the leeside.
Appendix Meanings of Uncommon Words/Terms Used in the Book
[In compiling this appendix, the author has relied to a considerable extent on the meanings of the terms, as given in two well-known publications: (1) Concise Dictionary of Physics and Related Subjects (1973) published by the Pergamon Press, Oxford, and (2) The Meteorological Glossary (2nd edition) of the American Meteorological Society (2000). However, complicated mathematical expressions, especially vector notations, have been avoided in order to make the meanings given easily intelligible to an average reader]
Words/terms
Meanings
Adiabatic change
Change occurring in a moving parcel of air which is not allowed to exchange properties with its environment When a body revolves in a circle about an axis perpendicular to the plane of its motion in a fixed frame of reference, it has an angular momentum given by the product of its mass, velocity and radius of the circle. Angular momentum can also be expressed in terms of the moment of inertia of the body by multiplying it with the angular velocity
Angular momentum
Absolute angular momentum
Anticyclonic circulation Baroclinic atmosphere
It is the angular momentum of a body in an absolute frame of reference and is obtained by adding the angular momentum of the body relative to the earth’s surface to the angular momentum of the earth’s surface relative to the absolute frame of reference A clockwise circulation around a high pressure area in the northern hemisphere The direction reverses in the southern hemisphere An atmosphere in which density varies along an isobaric surface (surface of equal pressure) due to horizontal temperature gradient
K. Saha, Tropical Circulation Systems and Monsoons, C Springer-Verlag Berlin Heidelberg 2010 DOI 10.1007/978-3-642-03373-5,
291
292
Appendix
Baroclinic instability
Instability of flow in a baroclinic atmosphere, which may allow growth of a perturbation if the wavelength of the perturbation and the horizontal temperature gradient meet certain critical conditions
Barotropic
A state of the atmosphere in which density remains constant along an isobaric surface.
Barotropic instability
Hydrodynamic instability of a barotropic atmosphere which may allow growth of a perturbation in a region across which the absolute vorticity changes sign. This is an inertial instability in which kinetic energy is the only form of energy transferred between the current and the perturbation.
Blackbody
A body which absorbs all heat radiation incident on it, remains in equilibrium with the radiation reaching and leaving it, and at a given steady temperature emits radiation with a flux density and spectral energy distribution characteristic of that temperature
Boundary layer
In meteorology, it is a thin layer of the atmosphere in contact with the earth’s surface in which the airflow strongly experiences the effects of the earth’s surface friction and vertical stability of the atmosphere A reversible working cycle of an ideal heat engine with maximum thermal efficiency in which the sequence of the cycle consists of isothermal expansion, adiabatic expansion, isothermal compression, and adiabatic compression to the initial state. It is the line integral of a fluid motion around the boundary of a closed surface
Carnot’s cycle
Circulation Cold surge
A sudden outbreak of strong cold winds at low levels of the atmosphere
Conditional instability
A state of instability of the atmosphere in which the lower layer (below about 3 km above surface) holding a lot of water vapour is vertically unstable, but the upper layer is stable, or vice versa Transport of heat energy in a medium by molecular vibration inside the medium when a temperature gradient exists between its two sides, usually in solids Process of heat transfer by actual physical movement of fluid elements in contact with a heated body In atmospheric motion, the coming closer of streamlines, which leads to an accumulation of atmospheric mass and rise of pressure
Conduction
Convection Convergence
Appendix
Coriolis force
Cyclonic motion
Cyclogenesis Depression
293
An apparent force (so called in honour of its discoverer, G.G. de Coriolis) exerted on a wind by the rotation of the earth. This force is always to the right of the wind direction in the northern hemisphere and to the left in the southern hemisphere Anticlockwise motion around a low pressure center in the northern hemisphere; clockwise in the southern hemisphere Process tending to generate cyclonic motion A low pressure area with a tangential windspeed of about 17 ms–1 at the top of the earth’s frictional layer
Diabatic change
Change in a parcel of air free to exchange properties with its environment Direct circulation A kinetic energy producing vertical circulation in which warm air rises and cold air sinks Divergence Separating out of streamlines leading to depletion of mass and fall of pressure Dry adiabatic lapse rate Rate of fall of temperature or pressure with height in a parcel of dry air, rising adiabatically in a dry atmosphere. Its value in the case of temperature is about –10◦ C/km. Dynamic instability Instability of a flow structure when its horizontal and vertical shear exceeds a certain critical limit ECMWF European Center for Medium-Range Weather Forecasts, located in UK ENSO
Abbreviation of El-Nino Southern Oscillation. An oscillation between sea surface temperature and pressure changes in the equatorial Eastern Pacific and corresponding changes in the equatorial Western Pacific and adjoining Eastern Indian Ocean
Entropy
A quantity characteristic of the thermodynamic state of a system. Mathematically, it may be expressed by the relation, dS = δQ/T, where S denotes entropy, Q a quantity of heat and T absolute temperature. Note that ‘dS’ denotes a total differential, while ‘δQ’ is simply a quantity of heat, not a total differential. In all naturally-occurring processes, entropy increases, except in cases which are reversible where it remains constant. An increase in the value of entropy signifies unavailability of useful energy
Equatorial circulation Equatorial heat source
Circulation along the equator A heat source characterized by warm, low pressure over the equatorial belt. It oscillates about its annual mean location following the seasonal movement of the sun
294
Appendix
Equatorial trough
The trough of low pressure of the equatorial heat source
Easterly wave
A westward-propagating wave in the tropical easterly tradewinds The latitudinal belts outside the boundary of the tropics
Extratropics FGGE
First GARP (Global Atmospheric Research Program) Global Experiment, 1979
Flux
Rate of flow of some atmospheric property, such as heat, momentum, energy, etc.
Gale
An unusually strong sustained wind of 14–23 ms–1
GARP
Global Atmospheric Research Program
General circulation GCM
Zonally – averaged annual – mean circulation of the earth’s atmosphere General Circulation Model
Geopotential meter (gpm)
Height in meter above the earth’s surface, taking into account the variation of the earth’s gravity with latitude
Geostrophic wind
It is a quasi-balanced wind between the pressure gradient force and the Coriolis force, blowing along the isobar A body becomes a heat source when it gets hotter than its immediate environment Reverse of a heat source; that is, a body becomes a heat sink when it gets colder that its immediate environment. It may be called a Cold Source Originally, the pole-to-equator direct circulation; now, the circulation between the subtropical ridge of high pressure and the equatorial trough of low pressure
Heat source Heat sink
Hadley circulation
High pressure area Hurricane
Hydrostatic approximation
Indirect circulation
Infrared
Area where pressure is higher than over its surroundings A deep tropical cyclone with central pressure (<980 mb) and tangential windspeed exceeding 32 ms–1 in the tropical Atlantic and Eastern Pacific Oceans An approximation made in the equation for vertical acceleration in a rotating fluid by balancing the vertical component of the pressure gradient force against the force due to gravity A circulation which generates available potential with warm air sinking and cold air rising; needed to drive a secondary kinetic energy – producing direct circulation in which warm air rises and cold air sinks In the frequency distribution of radiant energy from a blackbody, the part of the spectrum at frequencies shorter than the red line
Appendix
295
In-phase
Being in the same phase
Isentropic surface Instability
Surface of equal entropy A conditionally unstable state of the atmosphere which, if perturbed, will allow an amplification of the perturbation with time International Indian Ocean Expedition, 1962–1966
IIOE Intraseasonal oscillation Intertropical Convergence Zone (ITCZ) Isobar Isohyets Isotherm Jetstream Kinetic energy Lagrangian time Latent heat Longwave radiation
Oscillations within a seasonal oscillation A zone where the tradewinds of the two hemispheres converge, producing penetrative convection, clouding and precipitation A line joining places of equal barometric pressure A line joining places of equal amounts of rainfall A line joining places of equal temperature A narrow concentrated beam of strong winds with strong horizontal and vertical shear Energy of a body due to its motion Time counted with reference to a moving co-ordinate system Heat absorbed, or given up, by a solid, liquid, or vapour, at a fixed temperature during a change of phase Radiation emitted or absorbed by a body at the earth’s temperature at wavelengths longer than the visual range
Low pressure area Meridional wind
An area where pressure is lower than its surroundings Component of a wind along a meridian
Meridional circulation Midlatitudes
Component of a vertical circulation in a meridional plane Middle latitudes, usually between 30 and 60◦ latitudes
Midtroposphere
Middle part of the troposphere, usually between 5 and 9 km above sea level A divergent tradewind circulation which converges into the seasonally-migrating equatorial heat source, producing cloud and rain along the convergence zones and several other effects on prevailing airmass properties during advance and retreat The core or central part of a body
Monsoon
Nucleus Potential energy
Energy of a body by virtue of its position above a reference surface, or configuration of a system; usually measured by the work done in bringing the body from some reference level to the level of interest
296
Available potential energy Potential temperature
Equivalent potential temperature
Quasi-stationary Radiation
Ridge Saturated adiabatic lapse rate of temperature Shortwave radiation Streamline Subtropics
Appendix
The amount of potential energy available to do useful work out of the total potential energy stored in a body Temperature attained by a parcel of air at a given pressure surface, when it is lowered dry adiabatically to a reference surface, usually 1000 mb It is the temperature of a parcel of moist air at a pressure surface which is first lifted from its existing level to a level where all its moisture is precipitated out and the heat of condensation added to it, and then the air so warmed-up brought down dry adiabatically to a standard pressure surface, usually 1000 mb Stationary, or nearly so, usually applied to a wave, or front, etc. Radiant energy emitted by a heated body at a given temperature which travels through a medium without heating it Axis of maximum height in a height field. Also applicable to fields of pressure, temperature, etc. Rate of decrease of temperature of a parcel of saturated air when it is lifted or lowered from its existing level moist adiabatically Radiation at wavelengths shorter than those in visual range in the energy spectrum A line which is tangent to the instantaneous wind direction at any point in a wind field Belt of latitudes between the tropics and midlatitudes, approximately between 25 and 40◦ of latitude
Subtropical jet
An upper-level jetstream along the boundary between the tropics and the subtropics
Topography
Physical features and contours of elevation of the earth’s surface over a terrain The winds blowing over the tropical belt of the earth’s surface between the ridge of the subtropical high pressure and the equatorial trough of low pressure at low levels; so called, because they were used by early sailors while trading with tropical countries The warm equatorial belt of the earth’s surface, approximately between the 30◦ parallels of latitude Circulation over the tropical belt with tradewinds at low levels, and antitrades at high levels
Tradewinds
Tropics Tropical circulation
Appendix
297
Tropical Convergence Zone (TCZ)
The zone where the NE tradewinds converge into the circulation around the equatorial heat source on its poleward side in the Northern Hemisphere
Troposphere
The radiatively-convectively controlled lowest layer of the earth’s atmosphere in which temperature decreases with height
Tropopause
The level where the troposphere ceases and the temperature instead of decreasing with height remains either constant or increases with height
Trough
A line joining the lowest values in the distribution of an atmospheric property, such as pressure, temperature, etc.
Vortex
An intense cyclonic circulation about an axis perpendicular to the plane of the circulation Tendency of a fluid element in a circulating surface to rotate about an axis normal to the surface in the same direction as the circulation. In solid body rotation, vorticity is a measure of circulation per unit area
Vorticity
Relative vorticity Absolute vorticity Potential vorticity
Walker circulation
Vorticity relative to the earth’s surface Total vorticity of a fluid element obtained by adding relative vorticity to the vorticity of the rotating earth In the absence of friction and heat sources and sinks, the potential vorticity is a materially conservative property of a fluid particle and a function of absolute vorticity divided by the pressure depth between two adjacent potential temperature surfaces Steady forced atmospheric circulation between a heat source and a heat sink placed alternately along the equator
References
Abdullah AJ (1966) The Spiral Bands of a Hurricane: A Possible Dynamic Explanation. J Atmos Sci 23: 367–375 Aberson SD, Montgomery MT, Bell M, Black M (2006) Hurricane Isabel (2003) New Insights into the Physics of Intense Storms Part II. Bull Am Meteor Soc 87: 1349–1354 Akiyama T (1983) An Observational Study of Structure and Maintenance of the Baiu Front, Ph. D. Dissertation, Tokyo University, Tokyo Algue J (1904) The Cyclones of the Far East. Philippine Weather Bur, Manila, 283pp Alisov BP (1954) Die Klimat der Erde. Deutscher Verlag, Berlin Allan RJ (1983) Monsoon and Teleconnections Variability over Australia During the Southern Hemisphere Summers of 1973–1977. Mon Wea Rev 111: 113–142 American Meteorological Society (2000) Meteorological Glossary (2nd ed.). American Meteorological Society, Boston Ananthakrishnan R, Acharya UR, Ramakrishnan AR (1968) On the Criteria for Declaring the Onset of Southwest Monsoon over Kerala. Ind Met Dept Forecasting Manual No. IV-18.1 Andrews J (1932) Rainfall Reliability in Australia. Proc Linn Soc N S W 57: 95–100 Andrews J (1933) Seasonal Incidence and Concentration of Rainfall in Australia. Proc Linn Soc N S W 58: 121–124 Angell JK (1981) Comparison of the Variations in Atmospheric Quantities with Sea Surface Variations in the Equatorial Eastern Pacific. Mon Wea Rev 109: 230–243 Anthes RA (1982) Tropical Cyclones: Their Evolution, Structure and Effects. Am Meteor Soc Meteor Monogr 19: 1–208 Arnason G (1955) Large-Scale Vertical Velocity and Horizontal Divergence. MIT Technical Report, 16 Asnani GC (1993) Tropical Meteorology (vols. I & II). Published by the Author, Pune Badan-Dangon A, Dorman CE, Merrifield MA, Winnant CD (1991) The Lower Atmosphere over the Gulf of California. J Geophys Res 96: 16.877–16.896 Berson FA, Troup AJ (1961) On the Angular Momentum Balance in the Equatorial Trough Zone of the Eastern Hemisphere. Tellus 13: 66–78 Bhalme HN, Mooley DA (1980) Large-Scale Droughts/Floods and Monsoon Circulation. Mon Wea Rev 108: 1197–1211 Bilham EG, Relf EF (1937) The Dynamics of Large Hailstones. Q J R Meteor Soc 63: 149–162 Bjerknes J (1966) A Possible Response of the Atmospheric Hadley Circulation to Equatorial Anomalies of Ocean Temperature. Tellus 18: 820–829 Bjerknes J (1969) Atmospheric Teleconnections from the Equatorial Pacific. Mon Wea Rev 97: 163–172 Black ML, Willoughby HE (1992) The Concentric Eyewall Cycle of Hurricane Gilbert. Mon Wea Rev 120: 947–957 Blackadar AK (1957) Boundary Layer Wind Maxima and Their Significance for the Growth of Nocturnal Inversions. Bull Am Meteor Soc 38: 283–290
299
300
References
Blake DW, Krishnamurti TN, Low-Nam SV, Fein JS (1983) Heat Low over the Saudi Arabian Desert During May 1979 (Summer MONEX). Mon Wea Rev 111: 1759–1775 Blanford HF (1884) On the Connection of the Himalayan Snowfall with Dry Winds and Seasons of Drought in India. Proc R Soc Lond 37: 3 Blanford HF (1886) The Rainfall of India. Indian Meteor Mem 3: 658pp Bonner WD (1968) Climatology of the Low-Level Jet. Mon Wea Rev 96: 833–850 Bonner WD, Paegle J (1970) Diurnal Variation in the Boundary Layer Winds over the South Central United States in Summer. Mon Wea Rev 98: 735–744 Boucher K (1975) Global Climate. Wiley and Sons, New York, p126 Boyle JS, Chen T-J (1987) Synoptic Aspects of the Wintertime East Asian Monsoon. In Monsoon Meteorology (Eds: CP Chang, TN Krishnamurti), Oxford University Press, Oxford, pp. 125–160 Brenner IS (1974) A Surge of Maritime Tropical Air – Gulf of California to the Southwestern United States. Mon Wea Rev 102: 375–389 Bryson RA (1957) The Annual March of Precipitation in Arizona, New Mexico and Northwestern Mexico. Inst Atm Phys Univ Ariz Publ 6: 24pp Bryson RA, Hare FK (1974) The Climates of North America. In World Survey of Climatology (Ed-in-chief: HE Landsberg), vol. 11, chapter 1, pp. 1–47 Bureau of Meteorology (1956) Proceedings of Tropical Cyclone Symposium. Brisbane, p. 1955 Bureau of Meteorology (1978) Australian Tropical Cyclone Forecasting Manual. Australian Bureau of Meteorology, Melbourne Burpee RW (1972) The Origin and Structure of Easterly Waves in the Lower Troposphere of North Africa. J Atmos Sci 29: 77–90 Byers HR, Braham RR (1949) The Thunderstorm. US Government Printing Office, Washington, DC, 287pp Carlson TN (1969a) Synoptic Histories of Three African Disturbances that Developed into Atlantic Hurricanes. Mon Wea Rev 97: 256–276 Carlson TN (1969b) Some Remarks on African Disturbances and Their Progress over the Tropical Atlantic. Mon Wea Rev 97: 716–726 Chang C-P (1970) Westward Propagating Cloud Patterns in the Tropical Pacific as Seen from Time Composite Satellite Photographs. J Atmos Sci 27: 133–138 Chang C-P, Hou SC, Kuo HC, Chen GTJ (1998) The Development of an Intense East Asian Summer Monsoon Disturbance with Strong Vertical Coupling. Mon Wea Rev 126: 2692–2712 Chang C-P, Yi L, Chen GTJ (2000) A Numerical Simulation of Vortex Development During the 1992 East Asian Summer Monsoon Onset Using the Navy’s Regional Model. Mon Wea Rev 128: 1604–1631 Charney JG, Stern ME (1962) On the Stability of Internal Baroclinic Jets in a Rotating Atmosphere. J Atmos Sci 19: 159–172 Charney JG, Eliassen A (1964) On the Growth of the Hurricane Depression. J Atmos Sci 21: 68–75 Charney JG (1975) Dynamics of Deserts and Drought in the Sahel. Q J R Met S 101: 193–202 Charney JG, Quirk WJ, Chow S, Cornfield J (1977) A Comparative Study of the Effects of Albedo Change on Drought in Semi-Arid Regions. J Atmos Sci 34: 1366–1385 Charney JG, Shukla J (1981) Predictability of Monsoons. In Monsoon Dynamics, edited by James Lighthill and R.P. Pearce, pp. 99–109 Chen GTJ, Tsay CY (1978) A Synoptic Case Study of the “Meiyu” near Taiwan. Pap Meteor Res 1: 25–36 Chen GTJ, Chang C-P (1980) The Structure and Vorticity Budget of an Early Summer Monsoon Trough (Meiyu) over Southeastern China and Japan. Mon Wea Rev 108: 942–953 Chen L-X, Li W-L (1981) The Heat Sources and Sinks in the Monsoon Region of Asia. In Proceeding Symposium on Summer Monsoon in South East Asia 1981 Hongzhou, pp. 86–101 (in Chinese with English Abstract). People’s Press of Yunnan Province Kunming Cheng K (1963) China’s Meteorological Services. Today Wea 18: 366–372 Chin PC (1958) Tropical Cyclones in the Western Pacific and China Sea Area from 1884 to 1953. Roy Obsy Hong Kong Tech Mem 7: 84 charts
References
301
Cook C (1964) Korean Weather Service. Central Meteorological Office, Seoul, 41pp Crutcher HL, Meserve JM (1970) Selected Level Heights, Temperatures and Dew-Points for the Northern Hemisphere. NAVAIR-50-IC-52 (Available from Naval Weather Service Command, Washington, DC, 20772) Davidson NE, McBride JL, McAvaney BJ (1983) The Onset of the Winter Monsoon During Winter MONEX – Synoptic Aspects. Mon Wea Rev 111: 496–516 Davidson NE, McBride JL, McAvaney BJ (1984) Divergent Circulations During the Onset of the 1978–1979 Australian Monsoon. Mon Wea Rev 112: 1684–1696 Defant A (1961) Physical Oceanography (vol. 1). Pergamon Press, London Deppermann CE (1947) Notes on the Origin and Structure of Philippine Typhoons. Bull Am Meteor Soc 28: 399–404 Dickson RR (1983) Investigation of a Possible Feedback Loop Involving Eurasian Snow Cover, Indian Monsoon and Southern Oscillation. Proceedings of the 7th Annual Climate Diagnostic Workshop. NOAA, Washington, DC, pp. 437–443 Dickson RR (1984) Eurasian Snow Cover Versus Indian Monsoon Rainfall – An Extension of Hahn-Shukla Results. J Climatol Appl Meteor 23: 171–173 Diercks JW, Anthes RA (1976) A Study of Spiral Bands in a Linear Model of a Cyclonic Vortex. J Atmos Sci 33: 1714–1729 Dines WH (1912) Geophysical Memoirs 2. Meteorological Office, London Dines WH (1919) Geophysical Memoirs 13. Meteorological Office, London Ding Y-H, He S-X (1984) The Mean Circulation in the Tropics in Southeast Asia and West Pacific. Sci Bull 29: 414–416 Ding Y-H (1994) Monsoons over China. Kluwer Academic Publishers, Dondrecht/Boston/London, 419pp Ding Y-H (2004) Seasonal March of the East Asian Summer Monsoon (Ed: C-P Chang). World Sci 2: 3–53 Douglas MW, Maddox RA, Howard K, Reyes S (1993) The Mexican Monsoon. J Clim 6: 1665– 1677 Dunn GE (1940) Cyclogenesis in the Tropical Atlantic. Bull Am Meteor Soc 21: 215–229 Eliassen A (1952) Slow Thermally or Meridionally Controlled Meridional Circulations in a Circular Vortex. Astrophysica Norvegica 5: 2 Eliassen A (1959) On the Formation of Fronts in the Atmosphere the Rossby Memorial Volume. Rockefeller Institute Press, New York Estoque MA (1962) Vertical and Radial Motions in a Tropical Cyclone. Tellus 14: 394–402 Feng Z-Q, Chen L-G, Reiter ER (1984) The Heat Source Distribution in the Atmosphere over the Tibetan Plateau from May to August 1979. Paper Presented at the Third Conference on Mountain Meteorology, Portland Oregon, October 16–19 Ferrel W (1859) The Motions of Fluids and Solids Relative to the Earth’s Surface. Math Monthly 1: 140, 210, 300, 366, 397 Findlater J (1969a) A Major Low Level Air Current over the Indian Ocean During the Northern Summer. Q J R Met Soc 95: 362–380 Findlater J (1969b) Interhemispheric Transport of Air into Lower Troposphere over the Western Indian Ocean. Q J R Met Soc 95: 400–403 Fletcher RD (1945) The General Circulation of the Tropical and Equatorial Atmosphere. J Met 2: 167–174 Flohn H (1960) Equatorial Westerlies over Africa; Their Extension and Significance. Proceedings of Symposium on Tropical Meteorology in Africa (Ed: DJ Bargman). Munitalp Foundation, Nairobi Flohn H (1968) Contributions to Meteorology of the Tibetan Highlands Atmospheric Science Paper No. 130. Department Atmospheric Science, Colorado State University, Fort Collins, Colorado Flores JF, Balagot VF (1969) Climate of the Philippines. In World Survey of Climatology (Ed-inchief: HE Landsberg), Vol. 8: Climates of Northern & Eastern Asia (Ed: H Arakawa), Chap. 3, pp. 159–213
302
References
Fong SK, Wang AY (Eds) (2001) Climatological Atlas for Asian Summer Monsoon. Macau Meteorological and Geophysical Bureau and Macau Foundation, p. 318 Frank NL (1970) Atlantic Tropical Systems of 1969. Mon Wea Rev 98: 307–314 Frank NL (1971) Atlantic Tropical Systems of 1970. Mon Wea Rev 99: 281–285 Frank WM (1977) The Structure and Energetics of the Tropical Cyclone I. Storm Structure. Mon Wea Rev 105: 1119–1135 Freeman JC (1948) An Analogy Between the Equatorial Easterlies and Supersonic Gas Flows. J Met 5: 138–146 Fujita T (1981) Tornadoes and Downbursts in the Context of Generalized Planetary Scales. J Atmos Sci 38: 1511–1534 Fujita TT (1971) Application of ATS-III Photographs for Determination of Dust and Cloud Velocities over the Northern Tropical Atlantic. J Meteor Soc Japan 49: 813–820 Garcia E (1965) Distribucion de la precipitacion en la Republica Mexicana. Publ Inst Geogr 1: 175–191 Galvin JFP (2007) The Subtropical Jetstreams. Weather 62: 295–299 Gao Y-X, Xu S-Y (1962) Some Aspects of East Asian Monsoon. Science Press, Beijing Gao Y-X (1976) The Influence of Qinghai-Xizang Plateau and Land and Sea Distribution on the Climate of China – The Climatic Atlas of Plateau Meteorology, Academia Sinica, pp. 34–46 Gao Y-X (1979) The Phenomena of Qinghai-Xizang Plateau Monsoon. In The Meteorology of the Qinghai-Xizang Plateau, Science Press, Beijing, Chapter 6, 62–69 Gao Y-X, Tang M-C, Luo S-W, Shen Z-B, and Li C (1981) Some Aspects of Recent Research on the Qinghai – Xizang Plateau Meteorology. Bull Am Meteor Soc 62: 31–35 Gaviola E, Fuertes FA (1947) Hail Formation, Vertical Currents, and Icing of Aircraft. J Met 4: 116–120 Gentilli J (1971) Dynamics of the Australian Troposphere. In Climates of Australia and New Zealand in World Survey of Climatology (Ed-in-chief: HE Landsberg), vol. 13, Elsevier Publishing Co., Amsterdam, pp. 53–117 Gentry RC (1964) A Study of Hurricane Rainbands. National Hurricane Research Project Report no. 69, US Department of Commerce, Washington, DC, 85pp Gherzi E (1930) Typhoons in 1928, etc. Zikawei Observatory Shanghai, pp. 26–60, annually Gill AE (1980) Some Simple Solutions for Heat-Induced Tropical Circulation. Q J R Met Soc 106: 447–462 Gill AE (1982) Ocean-Atmosphere Dynamics. International Geophysics Series 30, Academic Press, Inc., New York Gray WM (1968) Global View of the Origin of Tropical Disturbances and Storms. Mon Wea Rev 96: 669–700 Gray WM, Shea DJ (1973) The Hurricane’s Inner Core Region II. Dynamics and Thermo-Dynamic Characteristics. J Atmos Sci 30: 1565–1576 Gray WM (1979) Hurricanes: Their Formation, Structure and Likely Role in the Tropical Circulation. In Meteorology over the Tropical Oceans (Ed: DB Shaw) Royal Meteorological Society, Bracknell, pp. 155–218 Griffiths JF (1972) Climates of Africa in World Survey of Climatology (Ed-in-chief: HE Landsberg), Elsevier Publishing Co., Amsterdam, vol. 10, pp. 1–604 Hadley G (1735) Concerning the Cause of the General Trade Winds. Phil Trans R Soc Lond 39: 58–62 Hahn DG, Shukla J (1976) An Apparent Relationship Between Eurasian Snow Cover and Indian Monsoon Rainfall. J Atmos Sci 33: 2461–2463 Hales JE, Jr (1972) Surges of Maritime Tropical Air Northward over the Gulf of California. Mon Wea Rev 100: 298–306 Hales JE, Jr (1974) Southwestern United States Summer Monsoon Source – Gulf of Mexico or Pacific Ocean. J Appl Meteor 12: 331–342
References
303
Halley E (1686) An Historical Account of the Trade Winds, and Monsoons, Observable in the Seas Between and near the Tropics, with an Attempt to Assign the Phisical Cause of the Said Winds. Phil Trans R Soc Lond 16: 153–168 Hastenrath S, Heller L (1977) Dynamics of Climate Hazards in Northeast Brazil. Q J R Meteor Soc 103: 77–92 Hawkins HF, Imbembo SM (1976) The Structure of a Small, Intense Hurricane, INEZ 1966. Mon Wea Rev 104: 418–442 Helfand HM, Schubert SD (1995) Climatology of the Great Plains Low-Level Jet and Its Contribution to the Continental Moisture Budget of the United States. J Clim 8: 784–806 Higgins RW, Yao Y, Yarosh ES, Janowiak JE, Mo KC (1997) Influence of the Great Plains LowLevel Jet on Summertime Precipitation and Moisture Transport over the Central United States. J Clim 10: 483–507 Hoecker WJ (1963) Three Southerly Low-Level Jet Systems Delineated by the Weather Bureau Special Pibal Network of 1961. Mon Wea Rev 91: 573–582 Holland GJ (1984a) On the Climatology and Structure of Tropical Cyclones in the Australian/Southwest Pacific Region, 1 Data and Tropical Storms. Austr Meteor Mag 32: 1–15 Holland GJ (1984b) On the Climatology and Structure of Tropical Cyclones in the Australian/Southwest Pacific Region II Hurricanes. Austr Meteor Mag 32: 17–31 Holland GJ (1984c) On the Climatology and Structure of Tropical Cyclones in the Australian/Southwest Pacific Region, III Major Hurricanes. Austr Meteor Mag 32: 33–46 Holland GJ, Nicholls N (1985) A Simple Predictor of El Nino? Trop Ocean-Atmos News-Lett 30: 8–9 Holton JR (1971) A Diagnostic Model for Equatorial Wave Disturbances: The Role of Vertical Shear of the Mean Zonal Wind. J Atmos Sci 28: 55–64 Holton JR (1979) An Introduction to Dynamic Meteorology (2nd ed.). Academic Press, New York, 391pp Hope JR, Neumann CJ (1970) An Operational Technique for Relating the Movement of Existing Tropical Cyclones to Past Tracks. Mon Wea Rev 98: 925–933 Hovermale JB, Livezey RE (1977) Three-Year Performance Characteristics of the NMC Hurricane Model Preprints 11th Technical Conference on Hurricanes and Trop Met Miami Amer Met Soc, pp. 122–124 Huges LA (1952) On the Low-Level Wind Structure of Tropical Storms. J Met 9: 422–428 Humphreys WJ (1940) Physics of the Air. McGraw-Hill, New York, 359pp India Meteorological Department (1943) Climatological atlas for Airmen-3 India. Meteorological Department, New Delhi Jelesnianski CP (1965) A Numerical Calculation of Storm Tides Induced by a Tropical Storm Impinging on a Continental Shelf. Mon Wea Rev 93: 343–358 Jelesnianski CP (1966) Numerical Computations of Storm Surges Without Bottom Stress. Mon Wea Rev 94: 740–756 Jelesnianski CP (1967) Numerical Computations of Storm Surges with Bottom Stress. Mon Wea Rev 95: 740–756 Johnson DH, Morth HT (1960) Forecasting Research in Africa. In Tropical Metorology in Africa (Ed: DJ Bargman), Munitalp Foundation, Nairobi, pp. 56–137 Jordan CL (1958) Mean Soundings for the West Indies Area. J Meteor 15: 91–97 Jordan CL (1961) Marked Changes in the Characteristics of the Eye of Intense Typhoons Between the Deepening and Filling Stages. J Meteor 18: 779–789 Kalnay E, Dimego G, Lord S, Kanamitsu M, Leetmaa A, Rao DB (1994) NMC Modeling and Data Assimilation Plans for 1994–1998. Preprints 10th Conference on Numerical Weather Prediction. American Meteorological Society, Portland, OR, pp. 143–148 Kanamitsu M, Krishnamurti TN (1978) Northern Summer Tropical Circulations During Drought and Normal Rainfall Months. Mon Wea Rev 106: 331–347 Keenan TD (1981) An Error Analysis of Objective Tropical Cyclone Forecasting Schemes Used in Australia. Austr Meteor Mag 29: 133–142 Keenan TD (1982) A Diagnostic Study of Tropical Cyclone Forecasting in Australia. Austr Meteor Mag 30: 69–80
304
References
Kendrew WG. (1953) The Climates of the Continents. Clarendon Press, Oxford, pp. 28–146 Khromov SP (1957) Die geographische verbreitung der Monsune. Petermanns Geogr Mitt 101: 234–237 Kiladis GN, Storch V, Loon V (1989) Origin of the South Pacific Convergence Zone. J Clim 2: 1185–1195 Klein WH (1957) Principal Tracks and Frequencies of Cyclones and Anticyclones in the Northern Hemisphere. US Weather Bureau Research Paper 40, 60pp Koteswaram P, George CA (1958) On the Formation of Monsoon Depressions in the Bay of Bengal. Ind J Met Geophys 9: 9–22 Kousky VE, Gan MA (1981) Upper Tropospheric Cyclonic Vortices in the Tropical South Atlantic. Tellus 33: 533–551 Krishnamurti TN (1961) On the Vertical Velocity Field in a Steady Symmetric Hurricane. Tellus 13: 171–180 Krishnamurti TN, Baumhefner D (1966) Structure of a Tropical Disturbance Based on Solutions of a Multi-Level Baroclinic Model. J Appl Meteor 5: 396–406 Krishnamurti TN, Kanamitsu M, Godbole R, Chang CB, Carr F, Chow JH (1976) Study of a Monsoon Depression(II) – Dynamical Structure. J Meteor Soc Japan 54: 208–225 Krishnamurti TN, Ardanuy P, Ramanathan Y, Pasch R (1981) On the Onset Vortex of the Summer Monsoon. Mon Wea Rev 109: 344–363 Krishnamurti TN, Subrahmanyam D (1982) The 30 to 50 Day Mode at 850 mb During MONEX. J Atmos Sci 39: 2088–2095 Krishnamurti TN, Surgi N (1987) Observational Aspects of Summer Monsoon. In Monsoon Meteorology (Eds: CP Chang, TN Krishnamurti), Oxford University Press, Oxford, pp. 3–25 Kuo HL (1959) Dynamics of Convective Vortices and Eye Formation. The Rossby Memorial Volume. Rockefeller Institute Press, New York Kurihara Y (1976) On the Development of Spiral Bands in a Tropical Cyclone. J Atmos Sci 33: 940–958 Kurihara Y, Bender MA, Tulega RE, Ross RJ (1993) Hurricane Forecasting with the GFDL Automated Prediction System. Preprints 20th Conference on Hurricanes and Tropical Meteorology. American Meteorological Society, San Antonio, TX, pp. 323–326 Lajoie FA, Butterworth, IJ (1984) Oscillations of High-Level Cirrus and Heavy Precipitation Around Australian Region Tropical Cyclones. Mon Wea Rev 112: 535–544 Landers H (1955) A Three-Dimensional Study of the Horizontal Velocity Divergence. J Met 12: 415–427 Lau K-M, Chang C-P (1987) Planetary Scale Aspects of the Winter Monsoon and Atmospheric Teleconnections. In Monsoon Meteorology (Eds: CP Chang, TN Krishnamurti), Oxford University Press, Oxford, pp. 161–202 Lau K-M, Wu HT, Yang S (1998) Hydrological Process Associated with the First Transition of the Asian Summer Monsoon: A Pilot Study. Bull Am Meteor Soc 79: 1871–1882 Lau K-M, Ding Y-H, Wang JT, Johnson R, Keenan T, Cifelli R, Gerlach J, Thiele O, Rikenbach T, Tsay SC, Lin PH (2000) A Report of the Field Operation and Early Results of the South China Sea Monsoon Experiment (SCSMEX). Bull Am Meteor Soc 81: 1261–1270 Lettau H (1967) Small to Large-Scale Features of Boundary Layer Structure over Mountain Slopes. Proceedings of Symposium on Mountain Meteorology. Atmospheric Science Paper 122, Colorado State University, 221pp Loewe F, Radok U (1950) A Meridional Aerological Cross-Section in the Southwest Pacific. J Met 7: 58–76 (Amendments 305–306) Lord SJ (1991) A Bogussing System for Vortex Circulations in the NMC Global Forecast Model Season Preprints, 19th Conference on Hurricanes and Tropical Meteorology. American Meteorological Society, Miami, FL, pp. 328–330 Love G, Garden G (1984) The Australian Monsoon of January 1974. Austr Meteor Mag 32: 185–194 Love G (1985a) Cross-Equatorial Influence of Winter Hemisphere Cold Surges. Mon Wea Rev 113: 1487–1498
References
305
Love G (1985b) Cross-Equatorial Interactions During Tropical Cyclogenesis. Mon Wea Rev 113: 1499–1509 Luo H, Yanai M (1984) The Large-Scale Circulation and Heat Sources over the Tibetan Plateau and Surrounding Areas During the Early Summer of 1979: Heat and Moisture Budgets. Mon Wea Rev 112: 966–989 Madden RA, Julian PR (1971) Detection of a 40–50 Day Oscillation in the Zonal Wind in the Tropical Pacific. J Atmos Sci 28: 702–708 Madden RA, Julian PR (1972) Description of Global-Scale Circulation Cells in the Tropics with a 40–50 Day Period. J Atmos Sci 29: 1109–1123 Madden RA (1986) Seasonal Variations of the 40–50 Day Oscillation in the Tropics. J Atmos Sci 43: 3138–3158 Malkus JS, Riehl H (1960) On the Dynamics and Energy Transformations in Steady-State Hurricanes. Tellus 12: 1–20 Mathur MB (1991) The National Meteorological Center’s Quasi-Lagrangian Model for Hurricane Prediction. Mon Wea Rev 119: 1419–1447 Matsuno T (1966) Quasi-Geostrophic Motions in the Equatorial Area. J Meteor Soc Japan 44: 25–42 Matsumoto S, Ninomiya K (1971) On the Mesoscale and Medium-Scale Structure of a Cold Front and the Relevant Vertical Circulation. J Meteor Soc Japan 49: 648–662 Matsumoto S (1972) Unbalanced Low-Level Jet and Solenoidal Circulation Associated with Heavy Rainfalls. J Meteor Soc Japan 50: 194–203 Maura AD, Shukla J (1981) On the Dynamics of Drought in Northeast Brazil: Observations, Theory and Numerical Experiments with a General Circulation Model. J Atmos Sci 38: 2653–2675 McBride JL, Zehr R (1981) Observational Analysis of Tropical Cyclone Formation – Part II: Comparison of Non-Developing Versus Developing Systems. J Atmos Sci 38: 1132–1151 McBride JL, Keenan TD (1982) Climatology of Tropical Cyclone Genesis in the Australian Region. J Climatol 2: 13–33 McBride JL (1983) Satellite Observations of the Southern Hemisphere Monsoon During Winter MONEX. Tellus 35A: 189–197 McBride JL, Nicholls N (1983) Seasonal Relationships Between Australian Rainfall and the Southern Oscillation. Mon Wea Rev 111: 1998–2004 McBride JL (1987) The Australian Summer Monsoon. In Monsoon Meteorology (Eds: CP Chang, TN Krishnamurti), Oxford University Press, Oxford, pp. 203–231 McCune S (1941) Climatic Regions of Korea and Their Economy. Geograph Rev 31: 95–99 McNoldy BD (2004) Triple Eyewall in Hurricane Juliette. Bull Am Meteor Soc 85: 1663–1666 Mean Annual Rainfall of Australia during Period, 1911–1940 (1962). Bureau of Meteorology, Australia Means LL (1952) On Thunderstorm Forecasting in the Central United States. Mon Wea Rev 80: 165–189 Miller BI, Chase PP (1966) Predictions of Hurricane Motion by Statistical Methods. Mon Wea Rev 94: 399–405 Miller BI (1967) Characteristics of Hurricanes. Science 157: 1389–1399 Montgomery MT, Bell MM, Aberson SD, Black M L (2006) Hurricane Isabel (2003): New Insights into the Physics of Intense Storms. Part 1. Bull Am Meteor Soc 87: 1335–1347 Mooley DA (1975) Vagaries of the Indian Summer Monsoon During the Last 10 Years. Vayu Mandal 5: 65–66 Mooley DA (1976) Worst Summer Monsoon Failures over the Asiatic Monsoon Area. Proceedings of Symposium Droughts in the Asiatic Summer Monsoon Area, Pune December 1972. Ind Nat Sci Acad, New Delhi Mooley DA, Parthasarathy B (1983) Droughts and Floods in India in Summer Monsoon Seasons, 1871–1980. In Proceedings of Symposium on Variations of the Global Water Budget (Eds: AS Perrot, M Beran, R Ratcliffe), Oxford August 1981, D. Reidel, Dodrecht, Holland, pp. 239–252
306
References
Mooley DA, Parthasarathy B (1984) Indian Summer Monsoon and East Equatorial Pacific Sea Surface Temperature. Atmos-Ocean 22(1): 23–35 Mooley DA, Shukla J (1987) Variability and Forecasting of the Summer Monsoon Rainfall over India. In Monsoon Meteorology (Eds: CP Chang, TN Krishnamurti), Oxford University Press, Oxford, pp. 26–59 Mosino Aleman PA, Garcia E (1974) The Climate of Mexico. In World Survey of Climatology (Ed-in-chief: HE Landsberg), Elsevier Publishing Co., Amsterdam, vol. 11, chap. 4, pp. 345– 404 Murakami T (1976) Cloudiness Fluctuations During the Summer Monsoon. J Meteor Soc Japan 54: 175–181 Murakami T (1981a) Orographic Influence of the Tibetan Plateau on the Asiatic Winter Monsoon Circulation Part-1: Large-Scale Aspects. J Meteor Soc Japan 59: 40–65 Murakami T (1981b) Orographic Influence of the Tibetan Plateau on the Asiatic Winter Monsoon Circulation Part-III: Short-Period Oscillations. J Meteor Soc Japan 59: 173–200 Murakami T, Ding Y-H (1982) Wind and Temperature Changes over Eurasia During the Early Summer of 1979. J Meteor Soc Japan 60: 183–196 Murakami T, Sumi A (1982) Southern Hemisphere Summer Monsoon Circulation During the 1978–1979 WMONEX. Part 1 Monthly Mean Wind Fields. J Meteor Soc Japan 60: 638–648 Murakami M (1983) Analysis of the Deep Convective Activity over the Western Pacific and SouthEast Asia, Part 1 – Diurnal Variation. J Met Soc Japan 61: 60–76 Murakami M (1984) Analysis of the Deep Convective Activity over the Western Pacific and SouthEast Asia, Part II – Seasonal and Intraseasonal Variations During Northern Summer. J Met Soc Japan 62: 88–108 Murakami M (1987a) Satellite Cloudiness in the Monsoon Area. In Monsoon Meteorology (Eds: CP Chang, TN Krishnamurti), Oxford University Press, Oxford, chapter 11, pp. 354–402 Murakami T (1987) Effects of the Tibetan Plateau. In Monsoon Meteorology (Eds: CP Chang, TN Krishnamurti), Oxford University Press, Oxford, chapter 8, pp. 235–270 Namias J (1972) Influence of Northern Hemisphere General Circulation on Drought in Northeast Brazil. Tellus 24: 336–343 NAVAIR (1966) Components of the 1000-mb Winds (or Surface Winds) of the Northern Hemisphere NAVAIR Document 50-IC-51, WASH, 74 Charts Neumann CJ (1972) An Alternate to the HURRAN Tropical Cyclone Forecast System. NOAA Tech Memo NWS SR-62, 22pp Neumann CJ, Hope JR, Miller BI (1972) A Statistical Method of Combining Synoptic and Empirical Tropical Cyclone Prediction Systems. NOAA Tech Memo NWS SR-63, 32pp Neumann CJ, Lawrence MB (1975) An Operational Experiment in the Statistical-Dynamical Prediction of Tropical Cyclone Motion. Mon Wea Rev 103: 665–673 Neumann CJ, Pelissier JM (1981a) Models for the Prediction of Tropical Cyclone Motion over the North Atlantic: An Operational Evaluation. Mon Wea Rev 109: 522–538 Neumann CJ, Pelissier JM (1981b) An Analysis of Atlantic Tropical Cyclone Forecast Errors, 1970–1971. Mon Wea Rev 109: 1248–1266 Newton CW (1966) Circulations in Large Sheared Cumulonimbus. Tellus 18: 699–713 Nicholls N, McBride JL, Ormerod RJ (1982) On Predicting the Onset of the Australian Wet Season at Darwin. Mon Wea Rev 110: 14–17 Nicholls N (1984b) The Southern Oscillation, Sea Surface Temperature and Internal Fluctuations in Australian Tropical Cyclone Activity. J Climatol 4: 661–670 Ninomiya K, Akiyama T (1971) The Development of the Medium-Scale Disturbance in the Baiu Front. J Meteor Soc Japan 49: 663–677 Ninomiya K, Akiyama T (1974) Band Structure of Meso-Scale Echo Cluster Associated with LowLevel Jet Stream. J Meteor Soc Japan 52: 300–313 Ninomiya K (1978) Heavy Rainfalls Associated with Frontal Depression. In Asian Subtropical Humid Region Part-1. J Meteor Soc Japan 56: 253–266
References
307
Ninomiya K, Yamazaki K (1979) Heavy Rainfall Associated with Frontal Depression in Asian Subtropical Humid Region Part II. J Meteor Soc Japan 57: 399–413 Ninomiya K, Koga H, Yamagishi Y, Tatsumi Y (1984) Prediction Experiment of Extremely Intense Rainstorm by a Very Fine-Mesh Primitive Equation Model. J Meteor Soc Japan 62: 273–295 Ninomiya K, Murakami T (1987) The Early Summer Rainy Season (Baiu) over Japan. In Monsoon Meteorology (Eds: CP Chang, TN Krishnamurti), Oxford University Press, Oxford, pp. 93–121 Ninomiya K, Akiyama T (1992) Multi-Scale Features of Baiu, the Summer Monsoon of Japan and the East Asia. J Meteor Soc Japan 70: 467–495 Nitta T (1970) Statistical Study of Tropospheric Wave Disturbances in the Tropical Pacific Region. J Meteor Soc Japan 48: 47–60 Nobre P, Shukla J (1996) Variations of Sea Surface Temperature, Wind Stress, and Rainfall over the Tropical Atlantic and South America. J Clim 9: 2464–2479 Norquist DC, Recker EE, Reed RJ (1977) The Energetics of African Wave Disturbances as Observed During Phase III of GATE. Mon Wea Rev 105: 334–342 Paegle J, Nogues-Paegle J (1997) Dynamic and Thermodynamic Aspects of South American LowLevel Jets. Preprints of Fifth International Conference on Southern Hemisphere Meteorology and Oceanography Pretoria, South Africa, 7–11 April 1997, pp. 326–327 Palmen E (1951) Mean Meridional Circulation Cells over the Globe in Winter. Q J R Met Soc 77: 337 Palmer CE (1952) Tropical Meteorology. Q J R Met Soc. 78: 126–163 Parthasarathy B (1984) Some Aspects of Large-Scale Fluctuations in the Summer Monsoon Rainfall over India During 1871–1978. Ph. D. Dissertation, University of Poona, Pune (India) Paterson LA, Bate PW (2001) The South Pacific and Southeast Indian Ocean Tropical Cyclone Season, 1999–2000. Austr Meteor Mag 50: 123–135 Pearce EA, Smith G (1990) The Times Books World Weather Guide, published by Times Books. Random House, Inc., New York Petterssen S (1956) Weather Analysis and Forecasting. Vol. 1: Motion and Motion Systems. McGraw-Hill Book Co. Inc., New York Pittock AB (1984) On the Reality, Stability and Usefulness of Southern Hemisphere Teleconnections. Austr Meteor Mag 32: 75–82 Portig WH (1976) The Climate of Central America. In World Survey of Climatology (Ed-in-chief: HE Landsberg), Elsevier Publishing Co., Amsterdam, vol. 12, chapter 7, pp. 405–478 Ramage CS (1966) The Summer Monsoon Circulation over the Arabian Sea. J Atmos Sci 23: 144–150 Ramage CS (1968) Role of a Tropical Maritime Continent in the Atmospheric Circulation. Mon Wea Rev 96: 365–370 Ramage CS (1971) Monsoon Meteorology. Academic Press, New York, pp. 1–296 Ramage CS, Raman CRV (1972) Meteorological Atlas of the International Indian Ocean Expedition. Vol. 2: Upper Air. US Government Printing Office, Washington, DC Ramamurty K (1969) Some Aspects of the ‘Breaks’ in the Indian Southwest Monsoon During July and August in Forecasting Manual, Part IV. India Meteorological Department, Pune, India Ramos RPL (1975) Precipitation Characteristics in the Northeast Brazil Dry Region. J Geophys Res 80: 1665–1678 Rao YP (1964) Interhemispheric Circulation. Q J R Met Soc 90: 190–194 Rao YP (1976) ‘Southwest Monsoon’ Monograph 1/76. India Meteorological Department, Pune, India Rao YP (1981) The Climate of the Indian Subcontinent. In World Survey of Climatology (Ed-inchief: HE Landsberg), Elsevier Publishing Co., Amsterdam, vol. 9, chapter 2, pp. 67–118 Rasmussen EM, Carpenter TH (1983) The Relationship Between Eastern Equatorial Pacific Sea Surface Temperature and Rainfall over India and Sri Lanka. Mon Wea Rev 111: 517–528 Reed RJ, Recker EE (1971) Structure and Properties of Synoptic-Scale Wave Disturbances in the Equatorial Western Pacific. J Atmos Sci 28: 1117–1133
308
References
Reed RJ, Norquist DC, Recker EE (1977) The Structure and Properties of African Wave Disturbances as Observed During Phase-III of GATE. Mon Wea Rev 105: 317–333 Rennick MA (1976) The Generation of African Waves. J Atmos Sci 33: 1955–1969 Riehl H (1945) Waves in the Easterlies and the Polar Front in the Tropics. Department of Meteorological, University of Chicago, Misc Rept 17, 79pp Riehl H (1954) Tropical Meteorology. McGraw-Hill, New York Riehl H (1977) Venezuelan Rain Systems and the General Circulation of the Summer Tropics, II: Relations Between Low and High Latitudes. Mon Wea Rev 105: 1421–1433 Rosen RD, Salstein DA (1981) Variations in Atmospheric Angular Momentum, 1 January 1976–31 December 1980 Rossby CG (1947) The Distribution of Angular Velocity in Gaseous Envelopes under Influence of Large-Scale Horizontal Mixing Processes. Bull Am Meter Soc 28: 53–68 Saha KR (1962) On the Formation of Large Hailstones. Ind J Met Geophys 13: 1–5 Saha KR (1970) Air and Water Vapour Transport Across the Equator in Western Indian Ocean During Northern Summer. Tellus 22: 681–687 Saha KR (1978) Some Recent Studies Concerning Tropical Monsoons in India. In Climate Change and Food Production (Eds: K Takahashi, MM Yoshino), University of Tokyo Press, Tokyo, pp. 401–409 Saha S, Saha KR (1980) A Hypothesis on Onset, Advance and Withdrawal of the Indian Summer Monsoon. Pure Appl Geophys (PAGeoph) 118: 1066–1075 Saha K, Sanders F, Shukla J (1981) Westward Propagating Predecessors of Monsoon Depressions. Mon Wea Rev 109: 330–343 Saha K, Chang C-P (1983) The Baroclinic Processes of Monsoon Depressions. Mon Wea Rev 111: 1506–1514 Saha K, Saha S (1993a) A Diagnostic Analysis of MONEX-1979 Onset Vortex over the Arabian Sea. Mausam 44: 321–328 Saha S, Saha KR (1996) Structure and Properties of Summer Monsoon Stationary Wave over Southern Asia: An Observational Study. Mausam 47: 133–144 Saha KR, Van den Dool HM, Saha S (1998) A View of Global Monsoons in Observed and Model Climatology. Mausam 49: 79–94 Saha K, Saha S (2000) The Australian Monsoon, Pt. 1: Climatological Features and the Asian Connection. Mausam 51: 127–154 Saha K, Saha S (2001a) The Australian Monsoon, Pt. 2: Depressions and Cyclones. Mausam 52: 333–350 Saha K, Saha S (2001b) African Monsoons, Pt. 1: Climatological Structure and Circulation. Mausam 52: 479–510 Saha K, Saha S (2002) African Monsoons, Part 2: Synoptic-Scale Wave Disturbances in the Intertropical Convergence Zone over North Africa. Mausam 53: 197–214 Saha K, Saha S (2004a) On the Monsoons of South America, Part 1: Climatological Structure and Circulation. Mausam 55: 41–72 Saha K, Saha S (2004b) On the Monsoons of South America, Part 2: Interaction with Extra-Tropical Disturbances and Formation of Tropical Depressions and Upper-Tropospheric Cyclonic Vortices. Mausam 55: 237–256 Saha K, van den Dool HM, Saha S (1994) On the Annual Cycle in Surface Pressure on the Tibetan Plateau compared to Its Surroundings. J Clim 7: 2014–2019 Sanders F, Burpee RW (1968) Experiments in Barotropic Hurricane Track Forecasting. J Appl Meteor 7: 313–323 Sanders F, Gyakum JR (1980) Synoptic-Dynamic Climatology of the “Bomb”. Mon Wea Rev 108: 1589–1606 Sanders F, Pike AC, Gaertner (1975) A Barotropic Model for Operational Tracks of Tropical Storms. J Appl Meteor 14: 265–280 Sands RD (1959) A Study in the Regional Climatology of Mexico with Precipitation as the Correlative Factor (Doctor’s Thesis) Clarke University, Worcestor, MA, 134pp
References
309
Satyamurti P, Nobre CA, Silva Dias PL (1998) South America. In Meteorology of the Southern Hemisphere (Eds: DJ Karoly, DG Vincent), American Meteorological Society Meteorological Monographs 27, no. 49, American Meteorological Society, Boston Schumann TEW (1938) The Theory of Hail Formation. Q J R Meteor Soc 64: 3–21 Schwerdtfeger W (1976) Climates of Central and South America. In World Survey of Climatology (Ed-in-chief: HE Landsberg), vol. 12, Elsevier Scientific Publishing Co., Amsterdam Selga M, Reppetti WC, Adams W (1931) Oceanographic Papers, vol. III (1–10), Bureau of Printing, Manila, 210pp Senn HV, Hiser HW, Bourret RC (1957) Studies of Hurricane Spiral Bands as Observed on Radar. Nat Hurr Res Project Rept No. 12, 13pp Senn HV, Hiser HW (1959) On the Origin of Hurricane Spiral Rain Bands. J Meteor 16: 419–426 Sheets RC (1969) Some Mean Hurricane Soundings. J Appl Meteor 8: 134–146 Sheets RC (1980) Some Aspects of Tropical Cyclone Modification. Austr Meteor Mag 27: 259– 280 Shukla J (1978) CISK-Barotropic-Baroclinic Instability and the Growth of Monsoon Depression. J Atmos Sci 35: 495–508 Shukla J (1979) Summary of a Lecture by Jule Charney, Prepared by J. Shukla and Published in the Proceedings of a Seminar on the Impact of GATE on Large-Scale Numerical Modeling of the Atmosphere and Ocean, August 24–28, 1979, pp. 166–169 Shukla J, Mintz Y (1982) The Influence of Land Surface Evapo-Transpiration on the Earth’s Climate. Science 214: 1498–1501 Shukla J, Paolino D (1983) The Southern Oscillation and Long-Range Forecasting of Summer Monsoon Rainfall over India. Mon Wea Rev 111: 1830–1837 Sikka DR (1980a) Some Aspects of large-Scale Fluctuations of Summer Monsoon Rainfall over India in Relation to Fluctuations in the Planetary and Regional Scale Circulation Parameters. Proc Ind Acad Sci (Earth and Planetary Sci) 89: 179–195 Sikka DR (1980b) Southern Hemispheric Influences and the Outset of Southwest Monsoon of 1979. FGGE Operations Rep 9(B): 23–26 Sikka DR, Gadgil S (1980) On the Maximum Cloud Zone and the ITCZ over Indian Longitudes During the Southwest Monsoon. Mon Wea Rev 108: 1840–1853 Sikka DR, Gray WM (1981) On the linkage of the genesis of the monsoon disturbances and cyclones in the north Indian Ocean with the passage of baroclinic waves across the south Indian Ocean. International Conference on the Scientific Results of the Monsoon Experiment, extended abstracts, Denpasar Bali Indonesia Simpson GC (1921) The Southwest Monsoon. Q J R Meteor Soc 47: 151–172 Simpson RH (1974) The Hurricane Disaster – Potential Scale. Weatherwise 27: 169–186 Smith RK (1980) Tropical Cyclone Eye Dynamics. J Atmos Sci 37: 1227–1232 Soliman K H (1958) On the Intertropical Front and the Intertropical Convergence Zone over Africa and Adjacent Oceans. Proceedings of Symposium on Monsoons of the World, New Delhi, pp. 135–142 Sud YC, Smith WE (1985) Influence of Local Land Surface Processes on the Indian Monsoon – A Numerical Study. J Climatol Appl Meteor 24: 1015–1036 Sukanto M (1969) Climate of Indonesia. In World Survey of Climatology (Ed-in-chief: HE Landsberg), Elsevier Publishing Co., Amsterdam, vol. 8, Chap. 4, pp. 215–229 Sumi A. Murakami T (1981) Large-Scale Aspects of the 1978–1979 Winter Circulation over the Greater WMONEX Region. Part 1: Monthly and Seasonal Mean Fields. J Meteor Soc Japan 59: 625–645 Sun Y (2002) A study of 1998 anomalous summer monsoon activity and its mechanism. Ph. D. Thesis, 280 pp. Available from National Climate Center, CMA Beijing (in Chinese) Sutcliffe RC (1947) A Contribution to the Problem of Development. Q J R Meteor Soc 73: 370–383 Tang MC (1979) The Pressure and Wind. In Meteorology of the Qinghai-Xizang Plateau. Science Press, Beijing, Chapter 3, pp. 23–28 Tang MC, Reiter ER (1984) Plateau Monsoon of the Northern Hemisphere: A Comparison Between North America and Tibet. Mon Wea Rev 112: 617–637
310
References
Taljaard JJ (1972) Synoptic Meteorology of the Southern Hemisphere (Ed: CW Newton), Meteorological Monographs no. 13, American Meteorological Society, Boston Tannehill IR (1927) Some Inundations Attending Tropical Cyclones. Mon Wea Rev 55: 453–456 Tao S-Y (Ed) (1980) Severe Rainstorms in China. Science Press, Beijing (in Chinese) Tao S-Y, Ding Y-H (1981) Observational Evidence of the Influence of the Qinghai Xizang (Tibet) Plateau on the Occurrence of Heavy Rain and Severe Convective Storms in China. Bull Am Meteor Soc 62: 23–30 Tao S-Y, Chen L (1987) A Review of Recent Research on the East Asian Summer Monsoon in China. In Monsoon Meteorology (Eds: CP Chang, TN Krishnamurti), Oxford University Press, Oxford Tech. Repr A345-T1(NTIS N82-11690) Environmental Research and Technology Inc., Concord, MA Tepper M (1958) A Theoretical Model for Hurricane Radar Bands. Preprints Seventh Weather Radar Conference. American Meteorological Society, Miami, pp. 56–65 Thomson J (1857) On the grand currents of atmospheric circulation. British Assoc Meeting Dublin (Unpublished). See Published paper at (1892) Phil Trans R Soc Lond (A) 183: 653–684 Thompson BW (1965) The Climate of Africa. Oxford University Press, London and New York Trewartha GT (1961) The Earth’s Problem Climates. University of Wisconsin Press, Madison, WI Troup AJ (1961) Variations in Upper Tropospheric Flow Associated with the Onset of the Australian Summer Monsoon. Ind J Met Geo Phys 12: 217–230 Van den Dool HM, Saha S (1993) On the Seasonal Re-distribution of Atmospheric Mass in a General Circulation Model. J Clim 6: 22–30 Virji H (1981) A Preliminary Study of Summertime Tropospheric Circulation Patterns over South America Estimated from Cloud Winds. Mon Wea Rev 109: 599–610 Vulquin A (1971) Arguments en faveur d’une mousson en Amazonie. Tellus XXIII: 74–80 Walisser DE, Gautier C (1993) A Satellite-Derived Climatology of the ITCZ. J Clim 6: 2162–2174 Walker GT (1910a) On the Meteorological Evidence for Supposed Changes of Climate in India. Mem India Meteor Dept 21: 1–21 Walker GT (1910b) Correlations in Seasonal Variations of Weather II. Mem India Meteor Dept 21: 22–45 Walker GT (1914) The Liability to Drought in India as Compared with that in Other Countries. Mem Indian Meteor Dept 21: 1–9 Walker GT (1915a) Correlation in Seasonal Variation of Weather IV. Sunspots and Rainfall. Mem Indian Meteor Dept 21(X): 17–60 Walker GT (1915b) Correlation in Seasonal Variation of Weather V. Sunspots and Temperature. Mem Indian Meteor Dept 21(X): 61–90 Walker GT (1915c) Correlation in Seasonal Variation of Weather VI. Sunspots and Pressure. Mem Indian Meteor Dept 21(XII): 91–118 Walker GT (1924) World Weather IX Mem. Indian Meteor Dept 24: 275–332 Wallace JM, Chang C-P (1969) Spectrum Analysis of Large-Scale Wave Disturbances in the Tropical Lower Troposphere. J Atmos Soc 26: 1010–1025 Wallace JM (1971) Spectral Studies of Tropospheric Wave Disturbances in the Tropical Western Pacific. Rev Geophys Space Phys 9: 557–612 Wallace JM, Hobbs PV (1977) Atmospheric Science – An Introductory Survey. Academic Press, New York, 467pp Wang J-Z, Leftwich PW (1984) A Major Low-Level Cross-Equatorial Current at 110E During the Northern Summer and Its Relation to Typhoon Activities. Sci Atmos Sin 8: 443–449 (in Chinese with English abstract) Wang B, Lin H (2002) Rainy Season of the Asia-Pacific Summer Monsoon. J Clim 15: 386–398 Watts IEM (1969) Climates of China and Korea. In World Survey of Climatology (Ed-in-chief: HE Landsberg), Elsevier Publishing Co., Amsterdam, vol. 8, pp. 1–117 Webster PJ, Magana VO, Palmer TN, Shukla J, Tomas RA, Yanai M, Yasunari T (1998) Monsoons: Processes, Predictability, and the Prospects for Prediction. J Geophys Res 103: 14451–14510
References
311
Wegener A (1911) Thermodynamics of the Atmosphere. J.A. Barth, Leipzig, p. 257 Wexler H (1947) Structure of Hurricanes as Determined by Radar. Ann NY Acad Sci 48: 821–844 Wiesnet DR, Matson M (1976) A Possible Forecasting Technique for Winter Snow Cover in Northern Hemisphere and Eurasia. Mon Wea Rev 104: 828–835 Williams KT (1970) Characteristics of the Wind, Thermal and Moisture Fields Surrounding the Satellite-Observed Meso-Scale Trade Wind Cloud Clusters in the Western North Pacific. Preprints of Papers, Symposium on Tropical Meteorology. American Meteorological Society, Honolulu, pp. D IV-1–D IV-6 Willoughby HE (1977) Inertia-Buoyancy Waves in Hurricanes. J Atmos Sci 34: 1028–1039 Willoughby HE (1978a) A Possible Mechanism for the Formation of Hurricane Rainbands. J Atmos Sci 35: 838–848 Willoughby HE (1978b) The Vertical Structure of Hurricane Rainbands and Their Interaction with the Mean Vortex. J Atmos Sci 35: 849–858 Willoughby HE (1979) Forced Secondary Circulations in Hurricanes. J Geophys Res 84: 3173–3183 World Meteorological Organization (1981) Scientific Results of the Air Mass Transformation Experiment (AMTEX) GARP Publication Series, no. 24 Xie PP, Arkin PA (1996) Analyses of Global Monthly Precipitation Using Gauge Observations, Satellite Estimates and Numerical Model Predictions. J Clim 9: 840–858 Yanai M (1961a) A Detailed Analysis of Typhoon Formation. J Meteor Soc Japan 39: 187–214 Yanai M (1961b) Dynamical Aspects of Typhoon Formation. J Meteor Soc Japan 39: 282–309 Yanai M (1963) A Comment on the Creation of Warm Core in Incipient Tropical Cyclone Typhoon. Res Lab Meteor Res Inst Tokyo Tech. Note 1 Yanai M (1964) Formation of Tropical Cyclones. Rev Geophys 2: 367–414 Yanai M, Nitta T (1967) Computation of Vertical Motion and Vorticity Budget in a Caribbean Easterly Wave. J Meteor Soc Japan 45: 444–466 Yanai M, Maruyama T, Nitta T, Hayashi Y (1968) Power Spectra of Large-Scale Disturbances over the Tropical Pacific. J Meteor Soc Japan 46: 308–323 Yasunari T (1979) Cloudiness Fluctuations Associated with Northern Hemisphere Summer Monsoon. J Meteor Soc Japan 57: 227–242 Yeh T-C, Gao Y-X (1979) Meteorology of the Tibetan Plateau. Scientific Publications Agency, Beijing (In Chinese) Zhou J, Lau KM (1997) Climatology of the South American Monsoon. Preprints of Fifth International Conference on Southern Hemisphere Meteorology & Oceanography. Pretoria South Africa, pp. 160–161
Author Index
A Abdullah, A. J., 78 Aberson, S. D., 73, 74 Akiyama, T., 140, 148 Algue, J., 164 Alisov, B. P., 27 Allan, R. J., 189 Ananthakrishnan, R., 98, 99 Andrews, J., 188 Angell, J. K., 115, 199 Anthes, R. A., 8, 62, 70, 79 Arkin, P. A., 236, 237, 264, 266, 267 Arnason, G., 39 Asnani, G. C., 26 B Balagot, V. F., 156, 161, 165 Bate, P. W., 189, 190, 191 Baumhefner, D., 40 Berson, F. A., 172 Bhalme, H. N., 115 Bilham, E. G., 55 Bjerknes, J., 4, 119 Black, M. L., 77 Blackadar, A. K., 282 Black, M., 78 Blake, D. W., 31 Blanford, H. F., 115, 117, 118 Bonner, W. D., 282 Boucher, K., 223 Boyle, J. S., 126, 130 Braham, R. R., 51, 52, 56 Brenner, I. S., 269 Bryson, R. A., 257, 268, 275, 279, 280, 286 Burpee, R. W., 84, 213 Butterworth, I. J., 189 Byers, H. R., 51, 52, 56
C Carlson, T. N., 195, 213, 220 Carpenter, T. H., 115, 119, 120 Chang, C. -P., 40, 47, 128, 130, 131, 140, 141, 142, 143 Charney, J. G., 3, 4, 30, 46, 121, 213 Chase, P. P., 84 Chen, G. T. J., 140, 141, 142, 166, 167 Chen, L., 126, 130, 136, 145 Cheng, K., 123 Chin, P. C., 164, 165 Cook, C., 153 D Davidson, N. E., 172, 182, 191 Deppermann, C. E., 66 Dickson, R. R., 118 Diercks, J. W., 78 Dines, W. H., 39 Ding, Y. -H., 105, 124, 137, 138, 140, 145, 166 Douglas, M. W., 256, 257, 269 Dunn, G. E., 39 E Eliassen, A., 46, 73 Estoque, M. A., 72 F Feng, Z. -Q., 105 Ferrel, W., 5 Findlater, J., 106, 136, 185 Fletcher, R. D., 16, 78 Flohn, H., 105, 138, 195, 207 Flores, J. F., 156, 161, 165 Fong, S. K., 137, 140 Frank, N. L., 213 Frank, W. M., 63, 65, 67, 68 Freeman, J. C., 44
313
314 Fuertes, F. A., 57 Fujita, T., 58–59 Fujita, T. T., 40 G Gadgil, S., 122 Galvin, J. F. P., 7 Gan, M. A., 241, 243, 247 Gao, Y. -X., 105, 123, 126, 128, 129, 138, 148 Garcia, E., 256, 257, 267, 268 Garden, G., 172 Gautier, C., 25 Gaviola, E., 57 Gentilli, J., 172, 188 Gentry, R. C., 78 George, C. A., 45 Gherzi, E., 164 Gill, A. E., 3, 14 Gray, W. M., 45, 49, 63, 64, 65, 66, 67, 72, 73, 248 Griffiths, J. F., 195 Gyakum, J. R., 131, 132, 133 H Hadley, G., 5, 6, 15–17, 18, 19, 22–24, 25, 30, 31, 105, 156, 169, 183–185, 233–235, 239, 295 Hahn, D. G., 115, 118 Hales, J. E, Jr., 269 Halley, E., 4, 24, 223 Hare, F. K., 257, 275, 279, 280, 286 Hastenrath, S., 239 Hawkins, H. F., 63, 64, 65, 69, 70 He, S. -X., 166 Helfand, H. M., 282 Heller, L., 239 Higgins, R. W., 269, 275, 282, 284 Hiser, H. W., 78 Hoecker, W. J., 282 Holland, G. J., 172, 189 Holton, J. R., 16, 40, 41 Hope, J. R., 84, 197 Hovermale, J. B., 84 Humphreys, W. J., 57 I Imbembo, S. M., 63, 64, 65, 69, 70 J Jelesnianski, C. P., 80, 81, 82 Johnson, D. H., 195 Julian, P. R., 122
Author Index K Kalnay, E., 85 Kanamitsu, M., 37, 115 Keenan, T. D., 189, 190, 191 Kendrew, W. G., 195, 223 Kiladis, G. N., 163 Klein, W. H., 27 Koteswaram, P., 45 Kousky, V. E., 241, 243, 247 Krishnamurti, T. N., 7, 10, 37, 40, 46, 73, 107, 115, 122, 166 Kuo, H. L., 72 Kurihara, Y., 78, 85 L Lajoie, F. A., 189 Landers, H., 39 Lau, K. M., 128, 130, 131, 145, 223 Lawrence, M. B., 85 Leftwich, P. W., 136 Lettau, H., 282 Li, W. -L., 166, 167 Lin, H., 137, 140 Livezey, R. E., 84 Loewe, F., 179, 181 Lord, S. J., 85 Love, G., 172, 191 Luo, H., 105 M Madden, R. A., 122 Malkus, J. S., 70 Mathur, M. B., 83 Matson, M., 118 Matsumoto, S., 148 Matsuno, T., 3, 14, 15 Maura, A. D., 239 McBride, J. L., 172, 189, 190, 191 McCune, S., 151, 152, 153 McNoldy, B. D., 75, 76 Means, L. L., 282 Miller, B. I., 67, 84 Mintz, Y., 121 Montgomery, M. T., 73, 74 Mooley, D. A., 115, 116, 117, 118, 119, 120 Morth, H. T., 195 Mosino Aleman, P. A., 256, 257, 267, 268 Murakami, M., 105, 166 Murakami, T., 105, 126, 128, 130, 145, 148, 172 N Namias, J., 239 Neumann, C. J., 83, 84
Author Index Newton, C. W., 54, 55 Nicholls, N., 172, 189 Ninomiya, K., 140, 145, 148, 149, 150 Nitta, T., 40 Nobre, C. A., 239 Nogues-Paegle, J., 238 Norquist, D. C., 195, 213 O Paegle, J., 238, 282 Palmen, E., 7 Paolino, D., 115, 120, 121 Parthasarathy, B., 115, 116, 117, 119, 120 Paterson, L. A., 189, 190, 191 Pearce, E. A., 223, 224 Pelissier, J. M., 83, 84 Petterssen, S., 20 Pittock, A. B., 172 Portig, W. H., 257 P Radok, U., 179, 181 Ramage, C. S., 20, 24, 26, 30, 169, 179, 223 Ramamurty, K., 122 Raman, C. R. V., 20, 179 Ramos, R. P. L., 241 Rao, Y. P., 24, 45, 100, 104, 107, 108, 112, 113, 114, 128, 136, 153, 185 Rasmussen, E. M., 119 Recker, E. E., 40, 41, 42 Reed, R. J., 40, 41, 42, 195 Reiter, E. R., 138, 275, 279, 282, 283, 285 Relf, E. F., 55 Rennick, M. A., 195, 313 Riehl, H., 25, 39, 40, 41, 68, 69, 70, 239, 272 Rosen, R. D., 122 Rossby, C. G., 5, 6, 14, 16 Q Saha, K., 30, 31, 37, 45, 47, 107, 173, 180, 184, 185, 191, 195, 213, 214, 217, 231, 233, 241, 242 Saha, K. R., 20, 21, 35, 36, 55, 56, 106, 110, 136, 185, 223 Saha, S., 20, 25, 30, 31, 35, 36, 37, 45, 107, 110, 173, 180, 184, 185, 191, 195, 213, 214, 217, 223, 231, 233, 241, 242 Salstein, D. A., 122 Sanders, F., 84, 131, 132, 133, 289 Sands, R. D., 268 Satyamurti, P., 223, 240 Schubert, S. D., 282
315 Schumann, T. E. W., 56, 57 Schwerdtfeger, W., 223 Selga, M., 160 Senn, H. V., 77, 78 Shea, D. J., 63, 73 Sheets, R. C., 63, 67 Shukla, J., 3, 4, 46, 115, 118, 119, 120, 121, 239 Sikka, D. R., 45, 119, 120, 122 Simpson, G. C., 106 Simpson, R. H., 62 Smith, G., 223, 224 Smith, R. K., 72 Smith, W. E., 121 Soliman, K. H., 31, 195, 209, 213, 214 Stern, M. E., 213 Sud, Y. C., 121 Sukanto, M., 156 Sumi, A., 172 Sun, Y., 136, 137 Surgi, N., 10, 115, 166 Sutcliffe, R. C., 39 R Taljaard, J. J., 223 Tang, M. C., 138, 275, 279, 282, 283, 285 Tannehill, I. R., 164 Tao, S. -Y., 136, 138, 140, 145 Tepper, M., 78 Thompson, B. W., 195 Thomson, J., 5 Trewartha, G. T., 131, 132 Troup, A. J., 172, 182 V Van den Dool, H. M., 223 Virji, H., 241 Vulquin, A., 223 W Walisser, D. E., 25 Walker, G. T., 4, 14, 15, 25, 115, 118, 119, 121, 156, 189 Wallace, J. M., 7, 40 Wang, B., 137, 140 Wang, J. -Z., 136 Watts, I. E. M., 126 Wegener, A., 55 Wexler, H., 78 Wiesnet, D. R., 118 Williams, K. T., 40 Willoughby, H. E., 72, 77, 78
316
Author Index
X Xie, P. P., 236, 237, 264, 266, 267
Yasunari, T., 122 Yeh, T. -C., 105, 123, 126, 128, 129, 138, 148
Y Yamazaki, K., 150 Yanai, M., 40, 41, 43, 105
Z Zehr, R., 191 Zhou, J., 223
Subject Index
A Aangular momentum, 5, 66, 122 Above-normal rainfall, 24 Absolute angular momentum, 5, 49, 291 Absolute vorticity, 43, 297 Adiabatic change, 38, 291 Adiabatic warming, 9, 39, 45, 50, 219, 247 Advection mechanism, 268 Afforestation, 120 Afghanistan, 92 Africa, 15, 24, 26, 28–30, 50, 90, 101, 103, 106, 182, 193, 195–198, 200–207, 209–221, 223–225 Agricultural calendar of climatic events, 123 Aircraft-observed profiles of wind, 78 Aircurrents fluctuations, 191 Air mass transformation experiment (AMTEX), 123, 130 Alaska, 260–261, 265–266, 276 Aleutian low, 140, 278 Amazon basin of South America, 15 Amazon river, 235 Andes mountains, 225, 256 Anticyclonic circulation, 15, 27, 92, 103, 105, 113, 126, 140, 162, 178–179, 192–193, 202, 204, 257–258, 262–263, 270, 277, 279, 291 Anticyclonic vorticity, 41, 289 Anticyclonic winds, 151 Appalachian mountains, 276, 280 Arabian Peninsula, 91, 103, 134 Arabian sea, 4, 23–24, 30–31, 35, 42, 45, 50, 91–94, 97–101, 103–105, 107–108, 113–114 Arakan, 89, 91, 102, 109–110 Arctic Circle, 276 Arctic ocean, 260, 277 Argentina, 225, 228–229, 235, 241
Asia, 4, 24, 26, 28–29, 35–36, 42, 89–91, 102, 106, 123–134, 136–137, 140, 141, 143, 146, 155, 157, 159, 161–162, 164, 166–167, 169, 182, 184, 197, 238, 279, 289 Astronomical high tide, 79 Atlantic ocean, 23, 28, 63, 195, 197, 200, 212–213, 215, 225–226, 229, 232, 235, 239–241, 245, 247–248, 250–251, 255, 257–258, 264, 266, 270, 272, 277–282, 288–289 Australia, 23, 26, 28, 30, 42, 101–103, 106, 129, 135, 155–157, 159, 161–162, 166–168, 171–174, 176–179, 182–189, 191–193, 223–225, 238 Axi-symmetric circulation system, 34 B Baiu front, 135, 141–142, 147–148, 152 Baja California, 75, 255 Balintang channels, 160 Bangladesh, 59, 80, 89, 91–92, 94–95, 98, 105, 110 Baroclinic atmosphere, 291 Baroclinic instability, 45, 107, 292 Baroclinic model, 85 Baroclinic wave activity, 3 Baroclinic wave disturbance, 128, 134, 140, 147, 152, 182, 219–221, 246–247, 288 movement of, 26 Baroclinic zone, 128, 209 Barograph, 68–69 Barotropic instability, 107, 292 Barotropic model (VICBAR), 85 Bashi channels, 160 Bay of Bengal, 4, 23, 35, 43, 45, 47, 80, 89, 91–97, 100–105, 107–109, 113–114, 164, 168, 179
317
318 Beta and Advection model (BAM), 84 Big wind, see Typhoons Blackbody, 167, 292 BOBBY – tropical disturbance, 191–192, 194 Boundary layer, 3, 11, 13, 94, 128, 130, 138, 283–284 Brazil, 225–226, 228–229, 232, 235, 238–241, 243, 245, 247–252 Break monsoon situation, 109 Buena Vista, 285 Bursters, 193 C Cameroons, 206, 219 Cape of Good Hope, 197 Caribbean sea, 225, 255, 258, 271–273 Carnot’s cycle, 11, 292 Carnot’s heat engine, 11 Charney-Stern theory, 213 Chinook wind, 280–281 CISK-barotropic-baroclinic instability, 45 Climates division on earth, 6 Cloncurry trough, 177 Clouding, 221 Cloud mass, 61, 77 Cloud-to-ground lightning, 54 Coastal mountain range, 89, 272, 276–277 Coastal topography, 79, 82 Cold surge, 128–130, 133, 191–292 Colombia, 225, 235, 256 Condensation heating, 8–9, 13, 16, 34, 45–47, 49, 72, 75 Conditional instability, 11, 292 Conduction, 292 Congo region, 103, 206 Conservation laws, 9–13 Carnot’s cycle, 11 conditional instability and convection, 11–12 coriolis control-variation, 12–13 direct and indirect circulations, 9–10 energy transformations, 10 shallow and deep convection, 12 Continental divide, 276 Continents and oceans, geographical distribution, 28 Convection, 8–9, 16–17, 19, 31, 38–39, 47, 50–51, 66–68, 72, 75–78, 94–96, 99, 120, 122–123, 129, 162, 185, 221, 240, 247, 268, 277 Convergence zone, 3, 16, 18–19, 72, 78, 97, 129, 139, 155–156, 162–163, 209, 277
Subject Index Coral sea, 23, 28, 162, 171, 175, 177–178 Coriolis force, 12, 293 rr Coriolis, 49, 66, 293 Crachin, 133 Crested butte, 285–286 Cross-equatorial current, 47, 102, 182 flow, 21, 28, 45, 47, 101, 177–179, 182, 203, 263 Cumulonimbus cloud, 16, 40, 51–54, 56–57 Cumulus clouds, 16, 40, 51, 193 Cyclogenesis, 12, 45, 47, 123, 147, 160, 164, 219, 289 Cyclone-anticyclone alternation, 24 Cyclone track and intensity prediction moels, 82–85 Cyclonic circulation, 15, 92, 101, 129, 140, 142, 177, 193, 232, 238, 262–263, 272, 279–280, 289 Cyclonic momentum, 66 Cyclonic motion, 293 Cyclonic storm, 33, 97, 107–108, 113, 289 tracks of, 100, 114 Cyclonic vortex, 107, 243, 251 Cyclonic vorticity, 93, 289 D Darwin pressure anomalies, 120 Data assimilation techniques, 85 Del operator, 44 Denakil desert, 197 Deviation wind’s meridional component, 233 Diabatic change, 8–9, 293 Differential heating, 5, 7, 14, 26–27, 34, 39, 102, 151, 160, 238, 279, 288 Direct circulation, 5, 10, 38, 72, 293 Double equatorial troughs formation, 96–97 Dry adiabatic lapse rate, 293 Dynamic instability, 213, 293 E East Africa, 4, 29, 90, 96, 197–198 Easterly wave, 34, 39–43, 44–47, 93, 121, 164, 272–273 Indian ocean region, 41–43 Tropical North Atlantic, 39–40 Tropical North Pacific, 40–41 Eastern Indian ocean, 15, 97, 119, 134, 168, 177, 182, 187, 189 See also Indian ocean Eastern-Mediterranean region, 92 Eastern Pacific ocean, 28, 50, 119, 168, 189, 235, 255, 262, 268–270 East of Continental Divide, 276
Subject Index ECMWF, 293 Electrical storm, 51 El Nino southern oscillation (ENSO), 171–172, 293 El Nino, 117, 119–120, 156, 168 E’ly waves, see Easterly wave Entropy, 293 Equator, 3, 5–6, 12, 14–17, 23, 26, 28, 42, 50, 92–93, 95–99, 101–103, 106, 129, 135–136, 155, 157, 161–162, 174, 177, 180, 182–185, 196–197, 200, 204, 206–207, 213–214, 226, 228–229, 234–235, 238–239, 242–243, 245, 250, 258, 266, 268–269 Equatorial circulation, 4, 14–16, 208, 293 Equatorial heat source, 17–19, 21, 25, 27, 101–102, 156, 293 Equatorial heat’s seasonal migration, 17–22 forcing for seasonal movement, 19–21 intraseasonal oscillation of monsoon, 21–22 monsoon’s origin, 17–18 wave structure’s origin, 18–19 Equatorial tropopause, 16 Equatorial trough, 3–4, 20, 22, 29, 33, 39, 93, 95–98, 101–103, 113, 129, 134, 137, 156, 160–165, 177, 202–204, 209, 214–220, 225–226, 239, 256 Equatorial westerlies formation, 96–97 Equatorward drift, 5 Ethiopian mountains, 211 Eurasian snow cover, 117–119 Explosive cyclone, 130–132 definition, 132 distribution of, 132–133 Extratropics, 25, 225, 294 Eye and eye-wall, 71–77 concentric multiple eye-walls, 75–77 hurricane eye formation, 71–72 meso-scale vortices evidence, 73–75 spiral bands around, 77–78 origin and direction of propagation, 78 structure, 77–78 Eyewall replacement cycle, 77 F Feeble wave, 72 FGGE, 172, 294 Finite-difference technique, 80 Fletcher model, 16 Fog, 92, 133, 155 Forecast model, 83, 85
319 Forest fires, ignition of, 54 Frozen Arctic ocean, 259, 277 Fujita-scale (F-scale), 59 Full-grown tornado, characteristics, 58 Funnel cloud, 58 G Galapagos island, 225 Gale force, 193, 289, 294 Ganga-Brahmaputra delta, 79–80 General circulation, 3, 5, 294 Geopotential meter, 294 Geostationary operational environmental satellite (GOES), 248 Geostrophic wind, 128, 283, 294 GFDL model, 83, 85 Gibson desert, 173 Global Atomspheric Reserach Program (GARP), 123, 130, 294 Global data analysis system (GDAS), 243 GMS-1 geostationary satellite, 166 Gobi desert, 124–125 GPS, 73 Gradient-level monsoon trough, 190 Gray’s survey, 50 Great Australian Bight, 171, 177, 179 Great dividing range (The), 29, 171, 177 Greenwich meridian, 30–31, 212 Guinea mountains, 207 Gulf of California, 255, 268–269 Gulf of Mexico, 61–62, 79–80, 255, 258, 268–271, 273, 285, 289 Guyanas, 238 H Hadley circulation, 5–6, 15–18, 30–31, 105, 156, 169, 183, 185, 233–234, 239, 294 See also Walker circulation Hadley’s single-cell model, 5, 16 Hailstorm, 34, 50–51, 55–56, 94 Heat energy, conservation of, 13 Heat low circulation, 94, 103, 110, 139–140, 142, 162, 202, 209, 220, 232, 240, 242, 269 Heat sink, 3, 5, 7, 9, 11, 14–15, 17–18, 21, 27, 89, 101, 138, 168, 238, 240, 279, 294 Heat sources and sinks, 7–9 definition of, 7 diabatic/adiabatic sources, 8–9 High pressure area, 101, 128, 162, 177, 251, 279, 294 High-speed jetstream, 128
320 Himalayan mountain, 4, 89, 91, 105, 110, 124, 127–128, 130 Himalayan Massif, 126 Hovmoller diagram for high-pass-filtered, 131 Huang-Ho river valley, 144 HURRAN analog model, 85 Hurricane, 4, 44, 61–70, 72–75, 78, 83, 85, 164, 214–217, 289, 294 Andrew, 61–62 Anita, 63, 67 Gordon, 215–217, 220 INEZ, 63–65, 69, 72 Isabel, 73–74 Juliette, 75 Katrina, 79 thermodynamics of, 72 Hybrid models, 83 Hydrodynamical equations, 41 Hydrostatic approximation, 294 I Icelandic low, 278 IIOE, 169, 179, 295 Inclined troughs, 22–24 occurrence of, 23 India Meteorological Department, 92, 98, 104, 111 Indian ocean, 23, 28, 41–42, 83, 89–97, 99–103, 111, 123–136, 159, 162, 168, 171, 177, 182, 184, 187, 189, 197, 200, 204, 207, 213, 221, 238 cyclones in, 33, 97 Indonesia, 26, 103, 155–159, 177 Inertia-gravity waves, 80 oscillations, 14 Infrared, 162, 166, 294 Instability, 3, 12, 94, 130, 160, 213, 295 Internal baroclinic jet theory, 213 See also Charney-Stern theory Intertropical convergence zone (ITCZ), 17–18, 25–29, 31, 45, 47, 72, 78, 91, 96, 101–106, 135, 137, 139–140, 155, 159, 160–165, 178, 206, 219, 232, 235, 237–240, 257, 262–263, 265–266, 269, 273, 277, 295 Intertropical front (ITF), 209 Intracloud discharges, 54 Intraseasonal oscillation, 4, 21–22, 25, 121, 172, 295 Isallobaric gradient, 21, 160, 192 Isallobaric lows movement, 39 Isallobars, 44
Subject Index Isentropic surface, 295 Isobar, 295 Isohyets, 207, 267–268 Isolines of vorticity, 41 Isotherm, 198, 295 J JASON – tropical disturbance, 191 Jelesnianski’s storm surge models, 80 Jetstream, 93, 95, 129, 149, 169, 179, 181–182, 209, 287, 295 K Kalahari-Namib desert, 195 Kal-Baisakhi, 59, 95 Kenya-Ethiopian mountains, 207 Kenya, 90, 207, 219 Kinematic relation, 20 Kinetic energy, 6, 9–10, 38, 51, 75, 78, 214, 295 Korean Strait, 124 L Lagrangian time, 33, 295 Land-sea thermal contrast, 37, 103, 146, 236 La Nina events, 119 Large hailstones formation theory, 56–57 Latent heat, 8, 11, 16, 33, 50, 252, 295 Latent heat of condensation, 48 Line Islands, 120 Longitudinal compression, 54 Longwave radiation, 186 Lower Equatorial (LE), 16 Lower-tropospheric cyclonic circulations, 204 Low level jet (LLJ), formation of, 238–239 M Madagascar area, 200, 205 Madden-Julian Oscillations (MJOs), 122 Mammatos, 60 Maritime continent, 26, 101, 123, 129, 136, 155–170, 182 climate of, 156–159 pressure, 156–157 rainfall, 158–159 relative humidity and cloudiness, 158 temperature, 157 element of, 158 equatorial trough, 160–165 geographical location and topography, 159–160 ocean currents, 160 Marra mountain, 197, 211 Mean sea level pressure (MSLP), 244
Subject Index Mediterranean sea, 30, 105, 197–198, 214 Meiyu–baiu front, 135, 138, 142, 152 Meridional circulation, 5, 7, 15–16, 31, 169, 212, 223, 233–234 features of, 234 Meridional component of wind, 35, 179–180, 183, 207 Meridional-vertical circulation, 16, 106 Meridional wind, 41, 295 Meso-scale disturbances and local storms, 50–60 storm’s energy source, 50–51 hailstorms, 55–57 thunderstorms, 51–55 tornadoes, 57–60 Meso-scale trough, 77 Meso-scale vortices, 73, 140 Mexican monsoon, 256 Mexican plateau, 258–259, 267, 270, 273, 278 Mexico, 75, 255, 257–259, 261, 263, 265, 267–268, 270–273, 283 Midlatitudes, 5, 7, 34, 39, 295 Midtroposphere, 179, 182, 295 Midtropospheric cyclones, 108 jet, 213 Mississippi delta, 80 Mississippi valley, 276, 285 Missouri rivers 277 Moderate resolution imaging spectroradiometer (MODIS), 74 Moist convection, 68, 71, 270 Moisture-bearing aircurrent, 269 tradewinds, 106 Moisture budget, 284–285 Moisture convergence zone, 112, 275, 283 Mojave-Sonoan desert, 29 Monsoon and hadley circulations, 19, 31 Monsoon circulation, 17–18, 25, 27, 89, 95, 101, 104, 123, 155, 183, 185, 197–198, 223, 226, 230–234, 240, 275, 277, 279, 288 cross-equatorial movement of, 95–96 properties of, 17, 233, 275 Monsoon convection zone, 235 Monsoon depression, 45–47, 107, 122, 189–190 development of, 45 formation of, 45–47 Monsoon over Africa, 195–222 climates of Africa, 198–207
321 equatorial trough, 209–212 equatorial westerlies, 207–208 midlatitude waves of southern hemisphere, 220 physical features and environment, 196–198 structure, development and movement of waves, 219–220 wave disturbance origin, 212–219 Monsoon over Australia, 171–194 air temperatures, 174–175 annual rainfall of, 187–188 atmospheric pressure, 176–177 depressions and cyclones, 189–191 hadley circulations, 183–185 location and physical features, 171–172 ocean surface temperature (SST), 173–174 onset of monsoon, 182–183 summer monsoon rainfall, 185–187 tropical-midlatitude interaction, 191–194 variability of australian rainfall, 189 wind and circulation, 177–182 Monsoon over Central America, 255–273 climate of, 257–264 heat sources and sinks, 256–257 rainfall, 264–270 weather’s characteristic features, 270–273 Monsoon over Eastern Asia, 123–154 airmass transformations, 130–133 extratropical latitudes, 143–145 Japan, 145–149 Korea, 149–154 changma season, 152 climatic zones, 152–154 historical background, 149 physical features and climate, 149–151 winter monsoon, 151–152 meteorological developments associated with China, 137–142 Tibetan plateau monsoon, 138–139 Meiyu (plum rain) front, 139–142 origin of, 134–136 physical features and climate, 124–125 season, 125–130 stationary states and jumps, 137 summer monsoon, 136–137 transition period, 133–134 Monsoon over North America, 275–290 appalachian mountain range, 288–289 characteristic features of east coast monsoon, 287–288 climatological background, 275–278 heat sources and sinks, 278–279
322 seasonal circulations, 279–286 storms and hurricanes, 289–290 w’ly wave disturbances, 287 Monsoon over South America, 223–252 climatological features, 227–236 physical features and environment, 224–227 quasi-stationary waves, 236–241 tropical cyclone, 248–252 tropical disturbances, 241–247 Monsoon over Southern Asia, 89–122 rainfall over Indian subcontinent, 110–111 summer monsoon, 111–114 dates of withdrawal of monsoon, 111 disturbances of withdrawal phase, 113–114 retreating monsoon rain over Tamil Nadu, 111–113 summer monsoon disturbances, 107–109 summer monsoon to Indian subcontinent, 98–106 advance over Indian ocean, 102–104 advance to western himalayas, 105–106 monsoon rainfall, 106 onset over the, 104–105 transition season, 93–98 ‘heat lows’ and ‘cold highs’, 94 Indian ocean (equatorial), 95–98 local storms, 94–95 western disturbances, 94 variability of Indian summer monsoon, 115–122 interannual variability, 115–121 intraseasonal variability, 121 winter season disturbances, 92–93 Monsoon’s co-existence with desert circulation, 29–31 Monsoon’s global and regional distribution, 25–29 extratropical monsoons, 26–27 tropical monsoons, 25–26 zonal and meridional anomalies, 27–29 Monsoon stationary wave, 211, 237 Monsoon trough zone, 105, 107, 110, 191, 234–235 Monsoon trough, 35, 42, 44–45, 47, 72, 97, 105, 107–111, 113, 122, 135, 182, 185, 190–191, 220, 234–235, 238 Monsoon withdrawal phase, characteristic features, 113 Mountain barriers, effects of, 29 Mount Everest, 124 Mozambique, 90, 197–198, 200, 205
Subject Index N Namib-Kalahari desert, 30 National Hurricane Center (NHC), 63, 83 National Hurricane Research Laboratory, 63 National Meteorological Center, 84 NCEP reanalysis, 31, 96, 100, 144, 173–174, 176, 257, 275 Net warming, 51 Nigeria, 55, 198, 206 Nile river, 197 Nile valley, 211, 215 NOAA satellite, 61–62 Nomograms, 82 North Atlantic ocean, 39, 83 North equatorial trough (NET), 92, 102, 134 Northern summer monsoon, 162 Northern winter monsoon, 162 Northers, 270–271 North Pacific ocean, 39 Nucleus, 250, 295 O Ocean currents, 28 Oceanic cyclones, 130 Oceanic monsoons, 28 Onset vortex, 107 Orographic effects, 289 Orography, 89, 103, 110, 171, 198, 209, 236 Outgoing longwave radiation (OLR), 185–186, 207 P Pacific composite typhoon, 65, 67 Pacific ocean, 23, 26, 28, 33, 42, 63, 105, 119, 124, 126–128, 130, 136, 138, 140, 142, 144, 146–147, 155, 159, 162, 164, 168, 171, 177, 189, 225, 235, 242, 255, 257–258, 262–263, 265–266, 268–270, 272, 277–278, 280–281 Pacific ocean anticyclonic circulation, 140 Pacific typhoon, 67 Pakistan, 30, 89–92, 103, 105, 110 Panama canal, 255 Paraguay, 225, 228 Patagonia desert, 30 Philippines, 26, 155–159, 164 Plains low level jet (LLJ), 282 Plateau monsoon, 138, 144, 275, 278–279, 282, 287 Polar regions, 266–267 Potential energy, 10–11, 38, 46, 51, 75 Potential temperature, 51, 69, 71, 73–74, 137, 296
Subject Index Potential vorticity, 289, 297 conservation of, 13 conservation principle of, 13, 29 Pressure tendency equation, 44 gradient, 20 laplacian, 43 Processes, 8–9 Q Qinghai-Xizang Plateau Meteorology Experiment (QXPMEX), 123 Quasi-axi-symmetric disturbance, 72 Quasi-Lagrangian model (QLM), 83 Quasi-stationary inclined equatorial trough, 24 Quasi-stationary monsoon, 42, 44–48, 107, 121, 210, 221, 243 Quasi-stationary monsoon low pressure, 221 Quasi-stationary trough, 34, 45, 217, 221, 240, 242 Quasi-stationary wave, 33–39, 41, 44, 171, 192, 214, 229, 236, 246–247, 250 in temperature fields, 36–37 in wind field, 35 structure of, 37–39 R Radar, 62, 73 Radial and vertical distributions, 63 Radial winds, vertical cross-section of, 66 Radiosonde, 62 Rain gauge stations, 115 Rankine vortex, 66 Rarefaction, 54 Red sea, 215 Relative vorticity, 13, 58, 75–76, 297 Rocky Mountain, 255, 258, 261, 265, 267, 275–276, 280–281, 283, 286 Rossby model, 5–6 S Saffir-Simpson scale, 61, 73 Saharan desert, 29–31, 195, 197 SANBAR, 83–84 Satellite cloud imagery, 40, 61, 93, 97, 163, 214, 221, 238, 240, 248 Saudi Arabia, 31, 90, 94 Saudi Arabian desert (Rub-al-Khali), 31 Seasonally-migrating equatorial heat, 25 Seasonally-reversing tradewind circulation, 24 Seasonal movement of heat, 26 Seasonal reversals, 26, 138, 275, 282, 287 Sea surface temperature (SST), 85, 117, 119, 131, 160, 198, 226–228, 239–240
323 Shortwave radiation, 296 Sierra Madre oriental, 255, 267, 270–271 Simpson desert, 173 Single-cell circulation model, 5 S’ly low level jet (LLJ), 280 Soil moisture, 117, 120 Somalia, 90, 94, 96, 102–103, 197 Somali coast, 103, 197 Somali jet, 103, 238 Somali monsoon current, 136 Sonora-Arizona desert, 266, 268–269 South China sea monsoon experiment (SCSMEX), 124 South equatorial trough (SET), 42, 92, 102, 134, 168 Southern oscillation, 117, 119, 168, 189 Southern oscillation index (SOI), 189 Southwest Atlantic convergence zone (SACZ), 3, 27, 235, 240 Southwestern Pacific ocean, 23, 167 Southwest Pacific convergence zone (SPCZ), 3, 28, 159, 162, 168, 178 Squalls, 34, 53, 55 Sri Lanka, 93, 99, 101, 103, 111 Stationary wave, subtropical, 237 Statistical-dynamical models, 85 Statistical hurricane intensity prediction scheme (SHIPS), 85 Steering principle, 83 Storm-relative tangential and radial winds, 73 Storm surge, 33, 78–82, 114, 164 aspects of, 79–80 mathematical models, 80–82 Streak or forked lightnings, 54 Subnormal rainfall, 24 Subsidence warming zone, 72 Subsynoptic-disturbances, 50 Subtropical anticyclones, 202–203, 207, 226, 263 Subtropical front (STF), 30–31, 209, 213–214 Subtropical high pressure, ridge of, 7, 16, 177, 273 Subtropical jet (STJ), 7, 169, 179, 181–182, 296 Sumatra coast, earthquake near, 83 Surface temperatures (SST), 257 Synoptic-scale convergence, 11 disturbances, 33, 50, 59, 171 stationary wave circulations, 39 tropical disturbances, classification of, 34 wave disturbances, 33, 50
324 T Tangential and radial components, 66 Tangential winds, 65, 72 vertical distributions of, 67 Tanzania, 90 Tasman sea, 177 Tennaserim ranges of Myanmar, 89 Thar desert, 30, 110 Thermally-driven circulation, 5 Three-dimensional global circulation models, 85 Thundercloud and vertical currents, cellular structure, 51 Thunderstorm, 34, 40, 50–51, 52, 55–56, 94, 155, 160 structure of, 54 Tibetan plateau, 4, 13, 47, 89, 100, 105, 111, 123–124, 126–128, 138–139, 144, 279 Tornadoes, 33–34, 50–51, 58–60, 95 Tradewind circulation, 6, 17–18, 25, 137, 163, 183, 223, 230 Tradewinds, 4–6, 16–18, 27, 33, 39–40, 93, 96, 103, 106, 127, 135–136, 139, 142, 155, 161–163, 177, 179, 198, 202–204, 207, 225–226, 230, 232, 238–239, 257, 262–263, 268–273, 281 Transition season, 93–97, 281–282, 286 Tropical belt, 5–6, 16, 27, 124, 142, 187, 225, 255 Tropical circulation, 1, 3, 5–7, 13, 33, 44, 61, 89, 123, 155–156, 163, 171, 195, 223, 296 poleward boundary of, 6–7 tradewinds, 4–6 Tropical convection, 122, 182 Tropical convergence zone (TCZ), 16–18, 24, 29, 45, 47, 49, 72, 101–106, 135, 137, 139–140, 161–162, 178, 213, 219, 232, 237–240, 277 Tropical cyclone, 4, 24, 48–50, 60–63, 75, 77, 82–85, 156, 164–165, 172, 189–194, 247–249, 251–252 Tropical cyclone’s observed structure, 61–71 pressure distribution, 68–69 radial and tangential components, 64–67 temperature distribution, 69–71 vertical motion, 67–68 wind structure, 63–64 Tropical disturbance, 8, 11, 19, 33, 40, 44, 61, 191–192, 241 Tropical killer cyclones, statistics of, 114 Tropical monsoon, definition of, 25 Tropopause, 51, 55, 64, 85, 167, 297
Subject Index Troposphere, 6, 16–17, 31, 37, 39, 46–47, 51, 65, 67–69, 72, 94, 99, 105, 129, 174–175, 179, 185, 199–200, 243, 247, 264, 269–270, 297 Tropospheric convergence, 40 Trough zone, 47, 49, 105, 129, 240 Tsunami (high tidal wave), 83 Typhoons, 4, 33, 39, 61, 63, 147, 156, 159–160, 164–165 U Underwater volcanoes, 83 Upper Equatorial (UE), 16 Upper-tropospheric cold cyclonic vortex, 246 cyclonic vortex, 241–244, 247, 250–251 warm ridge, 243 Uruguay, 225 V Vector Del operator, 20 Vegetation and albedo, 117, 120 Venezuela, 228, 235, 238–239 Vertical p-velocity, 150, 211, 233 Vic Ooyama Barotropic model (VICBAR), 84 Vorticity, 13, 75, 93, 141, 289 W Walker circulation, 15, 25, 119, 156, 297 Waves development, 43–50 meaning of, 43–44 quasi-stationary monsoon trough, 45–50 Weather-forming processes, 3 West Africa, 59 Western atlantic ocean, 50, 255, 278, 289 Western deserts, 30 Western disturbance, 92, 94 Western Ghats mountains, 103, 110 Western Pacific ocean, 15, 50, 123, 127, 136, 155, 162, 289 Westward-propagating tropical disturbances, 172 Winter monsoon experiment (WMONEX), 172 W’ly waves, 46–47, 97, 107–108, 113, 130, 171, 192, 250–251, 287, 289 subtropical, 46 Y Yangtze river, 137, 140, 146 Yellow sea, 124, 146, 150 Yield hailstones, 55–56 Z Zambia, 198, 205 Zimbabwe, 90, 198 Zonal-vertical circulation, 24, 37, 243, 247
