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- Herausgeber: John Wiley & Sons
- Kategorie: Wissenschaft und neue Technologien
- Sprache: Englisch
Climatology in Cold Regions A groundbreaking interdisciplinary study of cold-region weather systems and their vital role in predicting climate change across the globe Climatology in Cold Regions explores the complexities of land-atmospheric interaction across the Earth's cryosphere, systematically placing soil thawing, snow melting, surface diabatic heating, and other processes within the context of broader climatological models. Drawing from a wealth of new data, leading atmospheric scientist Chenghai Wang illustrates how cold-region weather systems can be parameterized to improve seasonal climate prediction and provide crucial insights into projected changes in climate over the next 50-100 years. The book opens with an introduction to the characteristics and classification of cold-region climatology, followed by a detailed description of the primary weather systems and land surface processes in cold regions. The core of the book presents a new approach for seasonal climate prediction using signals obtained from cryospheric processes, supported by a discussion of climate disasters and the impact of climate change on the ecology of cold regions. * Introduces a new way of modeling climate in cold regions * Offers novel approaches for assessing climate signals from cold regions in seasonal and sub-seasonal predictions * Presents new data on the role of cold-region climatology in forecasting and driving global temperature changes * Discusses the role of cold regions as the main source of global freshwater supply A significant contribution to climate research and beyond, Climatology in Cold Regions is essential reading for students, scientists, and researchers in the atmospheric sciences, meteorology, ecology, hydrology, and Earth sciences.
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Table of Contents
Cover
Title Page
Copyright Page
List of Figures
List of Tables
Foreword
1 Introduction
1.1 Scope of the Subject and Recent Highlights
1.2 Some Definitions and Terms of Reference
1.3 Characteristics of Cold Region Climatology
1.4 Classification of Cold Region Climatology
1.5 Impacts of Cold Regions on Global Climate
1.6 Contribution of Cold Regions to Energy Balance
1.7 Summary
2 Climate in Cold Regions
2.1 General Circulation
2.2 Low and High Frequency of Atmosphere Variability
2.3 Weather Systems
2.4 Radiation
2.5 Precipitation and Snow
2.6 Frozen Ground
2.7 Hydrology
2.8 Winter Monsoon
3 Land‐Surface Process in Cold Regions
3.1 Land‐Surface Characteristics in Cold Regions
3.2 Daily and Seasonal Cycle of Freeze–Thaw in Soil
3.3 Soil Freezing–Thawing and Snow‐Melting Processes and their Hydrothermal Effects
3.4 Impacts of Soil Freeze–Thaw Process and Snow Melting on Surface Diabatic Heating
4 Land–Atmospheric Interaction in Cold Region
4.1 Spring Predictability Barrier
4.2 Signal of Climate Prediction in Land Processes in Cold Regions
4.3 Improvement of Climate Prediction by Assimilating Land‐Surface Signals in Cold Regions
4.4 Soil‐Moisture Memory in Cold Regions
4.5 Soil‐Moisture Climate Coupling in Cold Regions
4.6 Snow–Climate Coupling in Cold Regions
5 Models and Modeling in Cold Regions
5.1 Snow and Freeze–Thaw Parameterizations
5.2 Hydrothermal Coupling Scheme in the Land‐Surface Model
5.3 Simulation of Land‐Surface Process in Cold Regions
6 Hydrological Processes in Cold Regions
6.1 Glacier Hydrology
6.2 Snowmelt Runoff
6.3 Frozen Soil and Hydrology
6.4 River Ice and Lake‐Ice Hydrology
6.5 Sea‐Ice Hydrology
6.6 Sea Ice
6.7 Characteristics of Sea Ice Regimes in China
6.8 Precipitation Rates on the Tibetan Plateau
6.9 Precipitation Recycling on the Tibetan Plateau
7 Ecology and Vegetation in Cold Regions
7.1 Overview
7.2 Ecology and Vegetation Characteristics on the Tibetan Plateau
7.3 Ecology and Vegetation Characteristics in Middle and High Latitudes
7.4 Ecology and Vegetation Characteristics in the Arctic Region
8 The Carbon and Nitrogen Cycle in Cold Regions
8.1 Carbon, Nitrogen, and Greenhouse Gases on the Tibetan Plateau
8.2 Carbon, Nitrogen, and Greenhouse Gases at High Latitudes
8.3 Carbon, Nitrogen, and Greenhouse Gases in the Polar Regions
8.4 Outlook
9 Climate Disasters in Cold Regions
9.1 Overview
9.2 Snowmelt Flood
9.3 Freezing Rain
9.4 Winter Storms
9.5 Ice Avalanches
9.6 Ice‐Jam Floods
9.7 Snow Avalanches
9.8 Freezing and Thawing Damage
9.9 What's Next
10 Climate Changes in Cold Regions
10.1 Climate Change in Cold Regions in the Past Decades
10.2 Projection of Future Climate in Cold Regions
10.3 Changes of Related Environment Response to Climate Change in Cold Regions
11 Challenges and Outlook
11.1 Challenges for Future Study
11.2 Conclusion
Appendix
A. Soil Hydraulic and Thermal Properties
B. Numerical Implementation
Bibliography
Index
End User License Agreement
List of Tables
Chapter 3
Table 3.1 Major Arctic ice caps (Field 1975; Govorukha 1988; Williams and Fe...
Chapter 5
Table 5.1 The root mean square errors of
soil temperature
(
ST
) and
soil mois
...
Table 5.2 The magnitude of variables in the heat conduction equation (Eq. 5....
Table 5.3 The magnitude of variables in the modified Richards equations (Eq....
Chapter 7
Table 7.1 Ecosystem cover area (Unit: 10
2
km
2
).
Chapter 8
Table 8.1 Carbon and nitrogen and its transport to atmosphere in cold region...
List of Illustrations
Chapter 1
Figure 1.1 Definition of cold regions. The areas colored purple are the cold...
Figure 1.2 Distribution of precipitation (mm) for winter (December
–
Jan...
Figure 1.3 Distribution of near‐surface air temperature (°C) for January (to...
Figure 1.4 Comparisons of precipitation (mm, bar) and near‐surface air tempe...
Chapter 2
Figure 2.1 The geopotential height (contour) and horizontal wind (vector) at...
Figure 2.2 Latitude–pressure cross‐sections (80–100E mean) of zonal temperat...
Figure 2.3 The variance of the unfiltered (left), 2.5–6‐day filtered (middle...
Figure 2.4 The mean geopotential height (gpm) at 30 hPa for the four mid‐sea...
Figure 2.5 The geopotential height (gpm, color) and horizontal wind (m s
−1
...
Figure 2.6 Comparison of geopotential height fields at the lower jet stream ...
Figure 2.7 MODIS satellite image 06 April 2007 showing the polar low which f...
Figure 2.8 Mechanism for decadal variability of stratospheric sudden warming...
Figure 2.9 The vertical profile structure of tropical cyclone‐like Tibetan P...
Figure 2.10 The (a) East–West transverse shear line and (b) north–south perp...
Figure 2.11 The various radiative mechanisms associated with aerosol–cloud i...
Figure 2.12 Distribution of summer (June, July, August) precipitation (mm da...
Figure 2.13 Distribution of winter (DJF) snow depth (unit: m) over northern ...
Figure 2.14 Distribution of permafrost and extent of permafrost and seasonal...
Figure 2.15 Distribution of climatological (1961–1990) seasonally frozen gro...
Figure 2.16 Distribution of snow melting in spring (MAM snow depth minus DJF...
Figure 2.17 Spatial distribution of sea‐level pressure color; unit: hPa and ...
Chapter 3
Figure 3.1 Mean sea‐ice extent for March (all shaded areas) and September (d...
Figure 3.2 Physiography of the Arctic lands, showing topography and major ri...
Figure 3.3 Distribution of Arctic polar desert (dark shading) and approximat...
Figure 3.4 Average number of weeks of snow cover in the northern hemisphere,...
Figure 3.5 Annual evapotranspiration (ET) for the terrestrial region. Time s...
Figure 3.6 Diagram of the soil freeze–thaw process.
Figure 3.7 Seasonal cycle of soil freeze–thaw process. Section of soil moist...
Figure 3.8 Diurnal cycle of the soil freeze–thaw process. Diurnal variations...
Figure 3.9 Variations of soil hydrothermal properties in the freeze–thaw pro...
Figure 3.10 Effects of the soil freeze–thaw (FT) process on soil‐moisture va...
Figure 3.11 Variations of monthly surface‐sensible and surface‐latent heat f...
Figure 3.12 Comparison of surface sensible heat (SH) flux (top panel) and su...
Figure 3.13 A schematic diagram summarizing the impacts of the soil freeze–t...
Chapter 4
Figure 4.1 The forecast skill (or forecast ability) of model runs based on F...
Figure 4.2 (a) Distribution of stand deviation (color) and climatological me...
Figure 4.3 Evaluations of standardized series of snow depth average over the...
Figure 4.4 Relation between soil freeze–thaw process and the Asian summer mo...
Figure 4.5 Series of soil moisture (principal component of the empirical ort...
Figure 4.6 (a) Evolution of the
in situ
observation of soil moisture (SM), p...
Figure 4.7 Distribution of summer (June, July, August) precipitation root me...
Figure 4.8 Soil moisture memory (unit: days) in spring (March–April–May aver...
Figure 4.9 Spatial distribution of soil moisture memory (days) during March,...
Figure 4.10 Soil moisture–precipitation coupling in the northern hemisphere....
Figure 4.11 Distribution of coupling index
k
between spring (MAM) soil moist...
Figure 4.12 Evaluation characteristics of the coupling relation between soil...
Chapter 5
Figure 5.1 Comparison of daily soil temperature (upper row; °C) and soil moi...
Figure 5.2 Comparison of daily soil temperature (upper row; °C) and soil moi...
Figure 5.3 Coupling relations between soil temperature and soil moisture dur...
Figure 5.4 Diurnal cycles of soil temperature (upper row; °C) and soil moist...
Figure 5.5 Schematic diagram of water–heat transport processes during soil f...
Figure 5.6 Schematic illustration of soil‐heat transport (left) and soil wat...
Figure 5.7 Impacts of soil water content and soil textures on hydraulic cond...
Figure 5.8 Vertical sections of daily soil temperature (°C) (first and third...
Figure 5.9 Vertical sections of daily soil moisture (expressed as liquid wat...
Chapter 6
Figure 6.1 Runoff components of continental glacial watershed (Ding 2017).
Figure 6.2 Water cycle in glacial areas.
Figure 6.3 The confluence of snow meltwater (rainfall) (adapted from Ding, 2...
Figure 6.4 Responses of snowmelt runoff to air temperature increases, +2°C, ...
Figure 6.5 Location map of the Spiti Basin.
Figure 6.6 Process of groundwater supply in a cold area.
Figure 6.7 Influencing factors of the hydrological freezing–thawing process ...
Figure 6.8 Research domains. A, B, C, and D represent four regions in the Ti...
Figure 6.9 Distribution of the mean precipitation rate (1970–2009; %) over t...
Figure 6.10 Evolution of precipitation rate in four regions (a) ERA40 reanal...
Figure 6.11 The evolution of seasonal precipitation rate (1970–2009), (a)–(d...
Figure 6.12 Monthly variability of precipitation rate with mean from 1970 to...
Chapter 7
Figure 7.1 Spatial patterns of vegetation phenology. Only the first cycle is...
Figure 7.2 Circumpolar maximum normalized difference vegetation index (NDVI)...
Figure 7.3 Plant physiognomy occurring in different tundra bioclimate subzon...
Chapter 8
Figure 8.1 The model and mainly influencing factors for carbon transfer in t...
Figure 8.2 Schematic diagram of annual carbon partitioning in diverse carbon...
Figure 8.3 A conceptual model of the relationship among the environmental co...
Figure 8.4 Key disturbance processes impacting northern high‐latitude soil o...
Figure 8.5 The distribution of soil organic carbon (kg m
−2
) over the t...
Figure 8.6 The current state of the Arctic carbon cycle based on a synthesis...
Figure 8.7 Key factors controlling gaseous nitrogen (N) input and output pat...
Chapter 9
Figure 9.1 Snowmelt flooding, from National Oceanic and Atmospheric Administ...
Figure 9.2 The formation of the snowmelt flood.
Figure 9.3 Freezing rain.
Figure 9.4 Freezing rain falls when snowflakes melt completely before reachi...
Figure 9.5 Distribution of freezing rain in China from 2000 to 2015 (unit: s...
Figure 9.6 Maximum number of hours with freezing rain in a single year, 1928...
Figure 9.7 Winter storm precipitation from NOAA, both freezing rain and slee...
Figure 9.8 Snow storm, heavy snow can immobilize a region and paralyze a cit...
Figure 9.9 Annual mean snowfall days (
P
> 0.1 mm and
T
< 0°C) and freezing (
Figure 9.10 Three stages in the formation of an ice avalanche. First, it nee...
Figure 9.11 Schematic diagram of river‐ice jam and associated backwater leve...
Figure 9.12 Reported historical decline in ice‐cover duration in rivers in t...
Chapter 10
Figure 10.1 Linear trend of January (upper) and annual mean (lower) surface ...
Figure 10.2 Time series of the maximum depth of frost penetration; normalize...
Figure 10.3 Absolute change in permafrost extent (relative to 1986–2005) in ...
Figure 10.4 Projected changes in mean annual snow‐water equivalent (SWE) ove...
Figure 10.5 The partial correlations between active layer thickness in the p...
Figure 10.6 The correlation coefficients between soil moisture (SM) and obse...
Figure 10.7 The estimated permafrost extent based on Climatic Research Unit ...
Figure 10.8 The partial correlations between active layer thickness (ALT) in...
Figure 10.9 The estimated mean permafrost extent from six models' ensemble m...
Figure 10.10 The changes in active layer thickness for the three representat...
Figure 10.11 The partial correlations between active layer thickness in pers...
Guide
Cover Page
Title Page
Copyright Page
List of Figures
List of Tables
Foreword
Table of Contents
Begin Reading
Appendix
Bibliography
Index
Wiley End User License Agreement
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Climatology in Cold Regions
Chenghai Wang
School of Atmospheric Sciences
Lanzhou University
Lanzhou, China
This edition first published 2023© 2023 John Wiley & Sons, Inc.
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List of Figures
Figure 1.1 Definition of colds regions
Figure 1.2 Distribution of precipitation in cold regions
Figure 1.3 Distribution of near‐surface air temperature in cold regions
Figure 1.4 Comparisons of precipitation and near‐surface air temperature
Figure 2.1 The geopotential height and horizontal wind
Figure 2.2 Latitude–pressure cross–sections of temperature zonal deviation and vertical circulation
Figure 2.3 The variance of geopotential height in winter and spring
Figure 2.4 Fields of mean 30hPa geopotential height
Figure 2.5 The geopotential height and horizontal wind
Figure 2.6 Contrasting geopotential height fields
Figure 2.7 NOAA‐9 satellite image of the polar low’s incipient stage
Figure 2.8 Decadal variability of stratospheric sudden warming events
Figure 2.9 The vertical profile structure of the tropical cyclone‐like Tibetan Plateau vortex
Figure 2.10 Shear lines over the Tibetan Plateau
Figure 2.11 The various radiative mechanisms associated with aerosol‐cloud interactions
Figure 2.12 Distribution of summer precipitation
Figure 2.13 Distribution of winter snow depth
Figure 2.14 Distribution of permafrost and extent of permafrost and seasonally frozen ground
Figure 2.15 Distribution of climatological seasonally frozen ground
Figure 2.16 Distribution of snow melting
Figure 2.17 Spatial distribution of sea‐level pressure
Figure 3.1 Mean sea‐ice extent for March and September
Figure 3.2 Physiography of the Arctic lands
Figure 3.3 Distribution of Arctic polar desert and approximate southern limit of tundra
Figure 3.4 Average number of weeks of snow cover in the northern hemisphere
Figure 3.5 Annual evapotranspiration for the terrestrial region
Figure 3.6 Diagram of the soil freeze–thaw process
Figure 3.7 Seasonal cycle of the soil freeze–thaw process
Figure 3.8 Diurnal cycle of the soil freeze–thaw process
Figure 3.9 Variations of soil hydrothermal properties in the freeze–thaw process
Figure 3.10 Effects of the soil freeze–thaw process on soil‐moisture variations
Figure 3.11 Variations of monthly surface‐sensible and surface‐latent heat fluxes during the freeze–thaw process
Figure 3.12 Comparison of surface‐sensible heat flux (top panel) and surface‐latent heat flux
Figure 3.13 A schematic diagram summarizing the impacts of the soil freeze–thaw process on surface diabatic heating in the Tibetan Plateau
Figure 4.1 The forecast skill of model runs to predict average value in the Niño‐3.4 region
Figure 4.2 Distribution of stand deviation and climatological mean of snow depth in the Tibetan Plateau in the cold seasons
Figure 4.3 Evaluations of standardized series of snow depth average over the Tibetan Plateau
Figure 4.4 Relation between soil freeze‐thaw process and ASM
Figure 4.5 Series of soil moisture
Figure 4.6 Evolution of the in situ observation of soil moisture, precipitation, surface‐sensible and surface‐latent heating in Maqu
Figure 4.7 Distribution of summer precipitation root mean square error
Figure 4.8 Soil moisture memory
Figure 4.9 Spatial distribution of soil moisture memory
Figure 4.10 Soil moisture‐precipitation coupling in the northern hemisphere
Figure 4.11 Distribution of coupling index k between spring soil moisture and summer precipitation over the Tibetan Plateau
Figure 4.12 Evaluation characteristics of coupling relation between soil moisture and precipitation
Figure 5.1 Comparison of daily soil temperature and soil moisture between observation and simulation in the freeze–thaw process
Figure 5.2 Comparison of daily soil temperature and soil moisture in Maqu
Figure 5.3 Coupling relations between soil temperature and soil moisture during soil freezing and soil thawing
Figure 5.4 Diurnal cycles of soil temperature and soil moisture
Figure 5.5 Schematic diagram of water‐heat transport processes during soil freeze‐thaw
Figure 5.6 Schematic illustration of soil heat transport and soil water transport
Figure 5.7 Impacts of soil water content and soil textures on hydraulic conductivities in the FCS scheme
Figure 5.8 Vertical sections of daily soil temperature and soil volumetric water
Figure 5.9 Vertical sections of daily soil moisture and soil ice
Figure 6.1 Runoff components of continental glacial watershed
Figure 6.2 Water cycle in glacial areas
Figure 6.3 The confluence of snow meltwater (rainfall)
Figure 6.4 Responses of snowmelt runoff to air temperature increases
Figure 6.5 Location map of the Spiti Basin
Figure 6.6 Process of groundwater supply in a cold area
Figure 6.7 Influencing factors of hydrological freezing–thawing process of river ice
Figure 6.8 Research domains
Figure 6.9 Distribution of the mean precipitation rate over the Tibetan Plateau
Figure 6.10 Evolution of precipitation rate in four regions
Figure 6.11 The evolution of seasonal precipitation rate
Figure 6.12 Monthly variability of precipitation rate
Figure 7.1 Spatial patterns of vegetation phenology
Figure 7.2 Circumpolar maximum normalized difference vegetation index of Arctic tundra
Figure 7.3 Plant physiognomy occurring in different tundra bioclimate subzones
Figure 8.1 The model and mainly influencing factors for carbon transfer in the earth surface ecosystems
Figure 8.2 Schematic diagram of annual carbon partitioning in diverse carbon pools in the Kalancohe humilis meadow plant–soil system representing the annual carbon fluxes
Figure 8.3 A conceptual model of the relationship among the environmental conditions and the contribution of variables to the soil organic carbon stocks
Figure 8.4 Key disturbance processes impacting northern high‐latitude soil organic carbon stocks and feedback between them
Figure 8.5 The distribution of soil organic carbon
Figure 8.6 The current state of the Arctic carbon cycle
Figure 8.7 Key factors controlling gaseous nitrogen input and output pathways in Arctic ecosystems
Figure 9.1 Snowmelt flooding
Figure 9.2 The formation of the snowmelt flood
Figure 9.3 Freezing rain
Figure 9.4 Freezing rain falls when snowflakes melt completely before reaching the surface
Figure 9.5 Distribution of freezing rain in China
Figure 9.6 Maximum number of hours with freezing rain in a single year
Figure 9.7 Winter storm precipitation from the US National Oceanic and Atmospheric Administration
Figure 9.8 Snow storm
Figure 9.9 Annual mean snowfall days and freezing days
Figure 9.10 Three stages in the formation of an ice avalanche
Figure 9.11 Schematic diagram of river‐ice jam and associated backwater level
Figure 9.12 Reported historical decline in ice cover duration in rivers in the northern hemisphere
Figure 10.1 Linear trend of January and annual mean surface air temperature
Figure 10.2 Time series of the maximum depth of frost penetration
Figure 10.3 Absolute change in permafrost extent in the northern hemisphere
Figure 10.4 Projected changes in mean annual snow‐water equivalent over northern‐hemisphere land
Figure 10.5 The partial correlations between active layer thickness in the persistent permafrost regions and snow depth
Figure 10.6 The correlation coefficients between soil moisture and observed maximum depth of frozen ground
Figure 10.7 The estimated permafrost extent based on Climatic Research Unit data and Kudryavtsev's method
Figure 10.8 The partial correlations between active layer thickness in the persistent permafrost regions and snow depth
Figure 10.9 The estimated mean permafrost extent from six models' ensemble mean for the three representative concentration pathways
Figure 10.10 The changes in active layer thickness
Figure 10.11 The partial correlations between active layer thickness in persistant permafrost regions
List of Tables
Table 3.1 Major Arctic ice caps
Table 5.1 The root mean square errors of soil temperature and soil moisture
Table 5.2 The magnitude of variables in the heat conduction equation
Table 5.3 The magnitude of variables in the modified Richards equations
Table 7.1 Ecosystem cover area
Table 8.1 Carbon and nitrogen and its transport to atmosphere in cold regions
Foreword
It is a difficult task to put fragments together. Because I have been lived and worked in cold and arid regions, my research interests have focused on climate in cold and arid regions for a long time. Witnessing the wealth of exceptional research results in this fields in the past decades, I have dreamed of assembling the results of research in cold regions into a book to promote the formation of a systematic discipline.
This dream came true with the assistance of my students. Some former students have now become my colleagues. They follow me to work and study in this field and their insightful opinions and contributions are the basis of this book. I am extremely grateful for all of their contributions to the quality of this book. I would like to thank specifically Dr. Kai Yang and Dr. Feimin Zhang. They sacrificed a lot of spare time for the completion of this book.
I sincerely thank the teachers who guided and unselfishly helped me enter this research field, Professors Youxi Gao, Dahe Qin, Guodong Cheng, and Professor Wenjie Dong. I was afraid that the book might not live up to their expectations, so I dropped the idea of inviting one of them to write the foreword at the last minute. I am also grateful to many experts and friends who previously supported me and gave me attention.
As this is a new subject, I am only the first visitor to this particular scene. I was almost lost in the splendid view, that was created by many scientists. I am aware that I have omitted a lot of important works, and have overlooked some citations. I hope that publishing this book reflects the Chinese adage: “throw a brick to meet jade”. I also expect feedbacks and corrections from peers.
Finally, I thank Dr. Collison Mandy and Ms. Rosemary Morlin for their careful and professional assistance with this book, one of the important reasons why I chose Wiley to publish it. I also thank Dr. Frank Weinreich for designing an eye‐catching book cover.
September, 2022Chenghai WangLanzhou University, China
1Introduction
Cold regions are special geographical units. Climate characteristics on a global scale have become well known, and regional climate features have attracted more and more attention. This chapter will define cold regions and types of cold region, to illustrate climate characteristics in cold regions.
1.1 Scope of the Subject and Recent Highlights
The climate in cold regions is a new, interdisciplinary, subject that is concerned with the climate and its changing in cold regions over the earth's atmosphere system. It focuses mainly on the high‐altitude and middle‐high latitude geosciences and their physics, chemistry, and fluid dynamics.
Many large‐scale climate characteristics and processes have been understood, particularly entering the twentieth century. The climate in limited and specific regions has recently attracted more attention.
There is mounting evidence that the human activity has been influenced on climate, ecology, and biological systems. For instance, researches revealed that temperatures in cold region rise twice times faster than in middle‐low latitude; the minimum temperature rises faster than the maximum temperature. One more concern is the glacial retreat and frozen ground degeneration. Particularly, more and more attention has been paid to Arctic ice melting and relevant Arctic “amplification.”
Cold is a relative concept to warm. The climate can be roughly divided according to latitude into tropical, sub‐tropical, and polar regions. Cold regions can be found in high‐altitude areas and high mountains. Cold means a region lacks of solar radiation, in other words, where surface‐energy negative balance, and absorbing less radiation than what is being released.
Owing to latitudinal differences in the distribution of solar radiation, there are latitudinal zones in which the climate varies regularly according to latitude. The globe is divided into multiple climate zones according to the characteristics of fundamental climate indicators (e.g. surface air temperature and precipitation).
1.2 Some Definitions and Terms of Reference
Cold regions are generally regarded as the cold environment in both polar regions and more temperate locations, which are closely related to the cryosphere. The areas covered by stable seasonal snow is usually considered to be cold regions. In that case, cold regions consist of glaciers, ice caps, stable snow, permafrost, sea ice, river ice, and lake ice. Based on this definition, the global area of cold region is about 0.6 × 108 km2, covering 40% of land area. This does not include unstable snow areas and parts of the seasonally frozen ground (SFG) area.
There is no universally accepted definition of cold regions. In the early stage, Köppen (1936) defined the cold region in Canada by the criterion that the near‐surface air temperature in the coldest month of a year is lower or equal to −3 °C and there are only four months with a monthly average temperature greater than 10 °C. Gerdel (1969) considered the annual mean temperature zero isothere for the division of cold regions and other regions. Wilson (1967) suggested a definition that cold regions with temperature and precipitation as essential elements. Yang (1997) used Köppen's criterion and defined the cold region in China as a places where rivers and lakes were frozen for more than 100 days. Based on these definitions, Chen et al. (2005) used the criterion that the near‐surface air temperature in the coldest month of the year is lower or equal to −3°C, that the number of months with a monthly mean temperature greater than 10°C is not more than 5 months, and that the annual mean temperature is lower or equal to 5°C to define the cold region in China, along with the distribution and boundary of permafrost, glaciers, stably seasonal snow and vegetation zone divisions.
Large amounts of permafrost, SFG, and snow are distributed in cold regions. A distinct feature of land surface in cold regions is that the soil freezes and then thaws. The seasonal variation of the thermal regime in soil is a representative indicator of the climate in a cold region. In climatology, similar to Gerdel's (1969) definition, the area north of the zero isothere of the January mean temperature is defined as a cold region (Figure 1.1). This region not only includes the area defined by Köppen as the circumpolar cold region, but also contains the areas with the coldest monthly mean temperature of nearly 0°C.
Figure 1.1 Definition of cold regions. The areas colored purple are the cold regions defined by Köppen's (1936) criterion. The areas where soil undergoes seasonally freeze–thaw is also regarded as the cold regions (slate blue). The thick and thin lines are the zero isotherms of the January temperature and annual mean temperature, respectively, which define the SFG and permafrost regions.
The circumpolar cold region generally covers the permafrost area and the cold region in mid‐latitudes covers the SFG. The freeze–thaw cycle in frozen ground can influence the soil water–heat transport and variations of surface diabatic heating, which makes the climate in this region unique. Consistent with essential awareness, this region is also covered with snow, while the maximum extent of the snow‐cover basically coincides with the zero isothere line of the January mean temperature. In the next two sections, the characteristics of climatology in cold regions and their discrepancies will be introduced.
1.3 Characteristics of Cold Region Climatology
According to the world map of the Köppen‐Geiger climate classification, cold regions cover the snow zone, and polar zone, while the characteristics of precipitation and temperature have distinct spatial differences. The climate in a region is closely related to the variations of precipitation and temperature, which decides the formation of the climate. The following sections will analyze the temporospatial characteristics of precipitation and near‐surface temperature.
Compared with the tropics or summer monsoon regions, precipitation in winter and summer in cold regions is obviously less, especially in winter. The annual mean precipitation in winter is only 100–150 mm; in the western region of North America, eastern Siberia, and northern China, the annual mean winter precipitation is less than 50 mm. In summer, the mean precipitation is about 200–300 mm (Figure 1.2). The distribution of precipitation has a distinct spatial difference in cloud regions; precipitation in high latitudes is greater than mid‐latitudes in both winter and summer, and winter precipitation in northern Europe is greater than summer precipitation.
Figure 1.2 Distribution of precipitation (mm) for winter (December–January–February, DJF) (upper panel) and summer (June–July–August, JJA) (lower panel) in cold regions in the northern hemisphere.
The near‐surface temperature in the cold regions has distinct seasonal differences, and temperature has zonal variations, decreasing from low latitudes to high latitudes (Figure 1.3), decreasing from low attitude to high altitude. The near‐surface temperature is related to the land surface processes (e.g. snow, frozen ground).
Figure 1.3 Distribution of near‐surface air temperature (°C) for January (top panel), July (middle panel), and annual mean (bottom panel) in the cold regions of the northern hemisphere.
1.4 Classification of Cold Region Climatology
The climate in different parts of cold regions has heterogeneity. Global cold regions are mainly located in high latitudes and high altitudes. According to the types of frozen ground, cold regions can be divided into two categories, the region of permafrost, the region of SFG. According to geographical location, it can be divided into polar regions, middle and high latitudes and the Tibetan Plateau (TP).
Comparing the precipitation and near‐surface temperature between high latitudes (permafrost) and SFG and high altitudes (TP) in cold regions (Figure 1.3), it is clear that near‐surface monthly temperature in the permafrost region is generally lower than that in the SFG region, and the near‐surface temperature in the TP is higher than that in the permafrost region, but lower than the SFG. The variations of near‐surface temperature have distinct dependence on the variation of latitude and altitude. Monthly precipitation in the SFG region is generally greater than that in the permafrost region, which is concentrated between August and October. In contrast, the precipitation in the permafrost region is concentrated between July and September. The annual variations of precipitation in the TP are different from the cold regions in high latitudes (Figure 1.4), which is concentrated between June and August with 60–80 mm, and its amount is more than that of the permafrost regi9on, but less than that in the SFG region. Comparing the annual variations of precipitation and near‐surface temperature in three different cold regions, these have distinct differences.
Figure 1.4 Comparisons of precipitation (mm, bar) and near‐surface air temperature (°C, lines) between high latitudes (permafrost and seasonally frozen ground regions, SFG) and high altitudes (e.g. the Tibetan Plateau (TP) cold regions.)
1.5 Impacts of Cold Regions on Global Climate
On average, the temporal scale of energy (heat), moisture, and matter exchange between land and atmosphere is shorter in tropical regions than in cold regions. In tropical regions, the interaction between oceanic surface and atmosphere has quick and short features. In tropical areas, convective precipitation and tropical cyclones (typhoons or hurricanes) frequently occur, with smaller monthly change features than in sub‐tropical regions. These probably result as an energy export to tropical regions. The natural function of the cold region is energy accumulation. According to this theory, cold regions are more likely to be energy hoarders for seasonal climate change.
1.6 Contribution of Cold Regions to Energy Balance
In addition to the energy from solar radiation, there is an inter‐conversion of energy in the climate system. The heat absorbed from surface radiation balance in colds region would be stored in soil, and adjusted by soil moisture and soil heat capacity of different soil kinds. Energy will be released to the atmosphere in summer when evaporation becomes significant. This process corresponds to ground freezing and thawing.
1.7 Summary
Cold regions play the role of “storage/memory” and “repeater” of energy in the climate system.
Cold regions contain the effective and crucial signals (e.g. soil moisture, snow, etc.) for climate prediction.
The moisture and energy cycle in cold regions are promotors for climate change.
Global changes intensify climate changes in cold regions.
Changes in the biology and vegetation in cold regions are more evident and vulnerable.
Climate changes in cold regions result from dynamics and hydrothermal processes.
Cold regions are the main and essential water sources of global fresh water, and is also the main source of global rivers.
Engineering and construction risks in cold regions are increasing.
Whether global warming has changed carbon and nitrogen cycles in cold regions.
What are the hazards and risks in cold regions?
2Climate in Cold Regions
The climate in cold regions is unique. Due to low temperature, the weather, climate, and their influence factors have obvious differences compared with subtropical and tropical regions. This chapter will present the characteristics of general circulation, the primary weather systems (including permanence and semi‐permanence), meteorological elements, land–surface (including the distribution and changes in frozen ground, hydrology, and ecology), and winter monsoon in cold regions.
2.1 General Circulation
There are two basic factors that determine the general circulation of Earth's atmosphere: the latitudinal change of energy received by the Earth–atmosphere system from the sun, and the global uneven distribution of angular momentum of the atmosphere. General circulation on Earth is mainly due to the uneven heating of Earth's surface. Earth's energy balance involves incoming solar radiation and outgoing thermal infrared radiation, which together make up the net radiation (NR). The global general circulation, pressure distribution, and wind patterns play an integral role in the heat balance of Earth and the formation of global ocean currents. In the global general circulation, warm air is transported from lower latitudes to higher latitudes and cold air from high latitudes to low latitudes. This heat exchange keeps global temperatures relatively stable. Low latitude region is net incoming energy throughout the year as a result of continuous heating, and high latitudes are colder as a result of continuous cooling due to net energy loss.
Cold region is a product of regional stationary sate in general circulation evolution. According to the results of geopotential height and horizontal wind at 500 hPa in January and July (Figure 2.1), the general circulation shows obvious seasonal variations. The position of the troughs and ridges in winter and summer is essentially stable or gradually changing, and occupies a considerable part of the year, while it is relatively short in the two transitional seasons (spring and autumn). In midlatitudes, the general circulation behaves as a flat westerly wind; both the geopotential height and the horizontal wind in winter (January) are larger than that in summer (July). That is, there are three troughs and ridges in the middle‐high latitude westerly in winter, four trough ridges in summer, the so‐called “winter three summer four”.
Figure 2.1 The geopotential height (contour) and horizontal wind (vector) at 500 hPa level in January (left panel) and July (right panel).
Due to the spatial discontinuity of global cold regions, some cold regions have regional and local scales, which induce local circulation/cells, for instance, the valley wind. A prominent regional scale cell is the monsoon cell on the southern Tibetan Plateau (TP), which is an indicator for the eastern Asian summer monsoon (Figure 2.2). Also, the high altitude of TP induces many regional cells in the west, east, north, and south directions, which influence the climate regime in eastern Asia. While wind passes across the underlying surface in cold regions, such as glaciers, snow, and frozen ground, in most cases, it would produce a downward flow along slopes, and produce a strong blast wind at valley export.
Figure 2.2 Latitude–pressure cross‐sections (80–100E mean) of zonal temperature deviation (shading, K) and vertical circulation (vector, m s−1; the vertical velocity multiplied by 5) calculated from ERA‐Interim (1981–2010 climatology) in May (left panel) and JJA (right panel). Black shading represents the topography.
The evolution of general circulation across cold region is greatly influenced by the variations of surface diabatic heating. As a typical cold region, TP is a huge, elevated heating source in summer, which can result in convergence in the lower atmosphere and divergence in the higher atmosphere in the eastern TP and the reversed structure in the western TP.
2.2 Low and High Frequency of Atmosphere Variability
Generally, the atmospheric motion can be divided into low‐ and high‐frequency variability according to its temporal scales. Comparing the variance of unfiltered, 2.5–6‐day filtered and 10–90‐day filtered geopotential height in winter and spring (Figure 2.3), the three centers are the North Pacific, North Atlantic, and Siberia, respectively. The evolution pattern of the 10–90‐day filtered geopotential height is more similar to the evolution pattern of the unfiltered geopotential height than that of the 2.5–6‐day filtered geopotential height. Meanwhile, the evolution of the 10–90‐day filtered geopotential height is much greater than that of the 2.5–6‐day filtered geopotential height. It means that the interannual variation of geopotential height in winter and spring is mainly dominated by the low‐frequency variation. It is also clear that the evolution of the 10–90‐day filtered geopotential height is the main contributor over the land around North Atlantic and Siberia. This suggests that land surface in cold regions plays an important role in the interannual variation of geopotential height in winter. This effect can persist from winter to spring, though the distribution of geopotential height evolution has some changes due to land‐sea thermodynamic differences.
Figure 2.3 The variance of the unfiltered (left), 2.5–6‐day filtered (middle), and 10–90‐day filtered (right) geopotential height at 100 hPa in winter (upper) and spring (down).
The 2.5–6‐day eddy is very active in the middle‐latitude North Pacific and North Atlantic, which is called the storm track. The 2.5–6‐day eddy involves strong momentum and heat flux, which causes the interaction between wave and flow in the storm track. The low‐frequency Rossby wave train triggered by the cold region in spring can persist during the summer and influence the summer climate.
2.3 Weather Systems
General circulation usually refers to the mean state of large‐scale and long‐term atmospheric movement or the change processes of a specific period. The horizontal scale is larger than several thousand kilometers; the vertical scale is larger than 10 km, and the temporal scale is longer than 106 hours. This large‐scale atmospheric movement not only restricts large‐scale weather changes, but is also one of the basic factors of climate formation. Abnormal changes in general circulation will inevitably lead to abnormal weather and climate. The so‐called weather system is a large part of the general circulation, and a particular weather process is the background of a specific general circulation. The typical weather systems include polar vortex, polar jet, blocking high, polar low, polar high, and related phenomena (e.g. sudden stratosphere warming). These will be introduced in detail as follows.
2.3.1 The Polar Circulation
Geographically, the regions north of 66.5°N and south of 66.5°S are defined as polar regions. The Arctic region is basically the ocean except Greenland, and the Antarctic region is mainly land. The Arctic Ocean is an ocean of icebergs, but it is not completely frozen even in winter. The mean atmospheric heat in polar region is net spending, thus the polar is the atmospheric cold source, medium and low latitudes of heat transports to polar regions via the average meridian circulation and the large eddies, causing the loss of heat in the polar region, forms cold air, and affects weather and climate in middle and low latitudes by heat exchange.
The mean winter circulation at sea level is dominated by three “centers of action”: the Icelandic Low off the east coast of southern Greenland, the Aleutian Low south of Alaska, and the Siberian High in central Eurasia (Figure 2.4). The mean Icelandic and Aleutian lows are maintained largely by their position downstream of major mid‐tropospheric stationary troughs, surface heating contrasts and regional cyclone development. The winter Siberian High is a low‐level system driven by radiative cooling. During summer, the Icelandic and Aleutian lows are much weaker and the Siberian High is replaced by mean low pressure. Weak low pressure also characterizes in the central Arctic Ocean in summer, compared with a pronounced high over the Beaufort Sea in spring. Winter cyclone activity on the Pacific side of the Arctic is most common in the vicinity of the AL.
Figure 2.4 The mean geopotential height (gpm) at 30 hPa for the four mid‐season months over 1981–2010, based on ERA5 data.
The polar vortex is a large area of low pressure and cold air surrounding both poles and it commonly exists near them. It is weaker in summer but stronger in winter. The term “polar vortex” dates back to at least 1853, but was revised in 1950, to become the “circumpolar vortex” and reverted to “polar vortex” in1959. The term “vortex” refers to the counter‐clockwise flow of air that helps to keep the colder air near the poles. Commonly, during winter in the northern hemisphere, the polar vortex will expand, sending cold air southward with the jet stream. The polar vortex is the strong, midlatitude, quasi‐zonal, stratospheric wind system that develops during winter as a result of the latitudinal gradient of ozone heating. In the literature of atmospheric science, the term “polar vortex” is mostly used as an abbreviation for circumpolar vortex and refers to a planetary‐scale westerly (west to east) flow that encircles the pole in middle or high latitudes. The latitude at which the zonal wind reaches its hemispheric maximum can be considered as the approximate border of a polar vortex. Vortex in the troposphere is much larger than the vortex in the stratosphere. Furthermore, the tropospheric vortex exists for the whole year, but the stratospheric polar vortex exists only from autumn to spring (Waugh et al. 2016).
2.3.2 Polar Vortex
In winter, the height field points to strong westerlies in the middle and upper stratosphere flowing around a deep cold vortex (Figures 2.5 and 2.6). Above 40 km (3 hPa, not shown), there is a polar westerly jet stream in midlatitudes. Peak mean zonal winds of 80 m s−1 actually occur around 60–70 km altitude (at 0.3–0.1 hPa level). The winter mean stratospheric vortex is broadly symmetric, but is less symmetric than its Antarctic counterpart. There are troughs over eastern Asia and eastern North America and a weak ridge located in western North America. In spring, the vortex center has shifted well off the North Pole to north‐central Eurasia.
Figure 2.5 The geopotential height (gpm, color) and horizontal wind (m s−1, vector) at 100 hPa in January (left panel) and July (right panel).
Figure 2.6 Comparison of geopotential height fields at the lower jet stream level (500 hPa) with low values in purple and the jet stream in white. (a) A single and more west to east jet stream encircling the pole contrasts with a more undulating configuration (b) with multiple low height centers (dark purple).
Source: NOAA / United States Department of Commerce / Public Domain.
In summer and above about 20 km, the cyclonic vortex has broken down. There is a polar easterly flow around a highly symmetric warm polar anticyclone. In summer, an easterly jet of about 60 m s−1 at 0.1 hPa is located at around 50–60°N. The October wind field illustrates the transition back toward the winter situation.
As discussed by Holton (2004) and Andrews et al. (1987), if there were no transports arising from the breaking of atmospheric waves propagating upwards from the troposphere, the zonal‐mean temperature of the stratosphere would become to a radiatively determined state, with the temperature distribution corresponding to an annually varying thermal equilibrium that follows the annual cycle in solar heating. The circulation would hence represent a zonal–mean flow in balance with the meridional temperature gradient (a thermal wind balance), with essentially no meridional or vertical circulation and no stratosphere–troposphere exchange. The existence of a westerly vortex in winter and an easterly vortex in summer is qualitatively that expected from radiative equilibrium, the climate in winter actually shows considerable departures from the radiatively determined state. In the 30–60 km region, the change in temperature from the winter pole to the summer pole is much smaller than the radiatively determined gradient. This is due to eddy transports that drive the flow away from a state of radiative balance. By contrast, the departure from radiative equilibrium in summer is small, implying a reduction of transport.
In the troposphere, these eddy transports are largely associated with moving synoptic‐scale waves (they are associated with migrating cyclones and anticyclones at the surface). By contrast, winter transports in the stratosphere are associated with the long planetary waves, especially for wavenumbers 1 and 2, which can penetrate into the stratosphere under specific conditions. Wavenumber refers to the number of atmospheric waves around a latitude circle. For any given day, Fourier transform methods can be used to break down the total circulation at a given level in the atmosphere with respect to the relative contributions of long planetary waves (lower wavenumbers), which move slowly or remain stationary with respect to the Earth's surface, and shorter waves (higher wavenumbers). A strong wavenumber two component, for example, would have a pronounced expression of two ridges and two troughs, each separated by 180° longitude. If we look at the circulation on a typical winter day, we can see that while the troposphere has a strong contribution from the higher wavenumbers, the circulation of the stratosphere is more symmetric, indicating that the shorter waves are not readily penetrating into the stratosphere. For long‐time averages, the effects of the shorter waves are marginal. But even with such time averaging, the winter circulation of the stratosphere is more symmetric than that of the troposphere. This can be related to the impacts of orography and land‐sea thermal contrasts (Pawson and Kubitz 1996).
Waves will penetrate (vertically propagate) into the stratosphere provided that the zonal‐mean wind is positive (the wind averaged around a latitude circle blows from west to east) but less than a critical value that depends strongly on the length of the waves. The longer the wave (that is, the lower the wavenumber), the stronger the zonal wind can be and still allow for vertical propagation. Since the zonal wind tends to be at a maximum near the tropopause in middle and high latitudes (much lower in altitude than the winter stratospheric jet), the shorter waves in winter tend to be “trapped” in the troposphere. In summer, the zonal–mean zonal winds in the stratosphere are easterly. There can be no vertical propagation of waves in these situations. This is consistent with the fact that the summer stratospheric circulation is highly symmetric with a thermal structure close to that expected from radiative equilibrium.
The typical thermal structure of the winter stratosphere features low temperatures in the vortex core in the lower stratosphere and increases outward from the axis of rotation. Conversely, there are high temperatures in the vortex core of the upper stratosphere and lower mesosphere that decrease outward. The vertical change in the temperature structure is attributed to warming through large‐scale subsidence in the mesosphere and radiative cooling in the lower stratosphere during the polar night (Gerrard et al. 2002). The stratospheric jet maximum is located at the level of the temperature transition region in the vertical (10–l hPa). Positive potential vorticity (PV) anomalies are found in the vortex core at the level of the thermal transition zone, with a cold trough below and a warm ridge above. The vortex winds, which are formed between the two thermal regimes, circulate around the core. PV represents the specific volume (volume per unit mass of air) times the scalar product of the absolute vorticity (the sum of relative and planetary vorticity) and the gradient of potential temperature. The PV maximum arises from the low density of the air at stratospheric levels (specific volume is large), the strong increase of potential temperature with height (strong stability), and the positive absolute vorticity.
Climatology of the stratospheric polar vortex in both hemispheres has been discussed based on areal extent (Baldwin and Holton 1988) and on elliptical diagnostics (Waugh 1997; Waugh and Randel 1999). Fitting an ellipse to a specified contour allows one to summarize various measures of the vortex. These include the equivalent latitude of the vortex edge (defining its area), the offset of the vortex from the pole, the longitude of its center and its elongation (aspect ratio or ellipticity). Waugh and Randel (1999) defined these parameters on the basis of selected PV contours for the 440 K (lower stratosphere) to 1300 K (upper stratosphere) potential temperature, average from October 1978 to April 1998.
The equivalent latitude of the vortex edge is lower (i.e. the vortex covers a larger area) in winter compared to autumn and spring. Area increases with height for January, the equivalent latitude is around 70°N in the lower stratosphere, compared to about 60°N in the upper stratosphere. Depending on height and season, the vortex is centered 6–18° off the pole. The vortex tilts westward with a height between 500 and 1300 K by about 60°. The vortex also tends to be elongated in shape. In November–December, there may be rapid eastward and then westward shifts of the vortex center in the lower stratosphere between 60 and 120°E and 60–90°W, apparently related to changes in the planetary waves. Similar but fewer shifts occur in late January to early February.
The Arctic vortex is smaller and breaks down earlier in summer than its Antarctic counterpart. It is also highly variable throughout its lifecycle, especially in late winter. This variability is partly caused by extreme events distorting the vortex. A distorted vortex was present in about half of the winters from January 1990 to August 1997 (Waugh and Randel 1999). There were also other periods when the vortex was symmetrical and quiescent. Interannual variability in the stratospheric polar vortex has been linked to the phase of the equatorial quasi‐biennial oscillation (QBO), as studied by Labitzke (1982). The QBO refers to an oscillation in the zonal winds of the equatorial stratosphere having a period that fluctuates between about 24 and 30 months. Stratospheric equatorial easterlies associated with the QBO favor a weaker, warmer vortex than equatorial westerlies, but the relationship has broken down in some winters.
The primary feature of the mid‐tropospheric (500 hPa) circulation in northern high latitudes is a well‐developed cyclonic vortex. The winter vortex is strongly asymmetrical. At 60°N, the mean flow is characterized by major troughs over eastern North America and eastern Asia and a weaker trough over western Asia (the Urals trough). A strong ridge is located over western North America, with much weaker ridges over the eastern Atlantic and central Asia. The lowest geopotential heights are found in the Canadian Arctic Archipelago. These features are related to dynamical forcing by orography (most importantly, the Rocky Mountains and the Himalayas which disturb the circulation), thermal forcing associated with the large‐scale land‐ocean distribution (warm oceans versus cold land) and radiative forcing: the stronger winter asymmetry at 500 hPa as compared to the stratospheric vortex, and the large relative shift in the vortex center. The more symmetrical nature of the stratospheric winter vortex relates to the fact that only the long planetary waves can propagate up to high levels. The three troughs and three ridges that are fairly well expressed at about 60°N indicate that the total circulation at this latitude has a strong wavenumber three component. In summer, the 500 hPa vortex is much weaker in response to the more even latitudinal distribution of radiation and temperature. It is also more symmetric than its winter counterpart, with the lowest pressure heights centered over the pole.
The mean thickness of the tropopause layer is only about 600 m over the Kara‐Laptev seas in January, whereas in summer it exceeds 1.6 km in a belt from northern Norway to northern Novaya Zemlya and eastward to Wrangel Island (Makhover 1983). There is a secondary minimum (maximum) in tropopause height in October (December–January). By contrast, the tropopause temperature shows a simple annual cycle with a minimum of −62 °C in January and a maximum of −49 °C in July, apparently following the annual cycle of temperature in the lower stratosphere as identified for the 100 hPa (Wilson and Godson 1962). Highwood et al. (2000) also suggested that tropopause temperatures in winter are lower in the Eurasian sector and higher in the Canadian‐Atlantic sector.
The existence of a double maximum/minimum in tropopause height, is attributed to several factors. The primary August maximum is attributed to the lag in the heating of the surrounding landmasses, while the second one in January is related to enhanced meridional heat transport by the planetary waves. The spring minimum is associated with the seasonal weakening of this transport. The autumn minimum relates to the cooling of the landmasses and a decrease in cyclone activity compared with late summer. Two patterns of the annual cycle in tropopause height are identified in different regions. Over North America and subarctic Siberia there is a simple annual wave (winter/summer–tropopause pressure maximum/minimum) while in the High Arctic, northern Europe, and western Siberia there is a double wave (spring/autumn – maxima, summer/winter – minima).
