- •Preface
- •1 A Voyage of Discovery
- •1.2 Goals
- •1.3 Organization
- •1.4 The Big Picture
- •1.5 Further Reading
- •2 The Historical Setting
- •2.2 Eras of Oceanographic Exploration
- •2.3 Milestones in the Understanding of the Ocean
- •2.4 Evolution of some Theoretical Ideas
- •2.5 The Role of Observations in Oceanography
- •2.6 Important Concepts
- •3 The Physical Setting
- •3.1 Ocean and Seas
- •3.2 Dimensions of the ocean
- •3.3 Sea-Floor Features
- •3.4 Measuring the Depth of the Ocean
- •3.5 Sea Floor Charts and Data Sets
- •3.6 Sound in the Ocean
- •3.7 Important Concepts
- •4.1 The Earth in Space
- •4.2 Atmospheric Wind Systems
- •4.3 The Planetary Boundary Layer
- •4.4 Measurement of Wind
- •4.5 Calculations of Wind
- •4.6 Wind Stress
- •4.7 Important Concepts
- •5 The Oceanic Heat Budget
- •5.1 The Oceanic Heat Budget
- •5.2 Heat-Budget Terms
- •5.3 Direct Calculation of Fluxes
- •5.4 Indirect Calculation of Fluxes: Bulk Formulas
- •5.5 Global Data Sets for Fluxes
- •5.6 Geographic Distribution of Terms
- •5.7 Meridional Heat Transport
- •5.8 Variations in Solar Constant
- •5.9 Important Concepts
- •6.2 Definition of Temperature
- •6.4 The Oceanic Mixed Layer and Thermocline
- •6.5 Density
- •6.6 Measurement of Temperature
- •6.7 Measurement of Conductivity or Salinity
- •6.8 Measurement of Pressure
- •6.10 Light in the Ocean and Absorption of Light
- •6.11 Important Concepts
- •7.1 Dominant Forces for Ocean Dynamics
- •7.2 Coordinate System
- •7.3 Types of Flow in the ocean
- •7.4 Conservation of Mass and Salt
- •7.5 The Total Derivative (D/Dt)
- •7.6 Momentum Equation
- •7.7 Conservation of Mass: The Continuity Equation
- •7.8 Solutions to the Equations of Motion
- •7.9 Important Concepts
- •8.2 Turbulence
- •8.3 Calculation of Reynolds Stress:
- •8.4 Mixing in the Ocean
- •8.5 Stability
- •8.6 Important Concepts
- •9 Response of the Upper Ocean to Winds
- •9.1 Inertial Motion
- •9.2 Ekman Layer at the Sea Surface
- •9.3 Ekman Mass Transport
- •9.4 Application of Ekman Theory
- •9.5 Langmuir Circulation
- •9.6 Important Concepts
- •10 Geostrophic Currents
- •10.1 Hydrostatic Equilibrium
- •10.2 Geostrophic Equations
- •10.3 Surface Geostrophic Currents From Altimetry
- •10.4 Geostrophic Currents From Hydrography
- •10.5 An Example Using Hydrographic Data
- •10.6 Comments on Geostrophic Currents
- •10.7 Currents From Hydrographic Sections
- •10.8 Lagrangian Measurements of Currents
- •10.9 Eulerian Measurements
- •10.10 Important Concepts
- •11.2 Western Boundary Currents
- •11.4 Observed Surface Circulation in the Atlantic
- •11.5 Important Concepts
- •12 Vorticity in the Ocean
- •12.2 Conservation of Vorticity
- •12.4 Vorticity and Ekman Pumping
- •12.5 Important Concepts
- •13.2 Importance of the Deep Circulation
- •13.3 Theory for the Deep Circulation
- •13.4 Observations of the Deep Circulation
- •13.5 Antarctic Circumpolar Current
- •13.6 Important Concepts
- •14 Equatorial Processes
- •14.1 Equatorial Processes
- •14.6 Important Concepts
- •15 Numerical Models
- •15.2 Numerical Models in Oceanography
- •15.3 Global Ocean Models
- •15.4 Coastal Models
- •15.5 Assimilation Models
- •15.6 Coupled Ocean and Atmosphere Models
- •15.7 Important Concepts
- •16 Ocean Waves
- •16.1 Linear Theory of Ocean Surface Waves
- •16.2 Nonlinear waves
- •16.3 Waves and the Concept of a Wave Spectrum
- •16.5 Wave Forecasting
- •16.6 Measurement of Waves
- •16.7 Important Concepts
- •17 Coastal Processes and Tides
- •17.1 Shoaling Waves and Coastal Processes
- •17.2 Tsunamis
- •17.3 Storm Surges
- •17.4 Theory of Ocean Tides
- •17.5 Tidal Prediction
- •17.6 Important Concepts
- •References
Chapter 16
Ocean Waves
Looking out to sea from the shore, we can see waves on the sea surface. Looking carefully, we notice the waves are undulations of the sea surface with a height of around a meter, where height is the vertical distance between the bottom of a trough and the top of a nearby crest. The wavelength, which we might take to be the distance between prominent crests, is around 50-100 meters. Watching the waves for a few minutes, we notice that wave height and wave length are not constant. The heights vary randomly in time and space, and the statistical properties of the waves, such as the mean height averaged for a few hundred waves, change from day to day. These prominent o shore waves are generated by wind. Sometimes the local wind generates the waves, other times distant storms generate waves which ultimately reach the coast. For example, waves breaking on the Southern California coast on a summer day may come from vast storms o shore of Antarctica 10,000 km away.
If we watch closely for a long time, we notice that sea level changes from hour to hour. Over a period of a day, sea level increases and decreases relative to a point on the shore by about a meter. The slow rise and fall of sea level is due to the tides, another type of wave on the sea surface. Tides have wavelengths of thousands of kilometers, and they are generated by the slow, very small changes in gravity due to the motion of the sun and the moon relative to earth.
In this chapter you will learn how to describe ocean-surface waves quantitatively. In the next chapter I will describe tides and waves along coasts.
16.1Linear Theory of Ocean Surface Waves
Surface waves are inherently nonlinear: The solution of the equations of motion depends on the surface boundary conditions, but the surface boundary conditions are the waves we wish to calculate. How can we proceed?
We begin by assuming that the amplitude of waves on the water surface is infinitely small so the surface is almost exactly a plane. To simplify the mathematics, we can also assume that the flow is 2-dimensional with waves traveling in the x-direction. We also assume that the Coriolis force and viscosity can be neglected. If we retain rotation, we get Kelvin waves discussed in §14.2.
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CHAPTER 16. OCEAN WAVES |
With these assumptions, the sea-surface elevation ζ of a wave traveling in the x direction is:
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where ω is wave frequency in radians per second, f is the wave frequency in Hertz (Hz), k is wave number, T is wave period, L is wave length, and where we assume, as stated above, that ka = O(0).
The wave period T is the time it takes two successive wave crests or troughs to pass a fixed point. The wave length L is the distance between two successive wave crests or troughs at a fixed time.
Dispersion Relation Wave frequency ω is related to wave number k by the dispersion relation (Lamb, 1945 §228):
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where d is the water depth and g is the acceleration of gravity. Two approximations are especially useful.
1.Deep-water approximation is valid if the water depth d is much greater than the wave length L. In this case, d L, kd 1, and tanh(kd) = 1.
2.Shallow-water approximation is valid if the water depth is much less than a wavelength. In this case, d L, kd 1, and tanh(kd) = kd.
For these two limits of water depth compared with wavelength the dispersion relation reduces to:
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The stated limits for d/L give a dispersion relation accurate within 10%. Because many wave properties can be measured with accuracies of 5–10%, the approximations are useful for calculating wave properties. Later we will learn to calculate wave properties as the waves propagate from deep to shallow water.
Phase Velocity The phase velocity c is the speed at which a particular phase of the wave propagates, for example, the speed of propagation of the wave crest. In one wave period T the crest advances one wave length L and the phase speed is c = L/T = ω/k. Thus, the definition of phase speed is:
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16.1. LINEAR THEORY OF OCEAN SURFACE WAVES |
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The direction of propagation is perpendicular to the wave crest and toward the positive x direction.
The deepand shallow-water approximations for the dispersion relation give:
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The approximations are accurate to about 5% for limits stated in (16.4, 16.5). In deep water, the phase speed depends on wave length or wave frequency.
Longer waves travel faster. Thus, deep-water waves are said to be dispersive. In shallow water, the phase speed is independent of the wave; it depends only on the depth of the water. Shallow-water waves are non-dispersive.
Group Velocity The concept of group velocity cg is fundamental for understanding the propagation of linear and nonlinear waves. First, it is the velocity at which a group of waves travels across the ocean. More importantly, it is also the propagation velocity of wave energy. Whitham (1974, §1.3 and §11.6) gives a clear derivation of the concept and the fundamental equation (16.9).
The definition of group velocity in two dimensions is:
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Using the approximations for the dispersion relation:
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For ocean-surface waves, the direction of propagation is perpendicular to the wave crests in the positive x direction. In the more general case of other types of waves, such as Kelvin and Rossby waves that we met in §14.2, the group velocity is not necessarily in the direction perpendicular to wave crests.
Notice that a group of deep-water waves moves at half the phase speed of the waves making up the group. How can this happen? If we could watch closely a group of waves crossing the sea, we would see waves crests appear at the back of the wave train, move through the train, and disappear at the leading edge of the group. Each wave crest moves at twice the speed of the group.
Do real ocean waves move in groups governed by the dispersion relation? Yes. Walter Munk and colleagues (1963) in a remarkable series of experiments in the 1960s showed that ocean waves propagating over great distances are dispersive, and that the dispersion could be used to track storms. They recorded waves for many days using an array of three pressure gauges just o shore of San Clemente Island, 60 miles due west of San Diego, California. Wave spectra were
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September 1959
Figure 16.1 Contours of wave energy on a frequency-time plot calculated from spectra of waves measured by pressure gauges o shore of southern California. The ridges of high wave energy show the arrival of dispersed wave trains from distant storms. The slope of the ridge is inversely proportional to distance to the storm. is distance in degrees, θ is direction of arrival of waves at California. After Munk et al. (1963).
calculated for each day’s data. (The concept of a spectra is discussed below.) From the spectra, the amplitudes and frequencies of the low-frequency waves and the propagation direction of the waves were calculated. Finally, they plotted contours of wave energy on a frequency-time diagram (figure 16.1).
To understand the figure, consider a distant storm that produces waves of many frequencies. The lowest-frequency waves (smallest ω) travel the fastest (16.11), and they arrive before other, higher-frequency waves. The further away the storm, the longer the delay between arrivals of waves of di erent frequencies. The ridges of high wave energy seen in the figure are produced by individual storms. The slope of the ridge gives the distance to the storm in degrees along a great circle; and the phase information from the array gives the angle to the storm θ. The two angles give the storm’s location relative to San Clemente. Thus waves arriving from 15 to 18 September produce a ridge indicating the storm was 115◦ away at an angle of 205◦ which is south of new Zealand near Antarctica.
The locations of the storms producing the waves recorded from June through October 1959 were compared with the location of storms plotted on weather maps and in most cases the two agreed well.
16.1. LINEAR THEORY OF OCEAN SURFACE WAVES |
277 |
Wave Energy Wave energy E in Joules per square meter is related to the variance of sea-surface displacement ζ by:
E = ρwg ζ2 |
(16.12) |
where ρw is water density, g is gravity, and the brackets denote a time or space
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Figure 16.2 A short record of wave amplitude measured by a wave buoy in the north Atlantic.
Significant Wave Height What do we mean by wave height? If we look at a wind-driven sea, we see waves of various heights. Some are much larger than most, others are much smaller (figure 16.2). A practical definition that is often used is the height of the highest 1/3 of the waves, H1/3. The height is computed as follows: measure wave height for a few minutes, pick out say 120 wave crests and record their heights. Pick the 40 largest waves and calculate the average height of the 40 values. This is H1/3 for the wave record.
The concept of significant wave height was developed during the World War II as part of a project to forecast ocean wave heights and periods. Wiegel (1964: p. 198) reports that work at the Scripps Institution of Oceanography showed
. . . wave height estimated by observers corresponds to the average of the highest 20 to 40 per cent of waves. . . Originally, the term significant wave height was attached to the average of these observations, the highest 30 per cent of the waves, but has evolved to become the average of the highest one-third of the waves, (designated HS or H1/3)
More recently, significant wave height is calculated from measured wave displacement. If the sea contains a narrow range of wave frequencies, H1/3 is related to the standard deviation of sea-surface displacement (nas , 1963: 22; Ho man and Karst, 1975)
where ζ |
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height from wave measurements.