Radiation stress

Breaking waves on beaches induce variations in radiation stress, driving longshore currents. The resulting longshore sediment transport shapes the beaches, and may result in beach erosion or accretion.

In fluid dynamics, the radiation stress is the depth-integrated – and thereafter phase-averaged – excess momentum flux caused by the presence of the surface gravity waves, which is exerted on the mean flow. The radiation stresses behave as a second-order tensor.

The radiation stress tensor describes the additional forcing due to the presence of the waves, which changes the mean depth-integrated horizontal momentum in the fluid layer. As a result, varying radiation stresses induce changes in the mean surface elevation (wave setup) and the mean flow (wave-induced currents).

For the mean energy density in the oscillatory part of the fluid motion, the radiation stress tensor is important for its dynamics, in case of an inhomogeneous mean-flow field.

The radiation stress tensor, as well as several of its implications on the physics of surface gravity waves and mean flows, were formulated in a series of papers by Longuet-Higgins and Stewart in 1960–1964.

Radiation stress derives its name from the analogous effect of radiation pressure for electromagnetic radiation.

Physical significance

The radiation stress – mean excess momentum-flux due to the presence of the waves – plays an important role in the explanation and modeling of various coastal processes:^[1][2][3]

• Wave setup and setdown – the radiation stress consists in part of a radiation pressure, exerted at the free surface elevation of the mean flow. If the radiation stress varies spatially, as it does in the surf zone where the wave height reduces by wave breaking, this results in changes of the mean surface elevation called wave setup (in case of an increased level) and setdown (for a decreased water level);
• Wave-driven current, especially a longshore current in the surf zone – for oblique incidence of waves on a beach, the reduction in wave height inside the surf zone (by breaking) introduces a variation of the shear-stress component S_xy of the radiation stress over the width of the surf zone. This provides the forcing of a wave-driven longshore current, which is of importance for sediment transport (longshore drift) and the resulting coastal morphology;
• Bound long waves or forced long waves – for wave groups the radiation stress varies along the group. As a result, a non-linear long wave propagates together with the group, at the group velocity of the modulated short waves within the group. While, according to the dispersion relation, a long wave of this length should propagate at its own – higher – phase velocity. The amplitude of this bound long wave varies with the square of the wave height, and is only significant in shallow water;
• Wave–current interaction – in varying mean-flow fields, the energy exchanges between the waves and the mean flow, as well as the mean-flow forcing, can be modeled by means of the radiation stress.

Definitions and values derived from linear wave theory

One-dimensional wave propagation

For uni-directional wave propagation – say in the x-coordinate direction – the component of the radiation stress tensor of dynamical importance is S_xx. It is defined as:^[4]

${\displaystyle S_{xx}={\overline {\int _{-h}^{\eta }\left(p+\rho {\tilde {u}}^{2}\right)\;{\text{d}}z}}-{\frac {1}{2}}\rho g\left(h+{\overline {\eta }}\right)^{2},}$

where p(x,z,t) is the fluid pressure, ${\displaystyle {\tilde {u}}(x,z,t)}$ is the horizontal x-component of the oscillatory part of the flow velocity vector, z is the vertical coordinate, t is time, z = −h(x) is the bed elevation of the fluid layer, and z = η(x,t) is the surface elevation. Further ρ is the fluid density and g is the acceleration by gravity, while an overbar denotes phase averaging. The last term on the right-hand side, ½ρg(h+η)², is the integral of the hydrostatic pressure over the still-water depth.

To lowest (second) order, the radiation stress S_xx for traveling periodic waves can be determined from the properties of surface gravity waves according to Airy wave theory:^[5][6]

${\displaystyle S_{xx}=\left(2{\frac {c_{g}}{c_{p}}}-{\frac {1}{2}}\right)E,}$

where c_p is the phase speed and c_g is the group speed of the waves. Further E is the mean depth-integrated wave energy density (the sum of the kinetic and potential energy) per unit of horizontal area. From the results of Airy wave theory, to second order, the mean energy density E equals:^[7]

${\displaystyle E={\frac {1}{2}}\rho ga^{2}={\frac {1}{8}}\rho gH^{2},}$

with a the wave amplitude and H = 2a the wave height. Note this equation is for periodic waves: in random waves the root-mean-square wave height H_rms should be used with H_rms = H_m0 / √2, where H_m0 is the significant wave height. Then E =  ¹⁄₁₆ρgH_m0².

Two-dimensional wave propagation

For wave propagation in two horizontal dimensions the radiation stress $\mathbf {S}$ is a second-order tensor^[8][9] with components:

${\displaystyle \mathbf {S} ={\begin{pmatrix}S_{xx}&S_{xy}\\S_{yx}&S_{yy}\end{pmatrix}}.}$

With, in a Cartesian coordinate system (x,y,z):^[4]

${\displaystyle {\begin{aligned}S_{xx}&={\overline {\int _{-h}^{\eta }\left(p+\rho {\tilde {u}}^{2}\right)\;{\text{d}}z}}-{\frac {1}{2}}\rho g\left(h+{\overline {\eta }}\right)^{2},\\S_{xy}&={\overline {\int _{-h}^{\eta }\left(\rho {\tilde {u}}{\tilde {v}}\right)\;{\text{d}}z}}=S_{yx},\\S_{yy}&={\overline {\int _{-h}^{\eta }\left(p+\rho {\tilde {v}}^{2}\right)\;{\text{d}}z}}-{\frac {1}{2}}\rho g\left(h+{\overline {\eta }}\right)^{2},\end{aligned}}}$

where ${\tilde {u}}$ and ${\displaystyle {\tilde {v}}}$ are the horizontal x- and y-components of the oscillatory part ${\displaystyle {\tilde {u}}(x,y,z,t)}$ of the flow velocity vector.

To second order – in wave amplitude a – the components of the radiation stress tensor for progressive periodic waves are:^[5]

${\displaystyle {\begin{aligned}S_{xx}&=\left[{\frac {k_{x}^{2}}{k^{2}}}{\frac {c_{g}}{c_{p}}}+\left({\frac {c_{g}}{c_{p}}}-{\frac {1}{2}}\right)\right]E,\\S_{xy}&=\left({\frac {k_{x}k_{y}}{k^{2}}}{\frac {c_{g}}{c_{p}}}\right)E=S_{yx},\quad {\text{and}}\\S_{yy}&=\left[{\frac {k_{y}^{2}}{k^{2}}}{\frac {c_{g}}{c_{p}}}+\left({\frac {c_{g}}{c_{p}}}-{\frac {1}{2}}\right)\right]E,\end{aligned}}}$

where k_x and k_y are the x- and y-components of the wavenumber vector k, with length k = |k| = √k_x²+k_y² and the vector k perpendicular to the wave crests. The phase and group speeds, c_p and c_g respectively, are the lengths of the phase and group velocity vectors: c_p = |c_p| and c_g = |c_g|.

Dynamical significance

The radiation stress tensor is an important quantity in the description of the phase-averaged dynamical interaction between waves and mean flows. Here, the depth-integrated dynamical conservation equations are given, but – in order to model three-dimensional mean flows forced, or interacting with, surface waves – a three-dimensional description of the radiation stress over the fluid layer is needed.^[10]

Mass transport velocity

Propagating waves induce a – relatively small – mean mass transport in the wave propagation direction, also called the wave (pseudo) momentum.^[11] To lowest order, the wave momentum M_w is, per unit of horizontal area:^[12]

${\displaystyle {\boldsymbol {M}}_{w}={\frac {\boldsymbol {k}}{k}}{\frac {E}{c_{p}}},}$

which is exact for progressive waves of permanent form in irrotational flow. Above, c_p is the phase speed relative to the mean flow:

${\displaystyle c_{p}={\frac {\sigma }{k}}\qquad {\text{with}}\qquad \sigma =\omega -{\boldsymbol {k}}\cdot {\overline {\boldsymbol {v}}},}$

with σ the intrinsic angular frequency, as seen by an observer moving with the mean horizontal flow-velocity v while ω is the apparent angular frequency of an observer at rest (with respect to 'Earth'). The difference k⋅v is the Doppler shift.^[13]

The mean horizontal momentum M, also per unit of horizontal area, is the mean value of the integral of momentum over depth:

${\displaystyle {\boldsymbol {M}}={\overline {\int _{-h}^{\eta }\rho \,{\boldsymbol {v}}\;{\text{d}}z}}=\rho \,\left(h+{\overline {\eta }}\right){\overline {\boldsymbol {v}}}+{\boldsymbol {M}}_{w},}$

with v(x,y,z,t) the total flow velocity at any point below the free surface z = η(x,y,t). The mean horizontal momentum M is also the mean of the depth-integrated horizontal mass flux, and consists of two contributions: one by the mean current and the other (M_w) is due to the waves.

Now the mass transport velocity u is defined as:^[14][15]

${\displaystyle {\overline {\boldsymbol {u}}}={\frac {\boldsymbol {M}}{\rho \,\left(h+{\overline {\eta }}\right)}}={\overline {\boldsymbol {v}}}+{\frac {{\boldsymbol {M}}_{w}}{\rho \,\left(h+{\overline {\eta }}\right)}}.}$

Observe that first the depth-integrated horizontal momentum is averaged, before the division by the mean water depth (h+η) is made.

Mass and momentum conservation

Vector notation

The equation of mean mass conservation is, in vector notation:^[14]

${\displaystyle {\frac {\partial }{\partial t}}\left[\rho \left(h+{\overline {\eta }}\right)\right]+\nabla \cdot \left[\rho \left(h+{\overline {\eta }}\right){\overline {\boldsymbol {u}}}\right]=0,}$

with u including the contribution of the wave momentum M_w.

The equation for the conservation of horizontal mean momentum is:^[14]

${\displaystyle {\frac {\partial }{\partial t}}\left[\rho \left(h+{\overline {\eta }}\right){\overline {\boldsymbol {u}}}\right]+\nabla \cdot \left[\rho \left(h+{\overline {\eta }}\right){\overline {\boldsymbol {u}}}\otimes {\overline {\boldsymbol {u}}}+\mathbf {S} +{\frac {1}{2}}\rho g(h+{\overline {\eta }})^{2}\,\mathbf {I} \right]=\rho g\left(h+{\overline {\eta }}\right)\nabla h+{\boldsymbol {\tau }}_{w}-{\boldsymbol {\tau }}_{b},}$

where u ⊗ u denotes the tensor product of u with itself, and τ_w is the mean wind shear stress at the free surface, while τ_b is the bed shear stress. Further I is the identity tensor, with components given by the Kronecker delta δ_ij. Note that the right hand side of the momentum equation provides the non-conservative contributions of the bed slope ∇h,^[16] as well the forcing by the wind and the bed friction.

In terms of the horizontal momentum M the above equations become:^[14]

${\displaystyle {\begin{aligned}&{\frac {\partial }{\partial t}}\left[\rho \left(h+{\overline {\eta }}\right)\right]+\nabla \cdot {\boldsymbol {M}}=0,\\&{\frac {\partial {\boldsymbol {M}}}{\partial t}}+\nabla \cdot \left[{\overline {\boldsymbol {u}}}\otimes {\boldsymbol {M}}+\mathbf {S} +{\frac {1}{2}}\rho g(h+{\overline {\eta }})^{2}\,\mathbf {I} \right]=\rho g\left(h+{\overline {\eta }}\right)\nabla h+{\boldsymbol {\tau }}_{w}-{\boldsymbol {\tau }}_{b}.\end{aligned}}}$

Component form in Cartesian coordinates

In a Cartesian coordinate system, the mass conservation equation becomes:

${\displaystyle {\frac {\partial }{\partial t}}\left[\rho \left(h+{\overline {\eta }}\right)\right]+{\frac {\partial }{\partial x}}\left[\rho \left(h+{\overline {\eta }}\right){\overline {u}}_{x}\right]+{\frac {\partial }{\partial y}}\left[\rho \left(h+{\overline {\eta }}\right){\overline {u}}_{y}\right]=0,}$

with u_x and u_y respectively the x and y components of the mass transport velocity u.

The horizontal momentum equations are:

${\displaystyle {\begin{aligned}{\frac {\partial }{\partial t}}\left[\rho \left(h+{\overline {\eta }}\right){\overline {u}}_{x}\right]&+{\frac {\partial }{\partial x}}\left[\rho \left(h+{\overline {\eta }}\right){\overline {u}}_{x}{\overline {u}}_{x}+S_{xx}+{\frac {1}{2}}\rho g(h+{\overline {\eta }})^{2}\right]+{\frac {\partial }{\partial y}}\left[\rho \left(h+{\overline {\eta }}\right){\overline {u}}_{x}{\overline {u}}_{y}+S_{xy}\right]\\&=\rho g\left(h+{\overline {\eta }}\right){\frac {\partial }{\partial x}}h+\tau _{w,x}-\tau _{b,x},\\{\frac {\partial }{\partial t}}\left[\rho \left(h+{\overline {\eta }}\right){\overline {u}}_{y}\right]&+{\frac {\partial }{\partial x}}\left[\rho \left(h+{\overline {\eta }}\right){\overline {u}}_{y}{\overline {u}}_{x}+S_{yx}\right]+{\frac {\partial }{\partial y}}\left[\rho \left(h+{\overline {\eta }}\right){\overline {u}}_{y}{\overline {u}}_{y}+S_{yy}+{\frac {1}{2}}\rho g(h+{\overline {\eta }})^{2}\right]\\&=\rho g\left(h+{\overline {\eta }}\right){\frac {\partial }{\partial y}}h+\tau _{w,y}-\tau _{b,y}.\end{aligned}}}$

Energy conservation

For an inviscid flow the mean mechanical energy of the total flow – that is the sum of the energy of the mean flow and the fluctuating motion – is conserved.^[17] However, the mean energy of the fluctuating motion itself is not conserved, nor is the energy of the mean flow. The mean energy E of the fluctuating motion (the sum of the kinetic and potential energies satisfies:^[18]

${\displaystyle {\frac {\partial E}{\partial t}}+\nabla \cdot \left[\left({\overline {\boldsymbol {u}}}+{\boldsymbol {c}}_{g}\right)E\right]+\mathbf {S} :\left(\nabla \otimes {\overline {\boldsymbol {u}}}\right)={\boldsymbol {\tau }}_{w}\cdot {\overline {\boldsymbol {u}}}-{\boldsymbol {\tau }}_{b}\cdot {\overline {\boldsymbol {u}}}-\varepsilon ,}$

where ":" denotes the double-dot product, and ε denotes the dissipation of mean mechanical energy (for instance by wave breaking). The term ${\displaystyle \mathbf {S} :\left(\nabla \otimes {\overline {\boldsymbol {u}}}\right)}$ is the exchange of energy with the mean motion, due to wave–current interaction. The mean horizontal wave-energy transport (u + c_g) E consists of two contributions:

• u E : the transport of wave energy by the mean flow, and
• c_g E : the mean energy transport by the waves themselves, with the group velocity c_g as the wave-energy transport velocity.

In a Cartesian coordinate system, the above equation for the mean energy E of the flow fluctuations becomes:

${\displaystyle {\begin{aligned}{\frac {\partial E}{\partial t}}&+{\frac {\partial }{\partial x}}\left[\left({\overline {u}}_{x}+c_{g,x}\right)E\right]+{\frac {\partial }{\partial y}}\left[\left({\overline {u}}_{y}+c_{g,y}\right)E\right]\\&+S_{xx}{\frac {\partial {\overline {u}}_{x}}{\partial x}}+S_{xy}\left({\frac {\partial {\overline {u}}_{y}}{\partial x}}+{\frac {\partial {\overline {u}}_{x}}{\partial y}}\right)+S_{yy}{\frac {\partial {\overline {u}}_{y}}{\partial y}}\\&=\left(\tau _{w,x}-\tau _{b,x}\right){\overline {u}}_{x}+\left(\tau _{w,y}-\tau _{b,y}\right){\overline {u}}_{y}-\varepsilon .\end{aligned}}}$

So the radiation stress changes the wave energy E only in case of a spatial-inhomogeneous current field (u_x,u_y).

Notes

1. Longuet-Higgins & Stewart (1964,1962).
2. Phillips (1977), pp. 70–81.
3. Battjes, J. A. (1974). Computation of set-up, longshore currents, run-up and overtopping due to wind-generated waves (Thesis). Delft University of Technology. Retrieved 2010-11-25.
4. Mei (2003), p. 457.
5. Mei (2003), p. 97.
6. Phillips (1977), p. 68.
7. Phillips (1977), p. 39.
8. Longuet-Higgins & Stewart (1961).
9. Dean, R.G.; Walton, T.L. (2009), "Wave setup", in Young C. Kim, Handbook of Coastal and Ocean Engineering, World Scientific, pp. 1–23, ISBN 981-281-929-0.
10. Walstra, D. J. R.; Roelvink, J. A.; Groeneweg, J. (2000), "Calculation of wave-driven currents in a 3D mean flow model", Proceedings of the 27th International Conference on Coastal Engineering, Sydney: ASCE, pp. 1050–1063, doi:10.1061/40549(276)81
11. Mcintyre, M. E. (1981), "On the 'wave momentum' myth", Journal of Fluid Mechanics, 106: 331–347, Bibcode:1981JFM...106..331M, doi:10.1017/S0022112081001626
12. Phillips (1977), p. 40.
13. Phillips (1977), pp. 23–24.
14. Phillips (1977), pp. 61–63.
15. Mei (2003), p. 453.
16. By Noether's theorem, an inhomogeneous medium – in this case a non-horizontal bed, h(x,y) not a constant – results in non-conservation of the depth-integrated horizontal momentum.
17. Phillips (1977), pp. 63–65.
18. Phillips (1977), pp. 65–66.

References

Primary sources

• Longuet-Higgins, M. S.; Stewart, R. W. (1960), "Changes in the form of short gravity waves on long waves and tidal currents", Journal of Fluid Mechanics, 8 (4): 565–583, Bibcode:1960JFM.....8..565L, doi:10.1017/S0022112060000803
• Longuet-Higgins, M. S.; Stewart, R. W. (1961), "The changes in amplitude of short gravity waves on steady non-uniform currents", Journal of Fluid Mechanics, 10 (4): 529–549, Bibcode:1961JFM....10..529L, doi:10.1017/S0022112061000342
• Longuet-Higgins, M. S.; Stewart, R. W. (1962), "Radiation stress and mass transport in gravity waves, with application to 'surf beats'", Journal of Fluid Mechanics, 13 (4): 481–504, Bibcode:1962JFM....13..481L, doi:10.1017/S0022112062000877
• Longuet-Higgins, M. S.; Stewart, R. W. (1964), "Radiation stresses in water waves; a physical discussion, with applications", Deep-Sea Research, 11 (4): 529–562, Bibcode:1964DSROA..11..529L, doi:10.1016/0011-7471(64)90001-4

Further reading

• Mei, Chiang C. (2003), The applied dynamics of ocean surface waves, Advanced series on ocean engineering, 1, World Scientific, ISBN 9971-5-0789-7
• Phillips, O. M. (1977), The dynamics of the upper ocean (2nd ed.), Cambridge University Press, ISBN 0-521-29801-6