Seawater intrusion and mixing in estuaries

Introduction

Estuaries are generally defined as semi-enclosed transition zones between river and sea. The intrusion of seawater in estuaries is mainly due to tides and buoyancy (related to the density difference between seawater and river water). Seawater intrusion in estuaries is an important phenomenon to man and nature: it limits fresh water availability for human and agricultural use and it determines the type of habitats and species that can develop in an estuarine environment. Besides, density driven currents and salinity play a role in estuarine turbidity and sedimentation processes.

We describe in this article the physical processes involved in seawater intrusion and mixing in estuaries and explain some simple methods for deriving quantitative estimates.

Seawater intrusion mechanisms

The fresh water discharge $Q_R$ and the salt flux $Q_S$ in an estuary are given by

$Q_R=-\lt A\overline{\overline{u}}\gt \equiv - \frac{1}{T}\int_0^T \int_{-B/2}^{B/2} \int_0^D u(x,y,z,t) dzdy dt$,

$Q_S=\lt A\overline{\overline{us}}\gt \equiv \frac{1}{T} \int_0^T \int_{-B/2}^{B/2} \int_0^D u(x,y,z,t)s(x,y,z,t) dzdydt$,

where $A(x,t)$ is the estuarine cross-section, $D(x,y,t)$ the instantaneous local water depth and $B(x,t)$ the estuarine width, $u$ the longitudinal current velocity and $s$ the salinity. The brackets stand for averaging over the tidal period $T$ (assuming a cyclic tide) and the overbars stand for averaging over the depth and the width. The coordinate $x$ follows the upstream positive longitudinal direction (along the thalweg), the coordinate $y$ the lateral direction and the coordinate $z$ the upward positive vertical direction.

We call $s_0 \equiv \lt \overline{\overline{s}}\gt$ the salinity averaged over the estuarine cross-section $A$ and the tidal period. We may then decompose

$Q_S=Q_{disp}- Q_R s_0, \quad Q_{disp} = \lt A\overline{\overline{u(s-s_0)}}\gt$ .

This decomposition singles out the fresh water discharge (the term $Q_R s_0$ ) as a mechanism for flushing seawater out of the estuary, while the term $Q_{disp}$ represents the sum of all processes contributing to seawater intrusion. These processes are:

• Horizontal circulations in the estuary (mainly the net relative displacement of water masses circulating between ebb and flood channels and the net relative displacements due to geometry-induced eddies, followed by lateral mixing of these water masses);
• Horizontal tidal straining (lateral mixing between water masses which are advected at different speeds, due to lateral gradients in the longitudinal velocity);
• Vertical circulation in the estuary, also called estuarine circulation (seawater intrusion induced by the density-driven net displacement of near-surface water relative to near-bottom water, followed by mixing over the vertical);
• Vertical tidal straining (vertical mixing between water masses advected at different speeds due to vertical gradients in the longitudinal velocity);
• Lateral mixing of water masses captured in "dead zones" with the main flow;
• Chaotic dispersion, related to the chaotic character of particle trajectories when travelling through a complex field of tide-generated eddies;
• Tidal pumping at the inlet (partial replacement of the ebb tidal prism with ‘new’ seawater flowing in from the nearshore zone during flood).

Random walk

Water parcels move some time forth and back in an estuary before they are evacuated offshore. We call $T_x$ the flushing time of water parcels in the estuary (the average residence time fluid parcels entering the estuary at the upstream boundary). During this time, water parcels also move in lateral and vertical directions, due to flow circulations and turbulent eddies. The time scale for vertical mixing, $T_z$, and the time scale for lateral mixing, $T_y$, are related to the vertical and lateral turbulent diffusion coefficients, $K_z$ and $K_y$, by the relationships $T_z=D^2/K_z$ and $T_y=B^2/K_y$, respectively. If the vertical and lateral mixing time scales are both much smaller than the flushing time $T_x$ of water parcels in the estuary, the longitudinal path of a water parcel follows a random walk. The longitudinal displacements $X_n$ in successive time intervals $T_n$ are uncorrelated, by choosing $T_n=T$ equal to an integer number of cyclic tidal periods such that $T_n \gt T_A$. The cross-sectional mixing time $T_A = T_y$, assuming that lateral mixing takes more time than vertical mixing, as is the case for most estuaries. In estuaries satisfying the condition $T_x \gt T_A$, salt transport by seawater intrusion processes, $Q_{disp}$, can be represented by a gradient-type transport formula,

$Q_{disp} = \lt A\overline{\overline{u(s-s_0)}}\gt = - A_0 K_x \frac{ds_0}{dx}$,

where $K_x$ is the longitudinal dispersion coefficient and $A_0=\lt A\gt$. The dispersion coefficient $K_x$ has the important property that it does not depend explicitly on the salinity distribution in the estuary , but only on the flow characteristics during the tidal cycle (which may be influenced by the salinity distribution, by the way). According to random walk theory ,

$K_x =\frac{\overline{X^2}}{2 T_n}$,

where $\overline{X^2}$ is the average of the squared successive random displacements,

$\overline{X^2}= \frac{1}{N} \sum_{n=1}^N X_n^2 , \quad N\gt \gt 1$.

The magnitude of the random displacements depends on the location $x$ in the estuary; the longitudinal dispersion coefficient $K_x$ is thus a function of $x$. This is illustrated in figure 1 for the Eastern Scheldt and Ems-Dollard estuaries.

Figure 1: Longitudinal dispersion coefficients at different locations in the Eastern Scheldt and Ems-Dollard estuaries.

Analytical expressions for the longitudinal dispersion coefficient

Under certain simplifying conditions it is possible to derive analytical expressions for the longitudinal dispersion coefficient. These assumptions are:

• the estuarine geometry does not vary strongly in $x$-direction over distances comparable to or larger than the tidal excursion;
• the cross-section of the estuarine main channel has approximately a rectangular shape.

We also have the condition $T_A\lt \lt T_x$. In the following we consider different seawater intrusion processes under these conditions and present an approximate analytical expression for the dispersive transport produced by each process.

Dispersion by residual circulation

Fist we consider seawater intrusion caused by estuarine circulation: the up-estuary near-bottom flow caused by the higher density of seawater relative to estuarine water. The estuarine circulation is represented by the velocity component

$u_0 (z)= \lt u\gt -\lt \overline{u}^z\gt$,

where the brackets $\lt u\gt$ stand for averaging over the tidal period (in fact, the averaging is done in a frame moving with the cross-sectional mean velocity), and $\overline{u}^z$ for averaging over the vertical. The longitudinal dispersive transport can be estimated by a procedure outlined by G.I.Taylor . The result is

$K_x = f_0^{(z)} T_z \overline{(u_0)^2}^z$ ,

with $f_0^{(z)} \approx 0.1$ .

For the dispersion coefficient related to lateral horizontal residual circulation a similar formula can be derived, replacing in the expression for $K_x$ everywhere $z$ by $y$.

Estuarine circulation is an important seawater intrusion mechanism in estuaries with a single deep (dredged) channel and small to moderate tide. Lateral circulations are important in wide natural estuaries with a complex geometry (meandering main channel , secondary channels, channel bars and tidal flats) and strong tides. The dominance of lateral circulations for seawater intrusion relative to vertical circulations appears in the analytical expression of $K_x$ through the much larger lateral mixing time $T_y$ compared to the vertical mixing time $T_z$. The presence of distinct ebb and flood channels is a major cause of lateral circulation in wide estuaries, see for example figure 2. However, density gradients related to seawater intrusion also produce lateral circulations, which contribute often even more to longitudinal dispersion than the vertical density-induced circulation .

Figure 2: Circulation cells in the Western Scheldt related to ebb- and flood-dominated channels.

Dispersion by tidal straining

If residual circulations are weak, dispersion is mainly caused by tidal straining, the relative displacement of water masses due to vertical and horizontal gradients in the tidal current velocity (In river flow, the usual term is shear dispersion). Seawater intrusion is primarily caused by vertical tidal velocity gradients in narrow deep estuaries, whereas lateral tidal velocity gradients are important in wide estuaries. We present formulas for vertical tidal straining; the expressions for lateral tidal straining are similar. The process of longitudinal dispersion through tidal straining is explained in figure 3.

The velocity component $u_1 (z,t)$ responsible for vertical tidal straining is

$u_1= u-\overline{u}^z$,

where $\overline{u}^z$ is the depth-average current velocity. By a procedure outlined by Holley, Harleman and Fischer , the following first-order estimate of the longitudinal dispersion coefficient is obtained:

$K_x \approx f_1^{(z)} \frac{T_z \lt \overline{ u_1^2}^z\gt }{1+( f_2^{(z)} T_z / T)^2 }$ ,

The coefficients $f_1^{(z)}, f_2^{(z)}$ depend on the velocity profile; order-of-magnitude estimates are $f_1^{(z)} \sim 0.1-0.2$ and $f_2^{(z)} \sim 0.5-1$.

Longitudinal dispersion produced by lateral tidal straining can be expressed by a formula similar as for vertical tidal straining. Dispersion by tidal straining is largest if the time scale for vertical or lateral mixing is on the order of $T/ f_2$. The time scale for vertical mixing is generally smaller than the tidal period and the time scale for lateral mixing is generally larger.

It should be realised that longitudinal dispersion is not simply the sum of transport processes related to circulation and straining. Circulation causes not only a net relative displacements of water masses in the estuary, but it also influences tidal straining.

Figure 3: Schematic representation of longitudinal dispersion by a vertical gradient in the tidal current $u(z,t)$. A patch of dye is introduced homogeneously over the vertical when flood starts at low water slack tide ($t=0$). The patch of dye is stretched by the tidal current while being mixed vertically by turbulent diffusion. The time scale for vertical mixing $T_z$ is assumed equal to the tidal period $T$. The tidal cycle is represented by four discrete time steps, illustrating the longitudinal spread of the patch of dye.

Dispersion by “dead zones”

The formula for lateral tidal straining includes the influence of "dead zones", if they are considered part of the estuarine cross-section and if they are distributed regularly along the estuary. Dead zones are areas along the main estuarine channel where water is not transported in longitudinal direction, for instance, tidal flats or lateral creeks. The longitudinal dispersion coefficient is given by an expression of the type

$K_x=f_{dz}rU^2T$,

where $U$ is the maximum tidal velociy and $r$ is the storage cross-section of the dead zones relative to the channel cross-section. The coefficient $f_{dz}$ depends on the mixing rate within the dead zone; in case of complete mixing during the tidal cycle $f_{dz}=1/12 \pi^2$, assuming that filling of the dead zones starts at low water (LW) . Even without any mixing, storage areas along the channel contribute to longitudinal dispersion, because of a non-zero phase shift $\phi$ that generally exists between horizontal and vertical tidal motion (between $u(t)$ and $d\eta /dt$, respectively, where $\eta(t)$ is the tidal level). The process is illustrated in figure 4. The coefficient $f_{dz}$ is now given by $f_{dz} = \sin \phi / (8 \pi)$.

A usual order of magnitude for the phase shift $\phi$ is 30-60 minutes times $2 \pi / T$, yielding $f_{dz} \approx 0.015$.

Figure 4: Dispersion by an intertidal "dead zone" along a tidal channel. The figure illustrates the case of a phase shift $\phi$ between the vertical tidal motion ($d\eta /dt$) and the horizontal tidal motion ($u(t)$). The paths of two fluid parcels are shown, both starting from the same location at low water ($t=0$). The first parcel (the square) follows the tidal motion in the channel, while the second parcel (the circle) is stored in the dead zone around high water in a time interval $[T/2-t,T/2+t]$. Without mixing, the latter parcel returns to the channel before the former parcel has arrived at the same location. The result is longitudinal dispersion: a net relative displacement of the two fluid during the tidal cycle.

Time scales for vertical and lateral mixing

A difficulty for practical use of the previous expressions for longitudinal dispersion, results from the uncertainty related to estimating the vertical and (especially) lateral mixing times, $T_z=D^2/K_z$ and $T_y=B^2/K_y$. In case of a logarithmic velocity profile, the vertical diffusion coefficient is given by $K_z=0.4z(1-z/D)u_*$, where $u_*\approx 0.05 U$ is the friction velocity and $U$ the flow velocity. This yields a longitudinal dispersion coefficient $K_x \approx 6u_*D$, for steady flow . However, in case of buoyant flows, vertical diffusion can be much slower (smaller $K_z$), leading to stronger longitudinal dispersion. Lateral diffusion depends strongly on the geometry of the estuary. The lateral diffusion coefficient is generally expressed as $K_y \approx \alpha u_* D$. An empirical estimate for moderately meandering channels is $\alpha \approx 0.6$ and a model estimate is $\alpha \approx 150 (B/ R)^2$ , where $R$ is a characteristic channel bend radius.

Chaotic dispersion

Dye experiments show that dispersion in wide estuaries with complex geometry generally proceeds in an irregular way, by advection through a field of geometry-induced tidal eddies. The result is very different from diffusion by a cascade of turbulent eddies of progressively decreasing size. Parts of the dye can be trapped within gyres with almost no diffusion, while other dye patches can be highly stretched in the flow direction. Strong stretching occurs in particular in the interface zones between tidal eddies. Zimmerman described the dispersion process in such systems as the result of Lagrangian chaos produced by the tidal whirlpool . Fluid parcels can be dispersed over the entire length of the estuary before lateral mixing has taken place. In this case, the random walk description of tidal dispersion is no longer valid. Zimmerman showed that longitudinal dispersion can still be described as a random process, even if turbulent mixing is completely absent. He called this random process “deterministic chaos” . In his model, fluid parcels are dispersed by moving along chaotic orbits through a lattice of tidal eddies. Most dispersion is generated by eddies with a size $\lambda$ comparable to the tidal excursion length $L$ . This suggests that the longitudinal dispersion coefficient for chaotic mixing should be proportional to

$K_x \propto UL$.

The size of the eddies also depends on the basin with $B$. A field study in Willapa Bay (US Pacific coast) suggests that chaotic dispersion could be described by a dispersion coefficient $K_x = 0.06 UB$ . If the lateral mixing time $T_y$ is comparable to or larger than the flushing time $T_x$, the representation of the dispersive transport $Q_{disp}$ by the product of a dispersion coefficient and the local salinity gradient is no longer valid.

Tidal pumping

The term tidal pumping is sometimes used for dispersion produced by circulations between ebb and flood channels in an estuary. Here we reserve the term tidal pumping to the phenomenon of estuarine outflow as an ebb tidal jet, see figure 5. The inflowing flood water therefore contains ‘new’ seawater from the sides of the ebb channel and from the near-bottom zone.

Figure 5: Schematic representation of seawater intrusing by tidal pumping.

We call $\alpha$ the rate of renewal of outflowing estuarine with seawater. For small fresh water discharge, the corresponding dispersion coefficient at the estuarine mouth is given by

$K_{mouth} = \alpha \frac{L^2 }{2T}$ ,

where $L=\int_{flood} u(t)dt$ is the tidal excursion at the estuarine mouth. Savenije derived the following semi-empirical expression for $\alpha$:

$\alpha=\frac{2800 \pi D}{l_A} \sqrt{Ri_E}$,

where $l_A$ is the convergence length of the estuary in the seawater intrusion zone (assuming $A(x) \sim exp(-x/l_A)$; $Ri_E$ is the Richardson estuary number $Ri_E=g \frac{\Delta \rho}{\rho} \frac{Q_R T^3}{\pi^2 B L^3}$; $\Delta \rho / \rho$ is the relative density difference seawater-fresh water and $B$ the estuarine width at the mouth. The value of $\alpha$ should not exceed 1. Tidal pumping is a major mechanism of seawater intrusion in estuaries during periods of high river flow.

Residence time scale

The residence time $T_r$ is defined as the average time a water parcel located at a distance $x$ from the sea boundary, will take to leave the estuary. If the fresh water discharge is zero or very small, and if the dispersion coefficient $K_x$ is assumed constant along the estuary, the residence time is given by random walk theory :

$T_r = \frac{x^2}{2K_x}$.

The flushing time $T_f=T_x$ is the average time it takes for a fresh water parcel to move through the estuarine zone to the sea. According to the random walk model, for small discharge $Q_R$,

$T_f = \frac{l^2}{2K_x}$ ,

where $l$ is the estuarine length. This is equivalent to $T_f = \frac{V_f}{Q_R}$ , where $V_f$ is the fresh water volume in the estuary.

Experimental determination of the longitudinal dispersion coefficient

The dispersion coefficient $K_x$ can be determined experimentally in situations where the freshwater discharge $Q_R$ is constant over a period longer than $T_x$. In that case the salinity distribution $s_0(x)$ is in equilibrium (time derivative equal to zero). The total residual salt flux $Q_S$ equals zero. We thus have

$AK_x \frac{ds_0}{dx}-Q_R s_0=0$.

Values of the dispersion coefficient can be derived from measurement of the residual discharge $Q_R$ and the salinity distribution $s_0(x)$. Examples of longitudinal dispersion coefficients determined in this way are shown in table 1. It should be noticed that the dispersion coefficient $K_x$ is a function of $x$ and $Q_R$. The dependence of $K_x$ on $Q_R$ has two causes. It is related in the first place to the influence of the salinity distribution on the velocity flow field $u(x,u,z,t)$; such an influence is due to salinity-induced density gradients. In the second place, it is related to the location of the freshwater-seawater transition zone. If this zone is situated near the estuarine mouth, the dispersion coefficient is strongly influenced by tidal pumping. This explains the high longitudinal dispersion coefficients for Rotterdam Waterway, Seine and Loire in table 1, which are determined for situations with important river flow $Q_R/hb$. The same holds, to a somewhat lesser degree, for the Elbe, Weser, Mekong and Sinnamary.

In estuaries with a complex geometry and river flow $Q_R/hb$ several orders of magnitude smaller than the tidal velocity, the influence of salinity-induced density gradients on the longitudinal dispersion coefficient is generally small. This is often the case for tidal lagoons with small river inflow.

If a non-buoyant dissolved substances is introduced in the estuary, it will be mixed over the cross-section after a time $T_A$ . From that time, the substance will be dispersed in longitudinal direction with the same dispersion coefficient as for salinity. For a permanent discharge of a non-buoyant dissolved substance, the same dispersion coefficient applies in the estuarine zones where the substance is mixed over the estuarine cross-section.

Longitudinal dispersion coefficients derived from observed salinity distributions at constant river discharge. The figures for tidal range, depth and width are representative values for the zone with the strongest salinity gradient. Data from , , , , , , .

estuary	tidal range $2a$ [m]	depth $D$ [m]	width $B$ [km]	discharge $Q_R$ [m3/s]	dispersion coefficient $K_x$ [m2/s]
Bay of Fundy (Canada)	10	20	20	150	300
Bristol Channel (UK)	8	30	20	480	60
Chao Phraya (Thailand)	2.5	7.2	0.6	30	330
Corantijn (Surinam)	4.3	6.5	3	500	230
Delaware (US)	1.5	6.6	7	300	300
Eastern Scheldt (Netherlands)	3	12.5	2	60	200
Elbe (Germany)	3.3	12	2.5	750	700
Ems-Dollard (Netherlands)	3	9	4	100	275
Gambia (The Gambia)	1.2	8.7	4	2	200
Hudson (US)	1.6	11,6	2.2	100	110
Humber (UK)	5.5	12	3	250	300
Incomati (Mozambique)	5.5	2.9	0.6	1	10
Limpopo (Mozambique)	2.6	7	0.2	5	150
Loire (France)	4.5	9	0.9	825	900
Mae Klong (Thailand)	2	5.2	0.2	30	200
Maputo (Mozambique)	6.7	3.6	1.3	10	100
Mekong-Co Chien+Cung Hau (Vietnam)	2.1	7	4	2125	570
Mekong-Tran De+Dinh Anh (Vietnam)	2.8	8	3	2250	530
Mersey (UK)	7.5	20	1	80	400
Potomac (US)	1.4	8.4	9	110	70
Pungue (Mozambique)	5	11.5	1.8	20	140
Rotterdam Waterway (Netherlands)	0.9	15	0.6	1000	1000
St. Lawrence (Canada)	3	74	48	8500	200
Seine (France)	5.5	8	0.8	440	800
Sinnamary (Guiana)	2.3	3.8	0.3	100	560
Solo (Indonesia)	1.1	9.2	0.17	10	240
Tha Chin (Thailand)	2.9	5.3	0.2	10	270
Thames (UK)	4.5	12	3	60	100
Weser (Germany)	3.8	9	2	324	1000
Western Scheldt (Netherlands)	3.8	16	3.5	100	200

References

The main author of this article is Job Dronkers Please note that others may also have edited the contents of this article. Citation: Job Dronkers (2016): Seawater intrusion and mixing in estuaries. Available from http://www.coastalwiki.org/wiki/Seawater_intrusion_and_mixing_in_estuaries [accessed on 2-04-2025] For other articles by this author see Category:Articles by Job Dronkers For an overview of contributions by this author see Special:Contributions/Dronkers J