Aquifer

An aquifer is a permeable geological formation containing and transmitting groundwater. A confined aquifer lies beneath a confining layer; an unconfined aquifer instead has a free water table.

An aquifer containing saline rather than fresh water. A deep saline aquifer can provide pore space for carbon capture and storage, but safe retention depends on reservoir connectivity, continuous caprock and pressure management.

An aquifer drainage hydrograph plots the time-dependent discharge from an aquifer. In the shallow constant-recharge model, early discharge grows linearly according to the rain-fed aquifer similarity solution. Sustained recharge $R$ on a finite length $L$ gives eventual steady discharge $RL$ when the other end is impermeable. After recharge stops, a separable aquifer drawdown has height proportional to $1/(t+t_0)$ and discharge proportional to $1/(t+t_0{)}^2$; this is generally a late-time description rather than an exact solution for arbitrary inherited height profiles.

An unconfined aquifer has a free groundwater table. In a shallow aquifer of thickness $h(x,t)$, hydrostatic pressure and Darcy's law give horizontal flux proportional to $-h{h}_x$.

The water table is the upper surface of the saturated zone in an unconfined aquifer, where liquid pressure equals atmospheric pressure in the simplest drainage model. Under the Dupuit approximation, its height determines hydrostatic pressure throughout the saturated region and therefore the lateral Darcy flux.

For pore fraction $n$, permeability $k$, liquid density $\rho$, viscosity $\mu$, rainfall recharge $R$, and a shallow groundwater thickness $h$,This is a nonlinear diffusion equation with diffusivity proportional to the local thickness.

A zero-height outlet can carry finite nonzero volume flux per unit width in a nonlinear Darcy flow model with $q=-Kh{h}_x$. If $h{h}_x\to J>0$ as $x\downarrow 0$, then $(h^2{)}^{\prime }\to 2J$, and integration gives $h\sim \sqrt{2Jx}$ and $h_x\sim \sqrt{J/(2x)}$. Thus the height vanishes while its slope diverges; the flux must be evaluated as a limit of the product. The divergent slope also identifies a region where the outer Dupuit approximation eventually needs a local correction.

For the dimensionless rainfall-driven Boussinesq equation for an unconfined aquifer, $h_t=(h{h}_x{)}_x+1$, balancing every time power in $h=t^{a}f(x/t^b)$ gives $a=b=1$. The similarity profile obeys $(f{f}^{\prime }{)}^{\prime }+\eta f^{\prime }-f+1=0$. A river drained to base level gives $f(0)=0$, while undrained storage far away gives $f(\infty )=1$. The solution approximates a finite aquifer before its drainage-affected length reaches the far boundary. If $J={\lim }_{\eta \downarrow 0}f{f}^{\prime }>0$, the dimensional drainage flux is $\sqrt{K}{R}^{3/2}Jt/\varphi$.

For $h_t=\kappa (h{h}_x{)}_x$ with a drained boundary $h(0,t)=0$ and a finite current vanishing at its nose, integration by parts gives $\dot{D}=-\kappa [h^2{]}_0^{\infty }/2=0$. The outlet can have finite volume flux because its contribution to the first-moment flux is multiplied by $x=0$. A positive prescribed boundary height instead gives $\dot{D}=\kappa h(0,t{)}^2/2$, so drainage is essential to this invariant.

The conserved first moment of a draining porous current fixes $\alpha +2\beta =0$ in a self-similar solution $h=a{t}^{\alpha }f(x/(d{t}^{\beta }))$. The Boussinesq equation for an unconfined aquifer supplies $\alpha =2\beta -1$, giving $\alpha =-1/2$, $\beta =1/4$. Normalize $\kappa a=d^2$ and the nose at $\eta =1$. Then $-f/2-\eta f^{\prime }/4=(f{f}^{\prime }{)}^{\prime }$ is solved byThe square-root outlet has finite drainage flux and the nose has zero flux. This is a particular similarity solution and its scaling family, not an exact description of arbitrary initial data.

For the dipole similarity solution of a draining porous current, $V=\varphi \int hdx=\varphi ad{t}^{-1/4}/18$. Thus $\dot{V}/V=-1/(4t)$. A shifted similarity origin replaces $t$ by $t+t_0$. Neither a positive rate nor a rate $-1/(2t)$ is compatible with this two-dimensional first-moment-conserving profile.

After recharge stops on a finite aquifer connected to a river, the late solution of the Boussinesq equation for an unconfined aquifer has the form $h\sim t^{-1}F(x/L)$. The profile solves a nonlinear eigenvalue problem, and river discharge decays as $t^{-2}$.

The Dupuit approximation treats flow in a shallow unconfined aquifer as predominantly horizontal, with hydrostatic pressure $p=\rho g(h-z)$ and horizontal pressure gradient independent of depth. Integrating Darcy's law over the saturated depth then gives a volume flux per unit width depending only on the groundwater height and its slope.

For saturated Darcy flow through a long dam occupying a channel of width $\alpha z^2$, let hydraulic conductivity be $K=k\rho g/\mu$ and free-surface height $h(x)$. The Dupuit approximation gives $Q=-K\alpha h^3{h}_x/3$. Fixed heads $h(0)=H$ and $h(L)=\eta$ give $h^4=H^4-(H^4-{\eta }^4)x/L$ and the displayed steady discharge. Porosity relates pore velocity to superficial Darcy velocity but does not independently multiply this discharge when $k$ is intrinsic permeability. For depth-dependent permeability, integrate the horizontal conductivity over the saturated cross-section first.

With permeability of a porous medium $k(z)=k_0(1+\beta z)$, the Dupuit approximation givesHere $\varphi$ is porosity and $R$ is recharge. The low-height limit has mobility proportional to $h$, whereas large heights have mobility proportional to $h^2$. This changes both the forced filling similarity for power-law diffusion and the separable draining profile for power-law diffusion.

Even when most of an unconfined aquifer with depth-dependent permeability is deep, its height vanishes at an absorbing outlet. Near that outlet, finite outward volume flux per unit width $Q$ gives $u_{b}h{h}_x\simeq Q$ and hence $h\simeq (2Qx/u_b{)}^{1/2}$. The high-depth outer approximation instead has $h\simeq [6Qx/(u_b\beta ){]}^{1/3}$. Their crossover lies at height of order $1/\beta$ and distance of order $u_b/({\beta }^{2}Q)$; the thin inner boundary layer transmits the same leading discharge.

A poroelastic aquifer couples pore-fluid storage and flow to elastic compaction of its solid skeleton. Changes in gravity, surface loading, or pore pressure can therefore change groundwater discharge even without changing rainfall.

For a skeleton with Young modulus $E$, undeformed solid fraction ${\varphi }_0$, grain-fluid density difference ${\rho }_s-\rho$, and reference gravity $g_0$, the hydrostatic poroelastic compaction length isIt is the depth over which self-weight changes the solid fraction appreciably.