Aquifer Mechanics - FORS(GEOL) 8730

Transcription

1 Aquifer Mechanics - FORS(GEOL) 8730 Todd Rasmussen (2-4300) John Dowd (2-2383) February 11, 2004 Course Summary This course focuses on understanding the mechanics of flow through subsurface media. Theories of flow include: confined homogeneous aquifers, leaky aquifers, delayed yield, effects of partially penetrating wells, unconfined (water table) aquifers, effects of boundaries and multiple wells, dual porosity media, fractured rock aquifers. Applications focus on using pump test data to identify the relative importance of flow and transport processes and to estimate hydraulic properties of aquifers. Textbook Kruseman GP, de Ridder NA, 1991, Analysis and Evaluation of Pumping Test Data, Second Edition, International Institute for Land Reclamation and Improvement, P.O. Box 45, 6700 AA Wageningen, The Netherlands 1

2 Contents 1 Introduction to Aquifer Testing Notation Definitions of Hydraulic Properties Common Errors Aquifer Test Preliminaries Ground-Water Flow Geometries Flow Parallel to Bedding Flow in Series Through Bedding Flow at an Angle to the Bedding The Ground-Water Flow Equation in Radial Coordinates 16 4 Flow in Confined Aquifers Unsteady Flow in Confined Aquifers (Theis Method) Unsteady Flow Problem Example Estimating the Theis Parameters Confined Aquifer Parameter Sensitivity Coefficients Aquifer Derivative Curves Flow in Leaky Aquifers Steady Flow in a Leaky Aquifer Unsteady, Leaky Flow (Hantush-Jacob Method) The Universal Confining Layer (by P. Stone)

3 1 Introduction to Aquifer Testing 1.1 Notation Physical Properties Aquifer thickness, b Aquitard thickness, b Aquifer porosity, n = Vv V b = n e + n i Volume of voids, V v Bulk volume, V b Aquifer effective porosity, n e Immobile zone porosity, n i Bulk density, ρ b = (1 n) ρ s Skeletal density, ρ s = 2.65 g/cm 3 Water content, θ = Vw V v Relative saturation, Θ = theta n = Vw V b Conductance Parameters Aquifer hydraulic conductivity, K = k γ µ Intrinsic permeability, k d 2 Pore diameter, d Fluid specific weight, γ = ρ w g = 9807 P a/m Fluid density, ρ w = 1 g/cm 3 Fluid dynamic viscosity, µ = P a s Aquifer transmissivity, T = K h b Horizontal component of hydraulic conductivity, K h 3

4 Vertical component of hydraulic conductivity, K v or K Anisotropy ratio, m = K h K v = K K Hydraulic conductance, C = Kv = K b b Hydraulic resistance, c = 1 C = b K Aquitard leakance, L = T c = T K b b C K = m b b K xx K xy K x z Hydraulic conductivity tensor, K = K yx K yy K y z K zx K zy K z z Flow Parameters Hydraulic head, h = z + p γ + v2 2 g Elevation, z Fluid pressure, p Fluid specific weight, γ = ρ g Fluid density, rho Gravitational constant, g = 9.807m/s 2 Hydraulic gradient vector, ı = [i x, i y, i z ] = gradh = [ h, h, ] h x y z Fluid flux vector, q = Q A = [q x, q y, q z ] = K h Fluid flow, Q = Vw t Cross sectional area, A Richards equation, q = D θ Aquitard flux, q = K h b = L h Storage Parameters Specific storativity, S s = θ h = [ Vw Vb ] h = Vw V 2 b V b h 1 V b V w h = Vw V b [ 1 V b V b h 1 V w ] V w h = γ n (β w β s ) = γ (α + n β w ) 4

5 Water compressibility, β w = 1 V w V w p = P a 1 = m 1 Aquifer compressibility, α = 1 V b V b p Effective stress, p Bulk volume, V b Skeletal compressibility, β s Barometric efficiency, BE = n/s s Tidal efficiency, T E = 1 BE Aquifer storativity, S = b S s Aquifer hydraulic diffusivity, D = K/S s =T/S Specific yield, S y = b θ h = b θ h Transport Parameters Solute effective dispersion coefficient, D e = D o + α v Molecular diffusion coefficient, D o Solute dispersivity, α Average fluid pore velocity, v Retardation factor, R = 1 + ρ b n K d Distribution coeficient, K d Cation exchange capacity, CEC Anion exchange capacity, AEC Organic carbon content, OC Specific surface area, S a Sorption kinetic parameter, k Additional Aquifer Parameters 5

6 Specific capacity, C p = Q/s Pumping rate, Q Aquifer drawdown, s Recharge rate, R Deep percolation rate, DP 1.2 Definitions of Hydraulic Properties Anisotropic: hydraulic conductivity is differ in different directions. We must represent the hydraulic conductivity, K, using a tensor Anisotropy ratio: Ratio of hydraulic conductivity in one principal direction to the hydraulic conductivity in another principal direction. Aquiclude: A geologic formation in which negligible fluid flow is possible Aquifer: A geologic formation that transmits appreciable quantities of water to a well Aquifer Compressibility: reciprocal of bulk modulus of elasticity of aquifer where is the effective stress. ranges from 10 6 P a 1 for clays, to P a 1 for rock Aquitard or confining layer: geologic formation that resists the movement of water between two aquifers. The layer has material properties K and b to distinguish them from aquifer properties. Flow through a confining layer can be assumed to be vertical: Conductance: hydraulic conductivity of resistive layer (K ) per unit thickness of the resisting layer (b ): Darcian flux: see flux Darcy s Law: relates hydraulic flux to hydraulic gradient and hydraulic conductivity 6

7 Dimension of Flow: Cartesian Flow: flux is different in each of the three cartesian directions Planar Flow: flux is different in two directions, but zero in one cartesian direction Linear Flow: flux is different in one cartesian direction, but zero in the other two Radial or Cylindrical Flow: flux is planar, radiating inward or outward from a central axis. Spherical Flow: flux is radiating uniformly in all three directions from a point Discharge: Water moving from the subsurface to the surface across the earth s surface Drawdown: water level decline due to pumping from well, where h i is baseline (initial) water level before pumping and h is observed water level during aquifer test. Flux: Volume of water (V w ) flowing per unit area (A) per unit time (t) Homogeneous Media: K is a constant is space Hydraulic Resistance: the reciprocal of the conductance (C) Hydraulic conductivity: Measure of ability of geologic media to transmit water, related in a general way to pore size and shape: C = constant of proportionality d = median pore or grain size γ specific weight of fluid used to measure total head µ = water dynamic viscosity 7

8 Hydraulic Diffusivity: ratio of hydraulic conductivity to specific storage, or, equivalently, of transmissivity to the storage coefficient Hydraulic Gradient: Change in total head per unit distance (h) Infiltration: Water passing across the earth s surface into the subsurface Isotropic: hydraulic conductivity is the same in all directions. It can be represented as a scalar Leakance, or Leakage factor Microporosity or matrix porosity: pores too small to see, such as voids between mineral grains or clay platelets. Macroporosity: visible pores, such as fractures, voids or vugs Nonlinear: K is not constant for all values of total head or fluid potentials Percolation: Water moving through the unsaturated zone Porosity: volume of voids (V v ) per unit bulk volume (V b ), ranging from 10 5 for dense granites to over 0.70 for volcanic tuffs. Precipitation: Water falling on the surface of the earth Recharge: Water moving across the water table from the unsaturated zone into the saturated zone Relative Saturation: Water content of material relative to saturated water content: Richards Equation: relates hydraulic flux to change in volume of water in storage Specific Capacity: water discharge from well divided by amount of drawdown Specific Storage: A measure of the volume of water released (or added) to storage, per unit volume of aquifer, per unit increase (or decrease) in fluid pressure. It 8

9 is the sum of the aquifer plus water compressibilities: where is the fluid specific weight, is the bulk compressibility, is the compressibility of water, and = - n is the compressibility of the mineral skeleton. Specific Yield: volume of water released (or added) from an unconfined aquifer per unit area of aquifer per unit decline (rise) in water table position. Similar to the storage coefficient for confined aquifers. Storage Coefficient: water released from storage for a given thickness of aquifer (b). Storativity: Identical to the storage coefficient Total Head: sum of elevation, pressure, velocity, osmotic and other potentials Transmissivity: Total hydraulic conductivity for a given aquifer thickness Water Compressibility: reciprocal of bulk modulus of elasticity of water where p is the water pressure. Water Content: volume of water (V w ) per bulk volume of aquifer (V b ) 1.3 Common Errors DANGER #1: q and v are not the same! q is specific flux (units of velocity) v is groundwater velocity (units of velocity) DANGER #2: Darcy s Law is not always valid! Darcy s law is only appropriate for laminar, and not turbulent flow, turbulent conditions arise in large voids (fractures, caves, etc.), and where velocities are high, such as near boreholes. 9

10 The Reynolds number helps to define the conditions for laminar flow d is void opening, is kinematic viscosity (approximately 10 6 m 2 /s), R is the ratio of inertial to viscous forces. When the inertial forces are too great, such as when R > 1, then flow can become turbulent. For turbulent conditions, the flow rate varies with the square root of the gradient. 1.4 Aquifer Test Preliminaries Definition of Test Objectives scientific vs engineering objectives identification of system behavior estimation of system parameters scale of interest local vs regional test short-term vs long-term effects hydraulic vs transport effects average or extreme behavior fluid vs solute migration Definition of Conceptual Model Define state of knowledge of area Type of aquifer Type of geologic media, homogeneous/heterogeneous, isotropic/anisotropic Geometries of aquifers and confining units Recharge and discharge areas 10

11 Inventory existing wells in area for baseline (reference) data Find local and regional flow gradients, horizontally and vertically Use the ground and surface water chemical characteristics to indicate of flow system Physical Installation Pumping well: Type of gravel or sand pack Depth of well Well and borehole diameters Length of screened interval and screen integrity Depth of pump, pumping capacity, and previous pumping records Measurement of pumped water rate and duration Location of discharge Observation wells: Number and location (distance from well and depth): Barometric pressure Precipitation Nearby surface water levels Reference ground water levels Aquifer test data: Discharge sampling (flow meter, weir, bucket, etc) and control (valves, voltage) Manual (steel tape, electrical sounder) vs automated (pressure transducer, float) Analog (strip charts) vs digital (computer) 11

12 Water level sampling rates and duration Well effects (cascading water, water quality changes) Aquifer test data interpretation Removing background effects Loading effects due to barometric pressure, rainfall, oceanic and earth tides Trends (seasonal or short-term) Fluid column density effects (salinity, temperature, dissolved gasses) Identifying aquifers type confined aquifer (leaky, double porosity, fractured rock) unconfined aquifer (delayed yield) Identifying other effects well bore storage (and excessive pumping rate) partial penetration boundaries Standard plots horizontal-axis: log t or log(t/r 2 ) vertical-axis: s or log s Derivative plots horizontal-axis: log t or log(t/r 2 ) vertical-axis: s/ t or s/( log t) 12

13 2 Ground-Water Flow Geometries 2.1 Flow Parallel to Bedding Let us assume that two permeable layers overlie each other with horizontal flow in each layer, and no flow between each layer. The flow in each layer is: Q 1 = Q 2 = h K 1 A a h b 1 x a x b (1) h K 2 A a h b 2 x a x b where A 1 = w b 1 and A 2 = w b 2. We sum the amount of flow in each layer: Q h = Q 1 + Q 2 Q h = K h A ha h b x a x b (2) where A = A 1 + A 2 = w (b 1 + b 2 ) = w b, and K h is the effective horizontal hydraulic conductivity. Substitution results in: K h A ha h b x a x b = K 1 A 1 h a h b x a x b + K 2 A 2 h a h b x a x b K h A = K 1 A 1 + K 2 A 2 A K h = K K A 2 A 2 A K h = K 1 b 1b + K 2 b 2b (3) We can also write this in terms of transmissivity: T e = T 1 + T 2 (4) because T e = K e b, T 1 = K 1 b 1, and T 2 = K 2 b 2. For multiple layers, we have: T e = n i=1 T i K e = n i=1 b i K i n i=1 b i (5) 2.2 Flow in Series Through Bedding In this case the flow is perpendicular to the beds: Q 1 = Q 2 = K 1 A ha h b z a z b (6) K 2 A h b h c z b z c 13

14 where A is the same for each bed. In this case, the flows are equal to each other: Q v = Q 1 = Q 2 K v A ha hc b = K 1 A ha h b b 1 = K 2 A h b h c b 2 (7) K v h a h c b = K 1 h a h b b 1 = K 2 h b h c b 2 where b = z a z c, b 1 = z a z b, b 2 = z b z c. Solving for the intervening head, h b, yields: h b = K 1 h c b 1 h a + K 2 b 2 K 1 b 1 + K (8) 2 b 2 Substituting into the previous equation yields: K v b h a h c = K 1 b 1 [ K v 1 b = b 1 + b 2 K 1 K 2 h a b K v = b 1 K 1 + b 2 K 2 K 1 b 1 h a+ K 2 b 2 h c K 1 + K 2 b 1 b 2 ] (9) or, by using the hydraulic resistance, c = K, we have: b c v = c 1 + c 2 (10) Thus, like transmissivity, the hydraulic resistance adds. For multiple aquifers this is: c v = n i=1 c i (11) K v = ni=1 ni=1 b i b i (12) K i 2.3 Flow at an Angle to the Bedding In general, we can define the horizontal and vertical components of flux using: q = Q A = [q h, q v ] = [K h i h, K v i v ] (13) where i h = h and i x v = h z The magnitude of the hydraulic gradient, i, and flux q, are: i 2 = q 2 = 14 i 2 h + i 2 v q 2 h + q 2 v (14)

15 The flux components can be determined from the direction, φ and magnitude, q, of flow: Substitution provides: q h = q v = q cos φ q sin φ (15) (K e i ) 2 = (K h i h ) 2 + (K v i v ) 2 (16) Simplification yields: K e = i 2 h i 2 K 2 h + i2 v i 2 K 2 v cos 2 φ K 2 h + sin2 φ K 2 v K h cos 2 φ + m sin 2 φ (17) where m = K v /K h is the anisotropy ratio. If the hydraulic gradient is at an angle of 45 to the bedding planes, we have: K e = K 2 h + K2 v 2 = K x 1 + m 2 2 (18) If the flux is at an angle of 45 to the bedding planes, we have q h = q v, so that K h i h = K v i v, and m = i x /i v, resulting in: K e = K h 2 m m 2 (19) 15

16 3 The Ground-Water Flow Equation in Radial Coordinates We derive the well flow equation by first defining a representative cylindrical volume with fixed height, b, an inner radius, r 1, and an outer radius, r 2 = r 1 + r. Flow across the inner surface, Q 1, and outer surface, Q 2, are: Q 1 = q 1 A 1 Q 2 = q 2 A 2 (20) with A i = 2 π r i b and q i = K ( ) h, and where A is the cross-sectional area of the r i cylinder perpendicular to the flow direction, q is the darcian flux across the cylindrical face, K is the hydraulic conductivity of the medium within the cylinder, i = h r is the hydraulic gradient, h is the hydraulic head, and r is the radial distance from the center axis of the cylinder. Combining terms yields: Q i = 2 π b K r i We now specify the mass balance equation as: ( ) h r i (21) Q = Q 1 Q 2 = V w t (22) where V w is the change in volume of water within the cylinder needed to balance the inflows with the outflows, and T is time. The change in the volume of water within the cylinder is related to the change in the water content, θ, within the cylinder using: where V b is the volume of the hollow cylinder, equal to: θ = V w V b (23) V b = V 2 V 1 = b π [ r 2 2 r 2 1 ] = b π r 2 (24) We can relate the hydraulic head to the water content using: S s = θ h (25) 16

17 where S s is the specific storage coefficient. Combining equations yields: 2 π b K [ ( ) h r 2 r 2 ( ) ] h r 1 = π b S s r 2 h t (26) r 1 which is the same as: 2 K [ r ( )] h r 2 K [ r ( h r r 2 = S s r 2 h t )] h = S s t )] (27) (28) 2 D [ r ( h r = S r 2 s h t (29) (30) where D = T S is the hydraulic diffusivity, and the partials are introduced to denote infinitesimal changes. We can simplify this further by noting that r2 r = 2 r: D [ r ( )] h r = h t (31) [ r r 2 ] h h D + 1r = h t (32) r2 r 2 h h + 1r r2 r = 1 h t (33) D (34) 17

18 4 Flow in Confined Aquifers 4.1 Unsteady Flow in Confined Aquifers (Theis Method) We start with the ground-water flow equation: T [ 2 h r + 1 ] 2 r h r = S h t [ 2 h D r + 1 ] 2 r h r = h t 2 h r r h r = 1 D h t (35) (36) (37) (38) 4.2 Unsteady Flow Problem Example Find the distance from a pumping well where the drawdown is 1 foot, given: Radius of pumping well, r o = 2 = ft Hydraulic conductivity, K = ft/min Aquifer saturated thickness, b = 10 ft Specific yield, S y = 0.1 ft 3 /ft 3 /ft Pumping duration, t = 3 days = 72 hrs = 4320 min Drawdown at pumped well, s o = 10 ft Derived Parameters: Aquifer Transmissivity, Storage Coefficient, and Diffusivity are: T = S = D = Kb = ( ft/min) (10 ft) = ft 2 /min S y b = [0.1 (ft 3 water)/(ft 3 soil)/(ft drawdown)] (10 ft soil) = 1 ft 3 /ft 2 /ft T/S= (K b)/(s y b) = K/S y = ft/min/0.1 ft 3 /ft 3 /ft = ft 2 /min (39) The well function argument is: u o = r2 o 4Dt = ( ft) 2 4 ( ft 2 /min) (4320 min) = (40) 18

19 This satisfies the Jacob condition: u < 0.01, so the well function at the pumped well is: W (u o ) = ln u = 6.55 (41) The pumping rate must be: Q = 4T s o W (u o ) = 4 ( ft 2 10 ft /min) 6.55 = ft3 /min = gpm (42) The well function at the observation well is simply: W (u 1 ) = W (u o ) s 1 s o = ft 10 ft The well function argument at the observation well is: = (43) u 1 = exp( W (u 1 ) = (44) The distance at which the drawdown is 1 foot is: r = 4 D t u 1 = 4 ( ft 2 /min) (4320 min) = 3.18 feet (45) 4.3 Estimating the Theis Parameters The following data are normally provided for an aquifer test: Q, aquifer pumping rate (constant) r, distance of observation well from pumping well (constant) s, drawdown vector in observation well t, time vector at which drawdowns are measured We can remove the constants Q and r from consideration by defining two normalized variables, a normalized drawdown, w, and a normalized time, tau : so that the Theis Solution becomes: w i = 4 π s i Q τ i = 4 t i (46) r 2 w i = W (u i) u i = 1 (47) T D τ i Our objective is to determine the following aquifer parameters: 19

20 D, aquifer diffusivity T, aquifer transmissivity S = T/D, aquifer storage coefficient It is clear that there are two parameters to determine, T and D. The storage coefficient is determined once D and T are known. Two unknowns require two equations, which we select to be the observed drawdowns at two separate times: w 1 = W (u 1) T w 2 = W (u 2) T (48) (49) u 1 = 1 D τ 1 (50) u 2 = 1 D τ 2 (51) (52) We can remove T from these equations by taking the ratio of the normalized drawdown at two different times: ln w 1 = W (u 1) (53) w 2 W u 2 w 1 W (u 2 ) = 1 (54) w 2 W u [ 1 ] w1 W (u 2 ) = 0 (55) w 2 W u 1 (56) which is now only a function of the unknown diffusivity, D. We solve for this unknown at each time step using Newton s method: D j+1 = D j F (D j) f(d k ) The functions, F (D) and f(d), are defined as: F (D j ) = ln [ ] w1 W (u 2 ) w 2 W u 1 (57) (58) 20

21 f(d j ) = df (D) dd (59) (60) Solving for f(d) is assisted by use of the chain rule: d [ln y] dd = d [ln y] dy y u u D (61) We can show that: f(d) = 1 D [ exp u2 W (u 2 ) exp u ] 1 W (u 1 ) (62) so that: D j+1 = D j 1 ln [ w 1 W (u 2 ) w 2 W (u 1 ) e u 2 W (u 2 ) e u 1 W (u 1 ) ] (63) Iteration proceeds until the aquifer diffusivity converges. Upon convergence, we calculate the aquifer transmissivity and storage coefficient for each time step using: T i = W (u i) w i (64) S i = T i D i (65) 4.4 Confined Aquifer Parameter Sensitivity Coefficients Confined aquifer sensitivity coefficients are used to relate changes in drawdown (raw, U, and normalized, U ) to each of the coefficients that determine drawdown. s = U r = U t = U T = U S = U D = Q 4 π T W (u) s = Q r 4 π T s = t s = T Q 4 π T Q 4 π T 2 e u r e u t e u W (u) T s = Q S 4 π T s = D Q 4 π T e u S e u T (66) The sensitivity between T and D is: D T = U T U S = e u W (u) S e u (67) 21

22 The sensitivity between T and S is: S T = U T U S = Note that T and S are independent of each other when: W (u) e u D e u (68) S T = 0 (69) which occurs when W (u) = e u. T and D are independent for these same conditions. In fact, drawdowns are insensitive to transmissivity at this point, because U T = 0 for this condition. This condition is satisfied at u = Aquifer Derivative Curves We calculate derivative curves using observed drawdowns: For normalized variables, this is: s = U t = s t = Q 4 π T e u t (70) w = U t w = 4 π s Q = e u T τ = W (u) T (71) (72) τ = 4 t r 2 (73) u = 1 D τ (74) The maximum of this function is found by setting the derivative of w to zero: (75) w = 2 w τ 2 = e u (u 2 1) T τ 2 = 0 (76) which occurs when u = 1. This implies that at the time of the maximum, t, and the value of drawdown at the maximum, w, can be used to determine aquifer parameters, D = 1/t, because: T = W (1) = w w w = S = W (u) T T D = τ w (77) 22

23 5 Flow in Leaky Aquifers 5.1 Steady Flow in a Leaky Aquifer For a leaky aquifer, we have: T [ 2 s r + 1 ] 2 r s r q = 0 (78) We account for leakage using: q = K b s = s c (79) Dividing by the aquifer transmissivity yields the equation: where L 2 = T c. 2 s r r s r s L 2 = 0 (80) This form is similar to the Modified Bessel Function Differential Equation (Abramowitz and Stegun, 9.6.1): z 2 2 w z 2 + z w z ( z 2 + v 2) w = 0 (81) for the conditions: v = 0 (82) s = w (83) z = r L (84) (85) The Modified Bessel Function solution for v = 0 is: w = K o (z) = for boundary conditions: o cos zt t2 + 1 dt = o lim z w = 0 lim z 0 z w z = 1 cos(z sinh t) dt (86) (87) 23

24 Our boundary conditions are: which yields the solution: lim r s = 0 lim r 0 s = r s r = Q 2 π T Q ( ) r 2 π T K o L (88) (89) Leakage into aquifers during well tests often result in steady flow conditions. This attribute is quantified using the specific capacity coefficient, C p : C p = Q s = 2 π ( T ) K r (90) o L 5.2 Unsteady, Leaky Flow (Hantush-Jacob Method) For unsteady flow conditions we have: We again account for leakage using: T [ 2 s r + 1 ] 2 r s r S s t q = 0 (91) q = K b s = s c (92) Dividing by the aquifer transmissivity yields the equation: where L 2 = T c. 2 s r r s r 1 D s t s L 2 = 0 (93) Using the same approach as the solution for the confined (Theis) equation: which is equivalent to: u 2 u [ 2 ] s u + s u + s 2 u r2 s 4 u L = 0 (94) 2 2 s s + u [u + 1] s u + u2 u r2 s 4 u L = 0 (95) 2 We can find an approximate solution for Jacob conditions, i.e., u << 1: u 2 2 s s + u s u + u2 u [ u 2 + v 2] s = 0 (96) 24

25 where: yielding: v = s = r 2 4 L 2 u2 (97) Q 4 π T K v(u) (98) which holds as long as u < 0.01 and r > u 2 L Hantush and Jacob found a solution for all u: s = Q 4 π T which is also known as the Walton solution. u exp [ x x ] r2 4 r L 2 dx (99) 5.3 The Universal Confining Layer (by P. Stone) An unrecognized protector of our ground water There is a unique geological unit economically critical, hydrologically essential, extremely widespread (ubiquitous among some investigated environments) that amazingly has not been formally recognized and described as an interrelated unit, and not named according to the rules for stratigraphic nomenclature. This unit, whose lithology and thickness varies among sites but is always highly impermeable, even when very thin, is found to lie atop the first usable drinking-water aquifer apparently virtually everywhere. Or, perhaps one should say, it is found in virtually every place where there has been some shallow stratigraphic and geo-hydrologic investigation that could reveal it; these of course being at sites where actual, or suspected, or potential, ground-water contamination has prompted such investigation. Surprisingly, this unit almost always perches some completely unconnected, sometimes ephemeral, always minor and inconsequential, zone or ground water atop it, that being the contaminated zone where present. Or else no water table whatsoever reportedly exists above the confining layer for the underlying aquifer in hydraulic nature even though it often occurs at an uncommonly shallow depth for such an aquifer. 25

26 This formation and confining layer is unique not only because of its near ubiquity among the widely scattered investigated sites, but because it encompasses various different geologic formations and, astoundingly, completely different geologic provinces also. What other geologic formation in the shallow subsurface encompasses both Appalachian Piedmont and Atlantic Coastal-Plain provinces, of vastly different ages and especially origins? Host materials range from soft sediments to residua from hard crystalline rocks (curiously, the same condition, if not the layer itself, is commonly reported for fractured hard-rock aquifers also). One might conjecture that this layer relates somehow to alluvial deposition, stream valleys certainly being one of the few depositional environments to span both Piedmont and Coastal-Plain geologic provinces. Or perhaps it is some unrecognized sub- C soil horizon or phenomenon, formed in-place. Will close examination of contamination-site investigations from other geologic provinces, in other parts of the country, reveal this same formation? Is it found still further afield, at least at sites where there is some exploratory interest able to find it? The economic importance of this stratum cannot be overestimated. By universally isolating the uppermost true aquifer from the host of dangerous materials at the ground surface above, or especially those dissolved in the shallowest perched or ephemeral ground water so nearby above, the environmental Cleanup costs that are avoided, or deemed unnecessary, are astronomical in their cumulative total, and represent some very considerable sum at any given site also. Other distinctive distributional characteristics include (I thank a colleague for this observation) the apparent tendency, judged from reported contaminant-plume geometry, for the confining unit to rise near property boundaries. Incidentally, I hope the identification of this physical mechanism lays to rest the suspicious sometimes voiced concerning the numerous ground-water contaminant plumes that seem never to migrate out from the responsible party s property to another s, ending instead somewhere near the boundary. The averted costs in liabilities add another vast value to this unit. One accepts as obvious that this geographic coincidence of property boundaries and 26

27 contaminant plumes and their controlling stratigraphy cannot be random coincidence and that property lines themselves cannot influence the underground. Could there in any way be some nonapparent control made by the underground environment on the boundaries? At first glance, of course, this would seem to be impossible. But perhaps it involves unknowing influence on decisions in platting, somewhat like the suspected subconscious reading of subtle topographic and other surficial signs by water witchers who then truly believe that they feel the divining rod move as they pass over an underground stream. What other reasonable explanation can there be for these fortuitously placed natural liners, and even natural slurry walls, containing the future spread of ground-water contaminants at the many sites that so desperately need them? How did the extremely widespread occurrence of this single stratum, or at least an analogous stratigraphic sequence, escape geologic investigation s attention prior to the proliferation of contamination-site hydrologic investigation? Dare one put forth an outrageous hypothesis? Could some action of the contaminants themselves create this layer, say perhaps by somehow chemically reducing permeability. What other physical or cultural characteristic of ground-water contamination, or of such sites, might cause or explain this prevailing presence? This remains as grist for future research. In any case, if this important zone exists almost everywhere and a host of apparently irrefutable professional sources have identified it at scores of sites it should be recognized and named formally, as are all significant geologic units no matter how emplaced. I have sought advice from colleagues regarding a descriptive or distinctive name for the layer or formation, with a diversity of suggestions received: From the artificial constructs Nofurac or Commonly Formations (no further action or continued monitoring only), to the Linnean binomial Confinus Maximus, to the Victorian-era-sounding The glorious [or great or grand ] hydraulic impedance. My personal favorite, and therefore the winner, is the Nocostus Aquitard, and thus it is given. 27