Geophysical Investigation of Seepage Beneathan Earthen Damby S.J. Ikard1, J. Rittgers1, A. Revil2,3, and M.A. Mooney4

AbstractA hydrogeophysical survey is performed at small earthen dam that overlies a confined aquifer. The structure of the dam has not

shown evidence of anomalous seepage internally or through the foundation prior to the survey. However, the surface topographyis mounded in a localized zone 150 m downstream, and groundwater discharges from this zone periodically when the reservoirstorage is maximum. We use self-potential and electrical resistivity tomography surveys with seismic refraction tomography to(1) determine what underlying hydrogeologic factors, if any, have contributed to the successful long-term operation of the damwithout apparent indicators of anomalous seepage through its core and foundation; and (2) investigate the hydraulic connectionbetween the reservoir and the seepage zone to determine whether there exists a potential for this success to be undermined.Geophysical data are informed by hydraulic and geotechnical borehole data. Seismic refraction tomography is performed to determinethe geometry of the phreatic surface. The hydro-stratigraphy is mapped with the resistivity data and groundwater flow patternsare determined with self-potential data. A self-potential model is constructed to represent a perpendicular profile extending outfrom the maximum cross-section of the dam, and self-potential data are inverted to recover the groundwater velocity field. Thegroundwater flow pattern through the aquifer is controlled by the bedrock topography and a preferential flow pathway existsbeneath the dam. It corresponds to a sandy-gravel layer connecting the reservoir to the downstream seepage zone.

IntroductionEarthen dams are susceptible to anomalous seepage

through their cores and foundations. Anomalous seepage,including elevated and/or concentrated seepage, is a cata-lyst for internal erosion and is therefore a primary causeof earthen dam failures (Foster et al. 2000, 2002). Failuremay occur upon the first filling of a reservoir or after manyyears of undetected seepage (Fell et al. 2003). Geotechni-cal and geophysical studies are typically performed afterseepage indicators manifest at the ground surface (Butleret al. 1990; Panthulu et al. 2001; Titov et al. 2002; Rozy-cki et al. 2006; Boleve et al. 2011). Recently, researchershave shown the improved capabilities of geo-electric mon-itoring, with and without the injection of conductive trac-ers, to localize preferential flowpaths (Titov et al. 2002;Ikard et al. 2012; Revil et al. 2012).

1Department of Geophysics, Colorado School of Mines, GreenCenter, Golden, CO.

2ISTerre, CNRS, UMR CNRS 5275, Universite de Savoie, LeBourget du Lac, France.

3Corresponding author: Department of Geophysics, ColoradoSchool of Mines, 1500 Illinois Street, Green Center, Golden, CO80401; (303) 273–3512; fax: (303) 273–3478; arevil@mines.edu

4Department of Civil and Environmental Engineering, ColoradoSchool of Mines, Golden, CO.

Received July 2013, accepted February 2014.© 2014, National Ground Water Association.doi: 10.1111/gwat.12185

Geophysical studies of dams are typically not per-formed if anomalous seepage or internal erosion indicatorsare not present. Perhaps such studies seem unnecessary ifthere is no seepage to contrast the “normal” seepage flowregime. However, it may be insightful to investigate nor-mal seepage in dams for the specific factors and hydraulicconditions that contribute to the overall long-term suc-cessful operation of the dam. Our primary goal in thispaper is to investigate the hydraulic connection betweenthe impounded reservoir and the downstream seepage dis-charge zone. We are looking for to a plausible explanationfor why a specific dam located in Colorado has shown along record of successful performance without developinganomalous seepage through its core and foundation?

To answer this question, we combine three methodsthat are commonly used for groundwater studies, but thatare each fundamentally sensitive to different properties ofthe subsurface: (1) 2D seismic refraction tomography ofthe P-wave velocity, (2) self-potential (SP) mapping, and(3) electrical resistivity tomography (ERT). The seismicrefraction method shows the position of the water table,the self-potential method is directly sensitive to the flowof groundwater, and ERT provides a way to localize theinterface between the bedrock and the variably saturatedembankment soil. Our interpretation of these geophysicaldata will be also informed by additional hydraulic andgeotechnical data that have been measured in piezometersin the dam cross-section and foundation.

238 Vol. 53, No. 2–Groundwater–March-April 2015 (pages 238–250) NGWA.org

(a) (b)

(c)

Figure 1. Description of the field site. (a) 1:24,000 topographic map of field site showing survey area, seepage location andposition of site photos relative to the dam crest and reservoir (map published by USGS, 1965 and inspected/revised 1994,reference code 39105-G2-TF-024). (b) Photo taken on the dam crest at the south abutment facing northeast towards location.(c) Photo taken on profile P7 facing west toward the toe access road and the dam crest.

Description of the Test Site

Localization and GeometryThe field site is a homogeneous earth-fill dam located

at the base of the Rocky Mountain foothills in JeffersonCounty, Colorado. The topography is shown in Figure 1aand photos of the dam crest from the right (south)abutment and the maximum cross-section of the dam areshown in Figure 1b and 1c. The topographic gradientwithin the survey area is mild and points towards the eastat the center of the dam crest. The topographic variabilityin a N-S direction along the strike of survey profiles P1-P6is minor. A 1:3,600 aerial photograph of the survey gridand the reservoir geometry is given in Figure 2. The damis 3.7 m wide at the crest and 427 m long, and is composedof homogenous earth fills directly borrowed from onsitematerials excavated from the reservoir basin (DenverWater 1988). It is 4 m tall at the maximum cross-sectionand the hydraulic height is 3.7 m. The upstream anddownstream slopes have a 2.5:1 grade. The impoundedreservoir (Figure 2) stores a maximum of 3.60 × 105 m3

of water at the high maximum pool elevation of 1800.4 m.The normal storage elevation of the reservoir is 1800.1 m.At this elevation the reservoir surface area is 1.1 × 105 m2

and the storage capacity is 2.5 × 105 m3.

Geology and Geotechnical PropertiesThe reservoir lies entirely within the upper portion of

the Laramie formation, although the west and southwestreservoir rims are faulted against the impermeable Pierreshale formation (Denver Water 1988). The engineering

Figure 2. Aerial photo of the survey area showing thelocation of access roads and the downstream seepage zone.The dam is 3.4-m high and positioned between the toe accessroad and the crest access road at the cross-section of profileP7. Station markers represent the position of the electrodesfor the electrical resistivity tomography and the self-potential(white filled circles) and the two seismic tomography fieldstations (red lines). Profile P2, positioned at the toe betweenprofiles P1 and P3, has been omitted for clarity.

design records for the dam describe the site as a four layersystem composed from top to bottom of (1) a maximumof 4 m of embankment fill compacted to 94% of thestandard Proctor maximum dry density, (2) a layer ofsandy to very sandy “natural” unconsolidated clay that

NGWA.org S.J. Ikard et al Groundwater 53, no. 2: 238–250 239

Table 1Model Parameter Summary for Modeled Soil

Textures

PropertyNatural

Clay Aquifer Bedrock

Percent gravel1 (% wt.) 0.00 50.0 0.00Percent sand1 (% wt.) 32.0 18.0 32.0Percent silt1 (% wt.) 24.0 0.00 24.0Percent clay1 (% wt.) 44.0 32.0 44.0Compaction1 0.94 0.94 1.18Residual water content2

(% vol), θ r

26.2 11.1 26.2

Porosity2 (−), φ = θ s 0.49 0.43 0.36Hydraulic conductivity2,3

(mm/hr), K s

2.0 5.0 0.04

Bulk density2 (kg/m3), ρ 1,360 1,510 1,690van Genuchten parameter2

α (m–1)0.36 5.7 × 10–3 3.0 × 10–3

van Genuchten parameter2

n (−)1.4 0.73 0.93

Formation factor4 (−), F 6.0 8.3 12.9Excess charge density5

(C/m3), QV

9.89 2.91 1864

Bulk conductivity6

(mS/m), σ

8.2 5.9 3.8

ηf denotes the dynamic viscosity of the pore water (10–3 Pa s at 25◦C).1From borehole grain size distribution data supplied by the dam owner.2From SPAW soils database of Saxton and Rawls (2006).3Based on prior geotechnical tests.4Using F = φ–m with m = 2.5 for clayey sediment (Revil et al. 2012).5Using C = −QV kρf g/

(σηf

), where σ = σ f /F with σ f = (4.9 ± 0.2)

× 10–2 S/m (pore water conductivity).6Using for the bulk conductivity σ = σ f /F with σ f = (4.9 ± 0.2) × 10–2 S/mat 20◦C.

is 0.6 to 2.3 m thick, (3) a layer of silty/clayey sandy-gravel that is 1.4 to 2.7 m thick, and (4) the impermeableclaystone bedrock of the Laramie formation (bedrock)(Denver Water 1988). Bedrock was encountered beneaththe dam crest in exploratory holes at shallow depths (4.0and 6.9 m). The permeability of the silty/clayey sandy-gravel layer above the bedrock is reported to be two ordersof magnitude greater than that for the bedrock, and oneorder of magnitude greater than the permeability of theoverlying unconsolidated clay sediments (Table 1). Theclayey sandy-gravel layer appears therefore as an aquiferconfined by the overlying and underlying impermeablehorizons. This aquifer has been partially to completelyexposed in the reservoir basin during construction of thereservoir and dam.

Grain size distributions from samples collected inthree exploratory boreholes (see position in Figure 2) willbe used later to parameterize the flow equation duringforward and inverse modeling. The boreholes were com-pleted at the north abutment, the center of the dam crest,and at the south abutment. Samples were available fromdepths between 0.6 and 5.5 m, and represent all of thelayer horizons above the impermeable bedrock. The masspercentages of the coarse and fine fractions are shown inFigure 3 and confirm the hydrostratigraphic units existing

above the bedrock. We assume that the embankment andnatural clay horizons have similar grain size distributionsbecause the embankment was constructed from thenatural clay materials that were excavated from thereservoir basin. Thus, the unconsolidated natural clayhorizon is represented by the grain size distributions thatcorrespond to depths less than 3.0 m that have a muchgreater fines fraction than samples from greater depths.The mean coarse fraction (sand plus gravel) of thesesamples is 32 ± 10%, and the mean fines fraction (clayplus silt) is 68 ± 10%. The samples collected at depthsgreater than 3.0 m represent the sandy-gravel aquifer andhave significantly increased coarse fraction. The meancoarse fraction of these samples is 82 ± 10% by mass.

We used hydrometer analyses of samples collectedbetween 1.8 and 3.0 m to determine the percentage ofsilt and clay present in the fines fractions shown inFigure 3. The resulting mass fraction gradation of theembankment and natural clay horizons consisted of 0%gravel and organic matter, ∼32% sand, ∼24% silt, and∼44% clay. The mass fraction gradation determined forthe sandy gravel layer underlying the embankment andnatural clay horizons is ∼50% gravel, ∼32% sand, and∼18% clay. Soil-water retention parameters for a twoparameter van Genuchten-Mualem model (van Genuchten1980) were estimated from these data by entering themass fraction distributions of each horizon into the SoilPlant Air Water (SPAW) database described by Saxtonand Rawls (2006). The resulting soil-water retentioncurves were then fitted by iteratively adjusting the modelparameters. The Saxton and Rawls (2006) database alsoprovided statistical estimates of bulk density and saturatedhydraulic conductivity which were used to compute theporosity and intrinsic permeability for each horizon abovethe impermeable bedrock. The bulk densities comparewell with the 94% of maximum dry density estimated

Figure 3. Mass fraction constituents of samples collectedin boreholes determined from grain size distributions. Thelayer below 3.0 m corresponds to the confined aquifer. Thepermeability of the aquifer is one order of magnitudegreater than the overlying natural clay and three orders ofmagnitude greater than the underlying bedrock aquitard.

240 S.J. Ikard et al Groundwater 53, no. 2: 238–250 NGWA.org

earlier. The parameters of the bedrock (rich in smectite)were estimated by assuming a permeability that wasthree orders of magnitude less than the clayey sandy-gravel aquifer. All of the geotechnical values describedpreviously are summarized in Table 1.

Anomalous SeepageAnomalous seepage was not observed or detected

through the cross-section or foundation of the dam priorto the survey. However, a localized zone of surface soilmounding is present between 150 and 180 m downstreamof the dam, with significantly greater soil moisturecontent than surrounding topography (see localizationin Figure 2). Furthermore, groundwater discharge hasbeen observed in this zone and correlated with the peakreservoir levels.

Hydrogeophysical InvestigationsThe primary goal of our study was to investigate the

hydraulic connection between the impounded reservoirand the downstream seepage discharge zone, and to pro-vide a plausible explanation for why the dam has showna long record of successful performance without develop-ing anomalous seepage through its core and foundation.The P-wave velocities and electric conductivities increasein each successive layer at depth, and the self-potentialis directly sensitive to groundwater flow, therefore thesemethods are ideal for such an investigation. The sur-vey layout for seismic, self-potential, and ERT profilesis shown in Figure 2.

Seismic P-Wave TomographyTwo 2D seismic profiles, S1 and S2, were collected

with a 24 channel SeimicSource DAQ Link III usinga 0.25 ms sample interval and twenty four vertical axis4.5 Hz center-frequency geophones. Seismic profile S1was established along the dam center line perpendicularto the crest inline with ERT profile P7 (Figure 2), andseismic profile S2 was established at the toe parallel to thecrest between ERT profiles P2 and P3. Geophones werespaced 2 m along these profiles, and a seismic source wastraversed every 4 m along the profiles and provided bydropping a 5.4 kg (12 lb) sledgehammer onto an aluminumplate from an elevation 1.5 m above the surface. Seismicsources were positioned with zero-offset from a givenreceiver location for each shot. Travel times of head wavesrefracted from the phreatic surface were observed at eachstation and inverted to recover the 2D distribution ofP-wave velocity in the subsurface beneath each of theprofiles.

Example raw seismic shot gathers are shown inFigure 4a and 4c with the interpreted first arrivalsindicated on these figures as red lines. An interpretationof observed shot gathers for profile S2 is also included inFigure 4b. The travel-time data were preprocessed priorto performing the inversion, as described below. Firstarrival times of the P-wave wave energy in each shotgather were selected using the built-in module in the

Vscope data acquisition software that was used duringthe acquisition. A floating pretrigger delay was usedduring data acquisition, and the relative “time-zero” foreach shot gather was determined by the onset of seismicenergy at the zero-offset receiver for each shot. Thisarrival time was then subtracted from each pick for agiven shot gather to obtain relative arrival times for usein tomographic modeling. 2D travel-time tomography ofthe P-wave velocity distribution was performed with theSeisOPT2D software which employs a synthetic annealingalgorithm in conjunction with a stochastic Monte-Carloprocedure to find a velocity model that yields a globalminimum value of the root-mean-square error betweenthe observed travel-time data and forward calculated data.Velocity constraints were not applied to the model duringthe inversion procedure, however, the resultant modeledvelocities are well within reasonable values for the presenthydrogeological setting, offering validity significant to thefinal obtained velocity model. The inversion results of thetwo profiles S1 and S2 are shown in Figures 5 and 6.

Self-Potential DataA larger survey consisting of 2D ERT and self-

potential profiles was completed at the field site betweenMarch and June 2011, the same time period when seismicrefraction data was collected. The survey consisted ofseven 2D profiles collected on the dam crest, downstreamslope, and the downstream topography. Crest-parallelprofiles P1-P6 began south of the dam centerline and wereterminated adjacent to the left (north) abutment. Data werecollected along profile P5 at two different times duringthe survey period and were offset by one 16 takeout cablereel between each survey. Crest-perpendicular profile P7began at the reservoir surface contact on the upstreamslope at the dam centerline, and extended 280 m outinto the downstream topography. The electrode spacingalong all profiles was 5 m. Profiles P1-P3 were offsethorizontally in a downstream direction by 10 m, andprofiles P4-P6 east of the access road were offset by 30 m.

A total of 778 self-potential stations were measured atthe field site using two nonpolarizing Pb-PbCl2 electrodesand a handheld Fluke 289 true RMS digital multimeter.A reference electrode was buried in an excavated pit nearthe right abutment of the dam (Figure 2) and assigneda potential of 0 V. All measured self-potential data wereadjusted to be relative to this reference electrode. Ateach station, a shallow hole was excavated to expose soilmoisture and reduce the contact resistance between theelectrode and the ground. The minimum and maximumresistances for the survey were 1 and 130 k� (measuredwith the voltmeter), respectively, and the mean resistancefor all stations was 20 k�, well below the internalimpedance of the voltmeter (100 M�). The potential dif-ference between the reference and roving electrodes wasmeasured before and after acquiring data along each 2Dprofile, and the electrode drift was computed and removedfrom each profile during processing. The electrical con-ductivity of the reservoir water was measured during thesurvey and was σ f = 457 μS/cm–1 (0.046 S/m) at 25◦C.

NGWA.org S.J. Ikard et al Groundwater 53, no. 2: 238–250 241

(a)

(b)

(c)

Figure 4. Seismic shot gathers. (a) Raw shot gather for seismic profile 1 collected along the dam crest with P-wave arrivaltime picks plotted as red lines. First arrival picks were used for the inverse tomographic modeling. (b) Graphic interpretationof a shot gathered from profile S2. (c) Shot gather for seismic profile S1.

DC Resistivity DataResistivity data were collected using an ABEM Ter-

rameter 4000 LUND imaging system with a Wenner-64array protocol. The measured resistances along eachprofile were used to produce 2D pseudo-sections of theapparent resistivity distribution beneath each profile.Pseudo-sections were inverted in 2D with RES2DINVusing a finite-element approach (Loke and Barker 1996).The topography has been taken into account in theinversion but is negligible in the parallel profiles. Thetopography is accounted for in the inversion of the per-pendicular profile and is included in the inverted results.

Interpretation of the Geophysical Data

Seismic P-Wave TomographyThe inverted P-wave velocity models and the corre-

sponding ray-path coverage are shown in Figures 5 and6. Warm colors (yellows and reds) indicate high veloc-ity zones and dense ray-path coverage, while cool col-ors (greens and blues) represent low velocity zones andreduced ray-path coverage. Because of dense source loca-tions along each line, the ray-path coverage is relativelydense, and the inverted velocities are relatively well con-strained within the model space. Piezometer locations and

242 S.J. Ikard et al Groundwater 53, no. 2: 238–250 NGWA.org

(a)

(b)

Figure 5. P-wave velocity inversion results for seismic profile 1 (S1) collected parallel to the crest along the downstream toeof the dam. (a) P-wave velocity distribution. Color scale indicates seismic velocity. (b) Ray-path coverage density plot showinga relative sensitivity distribution. Color scale represents the number of rays that intersect a given model element or pixel.The intersections with ERT profile P7 and seismic profile S2 are indicated at the top of the tomogram.

the interpreted depths to the water table are shown onthe figures as well. The interpreted position of approx-imately 95 to 98% water saturation within the capillaryfringe (black-dotted line) has been interpolated betweenthe piezometers and within the dam based on the measuredwater table depths and the inverted velocity distributions.It is worth noting that the seismic refraction method issensitive to the level of water saturation, and can there-fore be used to image the capillary fringe (Bardet andSayed 1993). P-wave velocities become sensitive to sat-uration levels above approximately 95%, and this can beseen in the tomograms as a sudden increase in veloc-ity within the immediate vicinity of the water table asmeasured in piezometers along the seismic profiles. Theinterpreted depth to greater than 95% saturation increasessharply in the downstream half of the dam, correspond-ing to flow conditions in good agreement with the strongnegative self-potential anomaly observed beneath the creston profile P7 (Figures 7 and 8) (see Titov et al. 2005, fora modeling of this effect). The velocity of the saturatedclayey sandy-gravel aquifer (indicated by the solid-blacklines in Figures 5 and 6) is much higher than that of theoverlying unconsolidated natural clay sediments, and isin good agreement with typically assumed P-wave veloc-ity ranges for saturated shales and clay sediments (1100to 2500 m/s). The capillary fringe is resolved above thephreatic surface in the velocity range 600 to 1100 m/s. The

velocity distribution beneath the dam crest seems homo-geneous in the embankment above the phreatic surface,and shows little to no lateral variations in the natural claysediments below the water table. The observed increase inP-wave velocity with depth is the result of increased satu-ration and decrease in porosity, from unconsolidated claysto more compact aquifer sediments into the compactedbedrock, and the depths of the sharp velocity interfacesobserved in the tomograms are in good agreement withthe depths of these layers interpreted from the grain sizedistributions and ERT (Figures 7 and 8).

Resistivity and Self-Potential DataFigures 7 and 8 show 2D electrical resistivity

tomograms collected across the crest and perpendicularto the crest at the dam centerline, and the correspondingself-potential data along each profile. The resistivitytomograms show each of the layers observed in thegeotechnical boreholes and P-wave velocity distributionsat depth, and show an increased thickness of the uppernatural clay layer due to the presence of the embankmentmaterials, which are assumed to be a slightly disturbedversion of the natural clay horizon as explained above.The subsurface electric conductivity increases with depth.The embankment dam and unconsolidated natural claylayers have resistivities greater than 50 �m. The highlyconductive, impermeable, claystone bedrock aquitard is

NGWA.org S.J. Ikard et al Groundwater 53, no. 2: 238–250 243

(a)

(b)

Figure 6. P-wave velocity inversion results for seismic profile S2 normal to the dam crest. (a) P-wave velocity distribution.(b) Ray-path coverage density plot, showing a relative sensitivity distribution. The color scale of the ray-path coverage plotrepresents the number of rays that intersect a given model element or pixel. Here, the approximate embankment/foundationcontact is plotted as a black-dashed line, and the intersections with ERT profiles P1, P2, and P3, and seismic profile S1are indicated at the top of the tomogram. Along both profiles S1 and S2, the black-dotted line indicates the depth at whichthe saturation levels within the capillary fringe begin to affect the P-wave velocity (saturation ∼95%). The solid-black lineindicates the interpreted phreatic surface. The average 2011 water table elevations from piezometer data are plotted as whitetriangles/lines. The maximum and minimum water elevations recorded during 2011 are plotted as black triangles.

characterized by low resistivities (less than 10 �m). TheERT data suggest that the aquitard surface is undulatorybeneath the dam crest (see Figure 7) in a crest-paralleldirection and approaches the surface downstream ofthe dam centerline along profile P7 (see Figure 8)perpendicular to the crest.

A general methodology to interpret self-potentialsignals can be found in Jardani et al. (2007) and thereader would find details regarding the underlyingphysics in Titov et al. (2002) and Revil et al. (2012). Theself-potential data are sensitive to groundwater flow in thesubsurface, and the self-potential anomalies are thereforesignificantly spatially correlated with the bedrock surfacetopography that is observed in the 2D ERT profiles.Self-potential data observed along the dam crest are pre-dominantly negative with respect to the reference station

and indicate a relatively strong component of ground-water seepage oriented vertically downward beneath andthrough the dam cross-section (see Figures 7 and 9).About 93% of the 315 self-potential stations measuredalong the dam crest showed self-potential readings lessthan 0 mV relative to the reference electrode. In contrast,the self-potentials observed downstream of the dam alongprofile P7 are primarily positive relative to the referenceelectrode, indicating that groundwater flow is directedpredominantly upward towards the surface. About 97% ofself-potential measurements along profile P7 are greaterthan 0 mV relative to the reference electrode, and theonly negative self-potential anomaly observed along thisprofile corresponds to stations measured on the dam crestand downstream slope. The increasing and decreasingtrends in the self-potential data corresponding to profile

244 S.J. Ikard et al Groundwater 53, no. 2: 238–250 NGWA.org

Iteration 5, RME error 6.0%

Elecrical resistivity (Ohm m)

1810

1800

1790

1780

1770

1760

1750

Electrical resistivity tomogram

Self-potential profile P1

Dam and natural clay material - Aquifer

Claystone bedrock - aquitard

20

15

10

5

0

-20

-15

-10

-5E

leva

tion

(m)

Self

-pot

entia

l sig

nals

(m

V)

Distance along the profile (m)

SOUTH NORTH

Reference potential

P7

(b)

(a)

0.0 50

5.0 10 20 50 100

100 150 200 250 300

Figure 7. Electrical resistivity tomography and self-potential measured on the crest (profile P1, taken at the crest of the dam).(a) Self-potential data on the crest were predominantly negative with respect to the reference electrode and were showingvery small spatial fluctuations with respect to those shown in Figure 8. (b) Electrical resistivity tomogram across the crest.The aquifer-aquitard boundary corresponds to the dash line.

P7 are spatially correlated with the bedrock peaksand troughs observed in the corresponding ERT data,respectively, and, excluding the first station (measured inthe reservoir water) and the positive anomaly observed inthe three stations at the eastern-most end of the profile,the largest positive self-potential anomalies are positionedover the two bedrock peaks in the observed seepage zone150 to 180 m downstream of the dam. For each bedrockpeak observed in this zone, the self-potential signalbecomes more positive (in a downstream direction) onthe upstream side indicating groundwater is being chan-neled towards the surface on the upstream slope of theseepage zone, and then decreases (also in a downstreamdirection) on the downstream slope as the groundwater isbeing channeled vertically downward to greater depths.Some of the groundwater in this vicinity intersects thesurface and results in the observed seepage at the groundsurface, and a significant volume of the groundwater inthis zone is also bifurcated around these bedrock peaks,as indicated by the strong negative self-potential anomalyA8 (see Figure 9a).

The self-potential data are shown in relation tothe surface and bedrock topography in Figure 9 andillustrate the high degree of correlation of positiveand negative anomalies with peaks and troughs of thebedrock surface. Negative anomalies are shown in shades

of green and blue and are representative of flow zones,but do not necessarily reflect vertical flow. The negativeanomalies tend to be positioned over bedrock troughs thatcreate continuous preferential flow channels through thesubsurface. The most negative self-potential anomaliesare shown along the dam crest (anomaly A1), and at thesouth end of profile five (anomaly A8). These negativeanomalies are associated with flow through the bedrocktrough beneath and parallel to the dam crest (anomalyA1), and the zone of groundwater bifurcation around thebedrock mounds observed in profile P7 (anomaly A8).Additional negative, albeit lesser amplitude, anomaliesare also observed downstream of the dam and providean indication of preferential groundwater flow channelsthrough the aquifer that are formed by the bedrocktroughs. Anomaly A2 at the north abutment is a reflectionof flow that is being channeled through a bedrock trough.Anomalies A4, A5, and A6 reflect the same, and areinterpreted to result from groundwater flow through acontinuous path that is formed by a bedrock trough andintersects the bedrock slope at the southern-most end ofprofile P5 that produces anomaly A8.

The self-potential anomalies are predominantly pos-itive downstream of the dam, and in contrast to thenegative anomalies, the positive anomalies indicate thatmuch of the groundwater flow beneath the dam is being

NGWA.org S.J. Ikard et al Groundwater 53, no. 2: 238–250 245

Iteration 5, RME error 5.1%

Elecrical resistivity (Ohm m)

Electrical resistivity tomogram

Aquifer

Claystone bedrock - aquitard

Ele

vatio

n (m

)

Distance along the profile (m)

1800

1790

1780

1770

1760

1750

Seepage and surface moundingDamcrest

Self-potential profile P7

Self

-pot

entia

l sig

nals

(m

V)

Positive anomalyDam toe

Infrastructure

30

20

10

0

0.0

5.0 10 20 50 100

40 80 120 160 200 240

-10

-20

Outlier

WEST EAST

Reference potential

P1

(a)

(b)

Figure 8. Profile P7 normal to the dam crest and intersecting the seepage zone 150 m downstream of the dam. (a) Self-potentialprofile. The positive anomaly at the west end of the profile was measured at the contact between the upstream dam slope andthe reservoir. A negative self-potential anomaly is present beneath the dam crest as reservoir water is channeled through theclayey-gravel aquitard below, and increases at the dam toe as seepage is channeled upward to a local bedrock plateau createdwhen the reservoir basin was constructed. (b) Resistivity tomogram of Profile 7. The aquifer-aquitard boundary correspondsto the dash line.

channeled vertically in a direction towards the surface.The trend of negative self-potential beneath the dam topredominantly positive self-potential downstream of thedam is consistent with the shallowing bedrock surfacegeometry shown in 2D ERT profiles (not shown here).The large positive anomaly A3 occurs in a zone where thebedrock is a shallow plateau that is relatively planar andfree of undulation and troughs. Anomaly A7 is positionedover the downstream discharge zone associated with thebedrock mounds shown in 2D ERT profile P7, and con-sists of two spatially localized, strong positive anomalies.Each individual anomaly is correlated with the one of thebedrock mounds, and both show a reduction in amplitudeimmediately downstream of the associated mound. Thesource of anomaly A9 is currently unknown. This anomalyis in the vicinity of drainage infrastructure at the southeastboundary of the survey area (see Figures 1 and 2), or mayalso result due to bedrock outcropping in the vicinity.

A 3D representation of the bedrock surface in thesurvey area was constructed by kriging the (x , y , z )coordinates that were hand-picked from the bedrocksurfaces shown in the 2D ERT profiles. The self-potential data is displayed over the bedrock surfacetopography Figure 9b to emphasize the anomalies thatare spatially correlated with the greater bedrockpeaks

and troughs. Self-potential anomalies are shown to behighly correlated with the bedrock topography. Negativeself-potential anomalies are correlated with the greatestbedrock surface depths, as well as on the downstreamslopes of bedrock peaks and plateaus, and the positiveself-potential anomalies are generally associated with theupstream slopes of bedrock peaks and on plateaus, wheregroundwater is flowing vertically upwards.

Laboratory InvestigationA sample of sediment was collected from the

reservoir bed at the south abutment during a low poolcondition, shortly after the survey was completed. Asample of reservoir water was also collected to usewith the sediment in quantifying the streaming potentialcoupling coefficient connected the electrical potential tothe hydraulic heads. Data were obtained by installingthe material at the base of water column, and filling thecolumn with the sample of reservoir water. The watercolumn containing the sample was then positioned inanother reservoir containing the water collected from thedam, to establish a hydraulic gradient across the sample.The sediment sample was allowed to achieve equilibriumwith the water in the column and the reservoir for 48 hrsprior to performing the experiment. The hydraulic gradient

246 S.J. Ikard et al Groundwater 53, no. 2: 238–250 NGWA.org

(a)

(b)

Figure 9. Self-potential data (in mV) at the field site. (a)Self-potential data with topography showing the downstreamslope of the dam and downstream topography. A1 to A9represent characteristic anomalies that are discussed in themain text. (b) Self-potential data with bedrock topographydetermined from electrical resistivity tomography.

in the experimental apparatus was varied in time and thepotential difference across the sample was measured as afunction of the hydraulic gradient with two nonpolarizingAg-AgCl electrodes and a Fluke 289 digital multimeter.The streaming potential coupling coefficient C is thekey parameter controlling the amplitude of the self-potential signals for a given head gradient. We foundC =−0.7 ± 0.2 mV/m, a value consistent with the trendbetween C and the pore water conductivity shown byRevil et al. (2003, their Figure 3).

Forward and Inverse Modeling of theSelf-Potential Field

A 2D steady-state numerical model of coupled elec-tric and groundwater flow was constructed in COMSOLMultiphysics 4.3a to model the self-potential data andtheir relation to seepage along profile P7 perpendicular tothe crest at the dam centerline. The geometry and depthsof the bedrock and aquifer layers of the model space wereobtained from ERT data acquired along profile P7. Theself-potential signals that were recorded in the field weremodeled and inverted to recover the 2D distribution of thecausative electrical current sources within the dam andsubsurface. The inverted electrical current sources wereused to quantify the Darcy velocity, as described below

(see Soueid Ahmed et al. 2013, for additional tests of thismethodology).

Computation of the Prior Groundwater Flow ModelFor the hydraulic problem we solve the governing

Richards equation (Richards 1931),

[Ce + SeS]∂H

∂t+ ∇· [−K∇ (H + z)] = 0, (1)

where z is the elevation head (m), H is the totalhydraulic head (m), C e denotes the specific moisturecapacity (in m–1) defined by C e = ∂θ /∂H where θ

is the volumetric water content (dimensionless), S e isthe effective saturation of the medium (dimensionless),related to the relative saturation of the water phase bySe = (

Sw − Srw

)/(1 − Sr

w

)(θ = S wφ where φ represents

the connected porosity of the soil material), S is thestorage coefficient (m–1), and t is time. The hydraulicconductivity K is related to the relative permeabilityk r and the hydraulic conductivity at saturation, K s,by K = k rK s. The effective saturation, specific moisturecapacity, relative permeability, and the volumetric watercontent are defined by (van Genuchten 1980),

Se ={

1/[1 + |αH |n]m

, H < 0

1, H ≥ 0, (2)

Ce =⎧⎨⎩ αm

1−m(φ − θr) S

1me

(1 − S

1me

)m

, H < 0

0, H ≥ 0(3)

kr =

⎧⎪⎨⎪⎩Sle

[1 −

(1 − S

1me

)m]2

, H < 0

1, H ≥ 0(4)

θ ={

θr + Se (φ − θr) , H < 0

φ, H ≥ 0(5)

where θ r denotes the residual volumetric water contentand α, n , m , and l (m = 1 – 1/ n) are material proper-ties. Equation 1 was parameterized with statistically basedestimates of saturated hydraulic conductivity, bulk den-sity, residual and saturated moisture contents, and soil-water retention curves corresponding to a clay texturedsoil. These estimates were obtained by fitting soil-waterretention curves given by the soils database described bySaxton and Rawls (2006).

We simulated the prior groundwater flow model byassigning separate permeability values to each hydro-stratigraphic layer in the model space, and assuming eachlayer was isotropic. The assigned permeability values forthe unconsolidated clay and the clayey sand-gravel aquiferlayers were obtained from Saxton and Rawls (2006) usingthe grain size distribution data available for each layer

NGWA.org S.J. Ikard et al Groundwater 53, no. 2: 238–250 247

(see Figure 3), which are assumed to be reasonably closeto ground truth. The permeability of the bedrock layerwas adopted from prior geotechnical reports that wereprovided by the dam owner. All model parameters arereported in Table 1.

The boundary conditions on the hydraulic componentof the prior model were applied as described below.A constant hydraulic head boundary was applied tothe reservoir and on the upstream slope of the dambelow the maximum reservoir elevation. A constant headequal to 3.3 m was applied to the reservoir. This valuecorresponds to the reservoir depth during the maximumstorage condition, which was in effect during the survey.A linear hydrostatic pressure head profile was applied tothe upstream slope in order to vary the hydraulic headalong this boundary based on the elevation of a givenpoint on the upstream slope relative to the maximumreservoir elevation. A seepage face was applied witha mixed boundary condition over the bedrock mounddownstream of the dam between x = 150 and 180 m,where groundwater seepage has been observed. All othermodel boundaries were assigned a specified flux equalto 0 m3/s. The upstream and downstream edges of themodel space were placed sufficiently far away from thedam so that the zero-flux boundaries assigned at the edgeswould not influence the simulated groundwater flow in thevicinity of the dam or the seepage zone.

Computation of the Self-Potential FieldThe self-potential ϕ (in V) is governed by a Poisson

equation (Jardani et al. 2007),

∇· (σ∇ϕ) = ∇· (QV u), (6)

which is obtained by combining the generalized Ohm’slaw, including the advective drag of the excess electricalcharges per unit volume of pore water in the diffuselayer (coating the surface of the minerals), QV (effectiveexcess charge density dragged by the flow of the porewater and expressed in C/m), and the Darcy (seepage)velocity u (in m/s). In Equation 6, σ (in S/m) denotesthe electrical conductivity of the porous material. Theright-hand side of Equation 6 corresponds to the self-potential source term associated with the Darcy velocitydistribution and accounts for the heterogeneities in thedistribution of the volumetric charge density QV . Thischarge density QV is the effective volumetric chargedensity due to the electrical diffuse layer coating themineral/pore-water interface and that is dragged alongwith the flow of the pore water (Jardani et al. 2007). Therelationship between the volumetric charge density andthe more classical streaming potential coupling coefficientC (in V/Pa) described in section “Resistivity and Self-Potential Data” is C = −QV kρ/ηf where ρ = 1/σ isthe electrical resistivity of the porous material (in �m)and ηf denotes the dynamic viscosity of the porewater (in Pa s) (Jardani et al. 2007). For pH ∼5to 8, Jardani et al. (2007) found that the empiricalrelationship log10 QV = −9.2 − 0.82 log10 k (QV in C/m3

and k in m2) holds for a broad range of porous rocksand soils.

The electrical component of the model consisted ofinsulation and electrical ground boundaries. An insulationboundary condition was applied to the reservoir bed,the upstream slope of the dam, and the downstreamtopography including the seepage face that was definedin the hydraulic model. An electrical ground boundarycondition was applied to all other model boundaries.

Inversion of the Self-Potential DataWe use the algorithm described in Jardani et al.

(2008) to perform the inversion of self-potential data from48 stations along Profile P7. The remaining data betweenstations 49 and 57 were excluded from the inverse modelfor two reasons. First, the spatial extent of the aquiferand bedrock horizons were obtained from ERT data alongProfile P7, and are not resolved in the Profile P7 soundingbeyond station 48 (see Figure 8). The aquifer geometryand groundwater flow patterns in this region cannot beaccurately modeled or estimated as a result. Second, thereis a large positive self-potential anomaly in this regionat the easternmost end of Profile P7 (anomaly A9, seeFigure 9a and 9b).

We determined the optimal value of the regularizationparameter by the L-curve criterion. The results of the2D inversion are shown in Figure 10. Figure 10acompares the observed and simulated data. The datamisfit least-square error between observed and simulateddata is 47.7 mV. Most of the data misfit is accountedfor by the large negative self-potential anomaly beneaththe dam, which is predominantly a result of seepageinside of the aquifer beneath the dam foundation. Theamplitude of this anomaly is roughly 25 mV over adistance of approximately 7 m, and it therefore accountsfor increased uncertainty relative to the more broadlydistributed anomalies, and therefore a majority of the totaldata misfit. Figure 10b shows the inverted Darcy velocitymodel giving the spatial distribution and orientation ofthe Darcy velocity vector at each model node on a grid ofnodes spaced 5 m in the horizontal direction and 1 m in thevertical direction. The Darcy velocity in each model layerwas determined from the inverted source current densityvectors and the associated excess volumetric chargedensities given in Table 1. The lengths of the arrowsare proportional to the Darcy velocity (in m/s). Theirdirections imply the predominant direction of groundwaterflow along the profile, and the color scale indicatesthe magnitude. The results of the inverse modelingcorroborate the overall interpretations of the geophysicaldata and show that groundwater flow along the 2D profileoccurs primarily within the confined aquifer. Seepage isguided along the aquifer-bedrock interface beneath thefoundation of the dam in a downstream direction whereit converges toward the surface topography between 150and 180 m where surface seepage has been previouslyobserved.

The Darcy velocity of the groundwater in the clayeysand-gravel aquifer was estimated from the inverse model.

248 S.J. Ikard et al Groundwater 53, no. 2: 238–250 NGWA.org

(a)

(b)

Figure 10. Results of 2D inversion for source current density vector. (a) Observed data vs. data simulated with the inverted2D model. The RMS data misfit is 47.7 mV. (b) Inverted model showing the 2D distribution of source current density vector inthe subsurface. The inverted model shows that flow is primarily beneath the dam in the confined aquifer layer, and convergestoward the topographic surface downstream of the dam in the location of observed topographic mounding, and groundwaterseepage. For an effective charge density of 2.9 C/m3, the mean and maximum inverted velocities in the clayey-gravel aquiferare 6.5 × 10–6 and 2.4 × 10–4 m/s, respectively. The aquifer/bedrock interface is drawn from the resistivity data.

The mean and maximum computed Darcy velocities ofthe groundwater within the aquifer beneath profile P7are 6.5 × 10–6 and 2.4 × 10–4 m/s, respectively, using avolumetric charge density value of 2.9 C/m for the aquifermaterials. The greatest flow velocities are observedbetween x = 80 m where the cross-section thickness ofthe aquifer layer appears to constrict, and x = 170 m wherethe elevation of the bedrock plateau is maximum along theprofile. The maximum velocity is spatially correlated withthe upstream slope of the first bedrock peak, at x = 130 malong the profile. The results of the inverse model supportthe conclusion that the flow occurs primarily through therelatively thin, confined aquifer and that the directionand velocity are predominantly controlled by the bedrocktopography.

ConclusionsWe performed a hydrogeophysical survey of an

earthen dam with a long record of successful operation andno observed anomalous seepage through the foundation orabutments. Our survey combined seismic P-wave tomog-raphy, electrical resistivity tomography, and self-potential.The geophysical data and borehole granulometric data col-lected during the construction of the dam point out a fourlayer hydrogeologic system comprised from top to bottomof (1) the embankment material, (2) unconsolidated nat-ural clay sediments, (3) a confined clayey-gravel aquifer,and (4) an impermeable bedrock (aquitard) (claystone ofthe Laramie formation). We mapped the bedrock surface

topography using ERT and investigated the flow charac-teristics with the self-potential data.

Our data show that a thin confined aquifer acts asa preferential flow pathway channeling the groundwaterthrough the subsurface along the bedrock contact, therebyreducing the propensity for seepage to enter the damcross-section and foundation. The clayey gravel aquiferappears to accept water directly from the reservoir basinchanneling groundwater beneath embankment foundation.The preferential flow directions are controlled by thetopography of the bedrock aquitard. The polarity of theobserved self-potential anomalies is correlated to thedepth-to-bedrock. The topographic mound and seepagezone 150 m downstream of the dam centerline is aresult of the bedrock topography, which nearly intersectsthe surface in this zone. Groundwater in the aquifer ischanneled along the bedrock toward the surface in thisvicinity and manifests as ponded water at the groundsurface during peak reservoir storage conditions.

Given the physical size of the dam relative to thesize of the reservoir, and the duration of its operation, theabsence of anomalous seepage through the dam cross-section and foundation is astonishing. Indeed, the damslong history of success is a result of the hydrogeologyat the site. The surface topography forms a natural basinthat is ideal for a reservoir, and the clayey-gravel aquiferoffers a path of least resistance for the water stored in thereservoir which reduces the potential for anomalous seep-age to form in the dam cross-section and foundation. Thedischarge zone downstream of the dam does not appear to

NGWA.org S.J. Ikard et al Groundwater 53, no. 2: 238–250 249

directly influence the development of anomalous seepagein the dam or its foundation, and there currently does notappear to be a significant potential for this seepage zoneto undermine the successful operation of the dam.

AcknowledgmentsThis work is supported by the NSF-funded projects

SmartGeo (IGERT: Intelligent Geosystems; DGE-0801692) and PIRE. We thank Denver Water forlogistical support and site access, K. Titov and ananonymous referee for their very constructive reviews.

ReferencesBardet, J.P., and H. Sayed. 1993. Velocity and attenuation

of compressional waves in nearly saturated soils. SoilDynamics and Earthquake Engineering 12: 391–420.

Boleve, A., F. Janod, A. Revil, A. Lafon, and J. Fry. 2011.Localization and quantification of leakages in dams usingtime-lapse self-potential measurements associated with salttracer injection. Journal of Hydrology 403: 242–252.

Butler, D., J.L. Llopis, T.L. Dobecki, M.J. Wilt, R.F. Corwin,and G. Olhoeft. 1990. Part 2: comprehensive geophysicalinvestigation of an existing dam foundation: engineeringgeophysics research and development. Geophysics 9. DOI:10.1190/1.1439782.

Denver Water. 1988. Design report and geotechnical informationfor Lower Long Lake Dam. Prepared by Dam SafetySection, Denver Water.

Fell, R., C.F. Wan, J. Cyganiewicz, and M. Foster. 2003. Timefor development of internal erosion and piping in embank-ment dams. Journal of Geotechnical and GeoenvironmentalEngineering 129, no. 4: 307–314.

Foster, M.A., R. Fell, R. Davidson, and C.F. Wan. 2002.Estimation of the probability of failure of embankmentdams by internal erosion and piping using event treemethods. Australian National Committee on Large Dams(ANCOLD) Bulletin 121, Melbourne, Australia.

Foster, M., R. Fell, and M. Spannagle. 2000. The statisticsof embankment dam failures and accidents. CanadianGeotechnical Journal 37, no. 5: 1000–1024.

van Genuchten, M.T. 1980. A closed-form equation for pre-dicting the hydraulic conductivity of unsaturated soils. SoilScience Society of America Journal 44, no. 5: 892–898.

Ikard, S.J., A. Revil, A. Jardani, W.F. Woodruff, M. Parekh,and M. Mooney. 2012. Saline pulse test monitoring with

the self-potential method to nonintrusively determine thevelocity of the pore water in leaking areas of earth damsand embankments. Water Resources Research 48: W04201.DOI:10.1029/2010WR010247

Jardani, A., A. Revil, A. Boleve, and J.P. Dupont. 2008.3D inversion of self-potential data used to constrain thepattern of ground water flow in geothermal fields. Journalof Geophysical Research 113: B09204. DOI:10.1029/2007JB005302

Jardani, A., A. Revil, A. Boleve, A. Crespy, J.P. Dupont, W.Barrash, and B. Malama. 2007. Tomography of the Darcyvelocity from self-potential measurements. GeophysicalResearch Letters 34: L24403.

Loke, M.H., and R.D. Barker. 1996. Rapid least-squaresinversion of apparent resistivity pseudosections using aquasi-Newton method. Geophysical Prospecting 44, no. 1:131–152.

Panthulu, T.V., C. Krishnaiah, and J.M. Shirke. 2001. Detectionof seepage paths in earth dams using self-potential andelectrical resistivity methods. Engineering Geology 59, no.3–4: 281–295.

Revil, A., M. Karaoulis, T. Johnson, and A. Kemna. 2012.Review: some low-frequency electrical methods for sub-surface characterization and monitoring in hydrogeology.Hydrogeology Journal 20, no. 4: 617–658. DOI:10.1007/s10040-011-0819-x

Revil, A., V. Naudet, J. Nouzaret, and M. Pessel. 2003. Princi-ples of electrography applied to self-potential electrokineticsources and hydrogeological applications. Water ResourcesResearch , 39, 5: 1114. DOI:10.1029/2001WR000916

Richards, L.A. 1931. Capillary conduction of liquids throughporous media. Physics 1: 318–333.

Rozycki, A., J.M.R. Fonticiella, and A. Cuadra. 2006. Detectionand evaluation of horizontal fractures in earth dams usingself-potential method. Engineering Geology 82: 145–153.

Saxton, K.E., and J.W. Rawls. 2006. Soil water characteristicestimates by texture and organic matter for hydrologicsolutions. Soil Science Society of America Journal 70, no.5: 1569–1578.

Soueid Ahmed, A., A. Jardani, A. Revil, and J.P. Dupont. 2013.SP2DINV: A 2D forward and inverse code for streaming-potential problems. Computers & Geosciences 59: 9–16.

Titov, K., A. Levitski, P. Konosavski, A. Tarasov, Yu. Ilyin, andM. Bues. 2005. Combined application of surface geoelectri-cal methods for groundwater flow modeling: A case history.Geophysics 70: H21–H31. DOI:10.1190/1.2076747

Titov, K., Yu. Ilyin, P. Konosavski, and A. Levitski. 2002.Electrokinetic spontaneous polarization in porous media:Petrophysics and numerical modelling. Journal of Hydrol-ogy 267: 207–216.

250 S.J. Ikard et al Groundwater 53, no. 2: 238–250 NGWA.org