&EPA
United States
Environmental Protection
Agency
Office of Research and
Development
Washington DC 20460
EPA/600/R-97/023
July 1997
Investigation of
Hydrogeologic Mapping to
Delineate Protection Zones
Around Springs
Report of Two Case Studies
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EPA/600/R-97/023
July 1997
INVESTIGATION OF HYDROGEOLOGIC MAPPING TO DELINEATE
PROTECTION ZONES AROUND SPRINGS
REPORT OF TWO CASE STUDIES
by
Mark E. Jensen
Utah Department of Environmental Quality
Salt Lake City, Utah 84114-4830
Mike Lowe
Utah Geological Survey
Salt Lake City, Utah 84109-1491
Michael Wireman
U.S. Environmental Protection Agency
Denver, Colorado 80202-2405
Cooperative Agreement
CR-817607
Project Officer
Steven D. Acree
Subsurface Protection and Remediation Division
National Risk Management Research Laboratory
Ada, Oklahoma 74820
National Risk Management Research Laboratory
Office of Research and Development
U.S. Environmental Protection Agency
Cincinnati, OH 45268 j
Printed on Recycled Paper
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NOTICE
The U.S. Environmental Protection Agency through its Office of Research and Development partially
funded and collaborated in the research described here under Cooperative Agreement CR-817607 to the
Utah Department of Environmental Quality. It has been subjected to the Agency's peer and administrative
review and has been approved for publication as an EPA document. The document contains copyrighted
material on pages 30 and 49. Mention of trade names or commercial products does not constitute endorse-
ment or recommendation for use.
All research projects funded by the U.S. Environmental Protection Agency that make conclusions or
recommendations based on environmentally related measurements are required to participate in the Agency
Quality Assurance Program. This project was conducted under an approved Quality Assurance Project Plan
and the procedures therein specified were used with the following exceptions. Geophysical logging of
boreholes specified in the plan was not performed; following initial investigation at the field sites, it was
determined that other techniques for acquiring the necessary information were more appropriate. In addi-
tion, the final springs chosen for study were different from those originally proposed and instrumentation
for measuring water temperature in the field was changed. Information, on the plan and documentation of
the quality assurance activities and results are available from the Principal Investigator.
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FOREWORD
The U.S. Environmental Protection Agency is charged by Congress with protecting the Nation's land,
air, and water resources. Under a mandate of national environmental laws, the Agency strives to formulate
and implement actions leading to a compatible balance between human activities and the ability of natural
systems to support and nurture life. To meet these mandates, EPA's research program is providing data and
technical support for solving environmental problems today and building a science knowledge base neces-
sary to manage our ecological resources wisely, understand how pollutants affect our health, and prevent or
reduce environmental risks in the future.
The National Risk Management Research Laboratory is the Agency's center for investigation of tech-
nological and management approaches for reducing risks from threats to human health and the environment.
The focus of the Laboratory's research program is on methods for the prevention and control of pollution to
air, land, water, and subsurface resources; protection of water quality in public water systems; remediation
of contaminated sites and ground water; and prevention and control of indoor air pollution. The goal of this
research effort is to catalyze development and implementation of innovative, cost-effective environmental
technologies; develop scientific and engineering information needed by EPA to support regulatory and
policy decisions; and provide technical support and information transfer to ensure effective implementation
of environmental regulations and strategies.
This report presents a discussion of hydrogeologic characterization tools that can be used to evaluate
potential zones of contribution to springs used in public water supplies and potential protection zones around
the springs. Techniques described in this document provide a cost-effective means for obtaining basic
information regarding hydrogeology and are applicable at many sites. It is published and made available by
EPA's Office of Research and Development to assist the user community.
Clinton W. Hall, Director
Subsurface Protection and Remediation Division
National Risk Management Research Laboratory
in
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ABSTRACT
Methods commonly used to delineate protection zones for water-supply wells are often not directly
applicable for springs. This investigation focuses on the use of hydrogeologic mapping methods to identify
physical and hydrologic features that control ground-water flow to springs to aid in delineating springhead
protection zones. Two public-supply springs were selected as study sites to represent diverse geologic
settings. One spring discharges from fractured dolomite and one from fractured siltstone, sandstone, and
shale.
Hydrogeologic mapping techniques, as applied in this study, are methods for mapping geologic or
hydrologic features or geophysical/geochemical signatures of subsurface features. These data are often
supplemented with information from subsurface investigations to aid in extrapolating surface results to
aquifer depths. In this investigation, geologic mapping, fracture-trace analysis, topographic analysis, catch-
ment area estimation, geochemical characterization, elemental isotope studies, and a tracer study were used
to locate and describe potential ground-water flow boundaries and pathways and develop conceptual mod-
els for site hydrogeology. Results were integrated to estimate the zones of contribution to each spring and
evaluated for use in the delineation of potential protection zones. Data from borings and hydraulic tests
supported this characterization by providing direct and indirect information regarding subsurface lithology
and hydraulic parameters.
In these case studies, results of hydrogeologic mapping allowed development of the conceptual
model for site hydrology and evaluation of potential ground-water flow controls. Although definitive ground-
water flow boundaries suitable for delineation of practicable protection zones could not be identified at
either site, these techniques provided sufficient information to support an initial evaluation of potential
protection zones. These results would also serve as a strong basis for additional investigations if more
detailed or reliable delineations were warranted.
This report was submitted in fulfillment of CR-817607 by the Utah Department of Environmental
Quality under the partial sponsorship of the U.S. Environmental Protection Agency. This report covers a
period from November 1990 to May 1993, and work was completed as of May 1993.
IV
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CONTENTS
i
Notice j ii
Foreword , iii
Abstract iv
Figures vii
Tables viii
Acknowledgements ....ix
Chapter 1. Introduction 1
Background 1 1
Hydrogeology of Springs ; 3
Objectives and Methodology 1 4
Wellhead Protection Criteria and Thresholds . 6
Chapter 2. Hydrogeologic Mapping Methods for Protection Zone Delineation 7
Surface Mapping of Hydrogeologic Features 7
Geologic Mapping • 7
Fracture-Trace Analysis 7
Topographic and Geomorphic Analysis 8
Catchment Area Estimation 8
Tracer Studies 8
Geochemical Characterization 8
Isotope Studies 9
Potentiometric Surface Mapping > 10
Geophysical Techniques 10
Methods Used to Support Hydrogeologic Mapping.; ; 10
Chapter 3. Results of Hydrogeologic Mapping < 12
Olsens Spring 12
Location and Description 12
Climate . 13
Previous Hydrogeologic Investigations 14
Geologic Setting [ 14
Surface Mapping of Hydrogeologic Features 14
Borings and Well Installations : 18
Physical and Chemical Properties of Discharge 20
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Field Parameters 20
Geochemical Characterization 21
Saturation Indices and Calcium/Magnesium Molar Ratios 25
Catchment Area Calculations 29
Tracer Study ". 29
Hydrogeologic Conceptual Model 31
Aquifer Tests 32
Zone of Contribution and Delineation of Spring Protection Zones 34
Sheep Spring 34
Location and Description 34
Climate 36
Previous Hydrogeologic Investigations 36
Geologic and Hydrologic Setting 36
Surface Mapping of Hydrogeologic Features 39
Borings and Well Installations 42
Physical and Chemical Properties of Discharge 44
Field Parameters 44
Geochemical Characterization 46
Catchment Area Calculations 47
Hydrogeologic Conceptual Model 48
Aquifer Tests 49
Zone of Contribution and Delineation of Spring Protection Zones 50
Chapter 4. Conclusions and Recommendations : 53
Evaluation of Hydrogologic Mapping Methods 53
Guidance for Application of Techniques 54
References 56
VI
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FIGURES
Number
1 Principal types of springs based on hydrogeologic characteristics 5
2 Location of Olsens Spring 12
3 Geologic map of Olsens Spring study area 15
4 Geologic cross section through Olsens Spring study area 17
5 Joint orientations in Olsens Spring study area i 17
6 Rose diagram depicting strike of joints in the surface drainage area surrounding
Olsens Spring ;' 18
7 Location of monitoring wells at Olsens Spring 19
8 Geologic and construction logs of monitoring wells at Olsens Spring 20
9 Combined monthly flow of seven springs in Mantua Valley 22
10 Temperature of water samples from Olsens Spring and mean monthly
air temperatures at Brigham City 23
11 Specific conductance and pH of water samples from Olsens Spring 23
12 Trilinear diagram depicting geochemical analyses of water samples from
Olsens Spring _ 26
13 Seasonal variation of calcite saturation index in water samples from Olsens Spring 28
14 Estimated catchment areas for Olsens Spring and combined Olsens/West
Mailings Springs 39
15 Instrumental intensity of eluted samples from Olsens/West Hallings Springs,
Olsens Spring, and Mud Spring : 31
16 Recovery data from aquifer tests at Olsens Spring , 33
17 Surface drainage basin surrounding Olsens Spring '. 35
18 Location of Sheep Spring study area 35
19 Geologic map of Sheep Spring study area 37
20 Potential zone of contribution to Sheep Spring and other springs in this area 38
21 Joint orientations in Sheep Spring study area 40
22 Rose diagram depicting strike of joints in Sheep Spring study area 40
23 Geologic cross section through Sheep Spring study area ...... 41
24 Location of monitoring wells installed at Sheep Spring....: 43
25 Construction logs of wells near Sheep Spring and representative geologic
log of Well 1 43
26 Temporal trend in water levels in wells at Sheep Spring.... : 44
27 Temperature of water samples from Sheep Spring and mean monthly
air temperature at St. George 45
28 Specific conductance and pH of water samples from Sheep Spring 46
29 Trilinear diagram depicting geochemical analyses of water samples from Sheep Spring... 48
30 Estimated catchment area for Sheep Spring 49
31 Recovery data from slug tests in wells near Sheep Spring ., 51
32 Example of potential protection zones for Sheep Spring based on preliminary
time-of-travel data 52
vii
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TABLES
Number
Page
1 Springs used for public water supply in the United States 2
2 Classification of springs based on average discharge rate 3
3 Location and elevation of Olsens Spring, West Hallings Spring, and associated
monitoring wells 13
4 Climatic data from weather stations near Olsens Spring 13
5 Ground-water elevations and estimated hydraulic gradients near Olsens Spring 21
6 Field parameters of water samples from Olsens Spring 22
7 Analytical results of water samples from Olsens Spring 24
8 Analytical results of samples from West Hallings Spring and two monitoring wells
near Olsens Spring 24
9 Analytical parameters and laboratory methods 25
10 Calcite and dolomite saturation indices calculated for water samples from Olsens
Spring 27
11 Calcum/magnesium molar ratios calculated for water samples from
West Hallings Spring and monitoring wells 28
12 Interpretation methods and transmissivity estimated from aquifer tests at
Olsens Spring 34
13 Location and elevation of Sheep Spring and monitoring wells 36
14 Field parameters of water samples from Sheep Spring 45
15 Analytical results of water samples from Sheep Spring 47
16 Hydraulic conductivity estimated from slug tests in wells near Sheep Spring 51
VIII
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ACKNOWLEDGMENTS
This study was conducted in cooperation with the Utah Geological Survey and the U.S. Geological
Survey, Water Resources Division, Salt Lake City, Utah. We thank Geoffrey W. Freethey, Larry Spangler,
and Joseph S. Gates of the U.S. Geological Survey, for their field work, numerous discussions, and reviews;
Wayne C. Downs, Robin R. Roberson, and Kris Jensen of the U.S. Environmental Protection Agency for
their efforts in compiling data regarding the use of springs in public water supplies; Gary Christenson, Utah
Geological Survey, for helpful review; and personnel of the Brigham City and Santa Clara city water depart-.
ments.
IX
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Chapter 1
INTRODUCTION
BACKGROUND
Throughout the course of human history, more interest has been generated in springs than in any other
ground-water feature (Davis and De Wiest, 1966). Historically, springs have determined the location of human
settlements in arid regions. Many cities and towns in the western United States are named for springs. For
centuries, hot springs and mineral springs were thought to be of medicinal or therapeutic value. Bottled "spring"
water is still considered by many people to be of higher quality than ordinary tap water. This belief is at least
partially responsible for the current high volume sales of bottled water. Unfortunately, this association of spring
water with purity is not based on fact. Spring water can contain higher concentrations of dissolved solids than
local public water supplies and many types of springs are more easily contaminated than other sources of drink-
ing water.
Over 3400 public water-supply systems in the United States are known to obtain part or all of their
drinking water from springs (Table 1). These systems provide drinking water for more than seven million
people. However, the average number of people served by a single public-supply system utilizing springs as a
water source is small. Public water supplies that utilize springs are more numerous in the western United States.
More than 200 public systems use springs in each of the states of Colorado, Idaho, Oregon, Utah, and Washing-
ton.
Springs are an important component of the water supply in many states. There are several states in which
more than 100,000 people are served by public systems that include springs. In the state of California, more than
four million people (14.2% of 1990 population) receive their drinking water from systems with springs. In both
Tennessee and Pennsylvania, more than 500,000 people receive their drinking water from such systems. In each
of the states of California, Hawaii, Tennessee, and Vermont, more than 10% of the population is served by
systems that include springs.
The Wellhead Protection Program was authorized by the 1986 Amendments to the Safe Drinking Water
Act. This program is designed to assist states in protecting areas around wells (or springs) within their jurisdic-
tion from contaminants that may have adverse effects on human health (Safe Drinking Water Act, section 1428[a]).
One of the six critical elements of a wellhead protection program is delineation of a scientifically valid protec-
tion area around public ground-water sources.
The U.S. Environmental Protection Agency (U.S. EPA) is charged with providing technical guidance to
state and local personnel on delineation of wellhead protection areas (WHPAs). As part of this effort, the U.S.
EPA has published several guidance documents including "Guidelines for Delineation of Wellhead Protection
Areas" (U.S. EPA, 1987), "Surface Geophysical Techniques for Aquifer and Wellhead Protection Area Delinea-
tion" (Violette, 1987), "Wellhead Protection Strategies for Confined-Aquifer Settings" (Kreitler and Senger,
1991), and "Model Assessment for Delineating Wellhead Protection Areas" (van der Heijde and Beljin, 1988).
These publications focus primarily on delineation of WHPAs around wells in granular porous-media aquifers.
The U.S. EPA has also published "Delineation of Wellhead Protection Areas in Fractured Rocks" (Wisconsin
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Table 1. Approximate Number of Springs Used for Public Water Supply in the United States (Various State
Personnel, Personal Communication to M. Wireman, 1992)
State
Alaska
Arkansas
California
Colorado
Connecticut
Georgia
Hawaii
Idaho
Illinois
Kansas
Maine
Maryland
Massachusetts
Michigan
Montana
Nebraska
Nevada
New Hampshire
New Jersey
New Mexico
New York '
North Carolina
North Dakota
Ohio
Oklahoma
Oregon
Pennsylvania
Rhode Island
South Dakota
Tennessee
Texas
Utah
Vermont
Virginia
Washington
West Virginia
Wisconsin
Approximate
No. of Systems
15
13
106
226
11
31
67
329
4
18
137
48
15
1
200
2
49
25
10
61
100
5
2
2
10
600
302
3
13
75
20
325
189
123
262
40
20
Approximate
Pop. Served
4,000
42,000
4,238,688
74,284
5,813
239,210
124,000
58,382
15,625
20,000
32,756
44,333
5,318
1,000
15,000
644
21,171
1,000
500
49,172
25,000
250
100
48,300
38,870
45,000
518,443
495
62,910
614,470
50,000
86,954
63,581
233,566
250,000
83,703
500
Total Pop. 1990 Census 5
550,000
2,351,000
29,760,000
3,294,000
3,287,000
6,478,000
1,108,000
1,007,000
11,431,000
2,478,000
1,228,000
4,781,000
6,016,000
9,295,000
799,000
1,578,000
1,202,000
1,109,000
7,730,000
1,515,000
17,990,000
6,629,000
639,000
10,847,000
3,146,000
2,842,000
11,882,000
1,003,000
696,000
4,877,000
16,987,000
1,723,000
563,000
6,187,000
4,867,000
1,793,000
4,892,000
V0 of Total Pop.
0.7
1.8
14.2
2.3
0.2
3.7
11.2
5.8
0.1
0.8
2.7
0.9
<0.1
<0.1
1.9
<0.1
1.8
<0.1
<0.1
3.2
0.1
<0.1
<0.1
0.4
1.2
1.6
4.4
<0.1
9.0
12.6
0.3
5.0
11.3
3.8
5.1
4.7
<0.1
TOTALS:
3,459
7,115,038
194,560,000
3.7%
Note: States that are not listed in this table either made no use of springs in public water supplies or had no
compiled statistics regarding springs at the time of data collection. Data were obtained from personnel in each
state and are considered to be estimates of actual usage.
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Geological and Natural History Survey, 1991), which focuses on delineation of protection areas for public wells
developed in fractured-rock aquifers.
Hydrogeologic methods commonly used to delineate protection zones around water-supply wells are of-
ten not directly applicable to springs. Aquifer tests generally cannot be conducted using springs alone. How-
ever, such tests have been reported (e.g., Clarke, 1989). Tests to estimate hydraulic parameters of the aquifer
supplying water to a spring usually require the use of nearby wells. Many public-supply springs are in remote
areas where there are no monitoring or production wells nearby, and very little aquifer data are available.
Hydrogeologic mapping techniques (e.g., geologic mapping, fracture-trace analysis, geochemical character-
ization, etc.) offer relatively low-cost methods for obtaining aquifer characterization data.
HYDROGEOLOGY OF SPRINGS
A spring is best defined as a concentrated discharge of ground water appearing at the ground surface as a
current of flowing water (Todd, 1980). Springs occur in many forms. Bryan (1919) divided all springs into
(1) those resulting from gravitational forces and (2) those resulting from non-gravitational forces. Gravity
springs result from ground water flowing under hydrostatic pressure. Non-gravitational springs include those
associated with volcanism and fractures that extend to great depths in the earth's crust. Most non-gravitational
springs are thermal springs. The vast majority of springs are gravity springs.
Springs can be classified in a number of ways. Characteristics used to classify springs include discharge
rate, type of aquifer (geology and structure), water quality, variability, and type of openings through which the
water issues. The most common classification of springs, based on discharge rate, was developed by Meinzer
(1923). This classification (Table 2) divides springs into eight orders of magnitude based on average discharge
rate.
Table 2. Classification of Springs Based on Average Discharge Rate (Meinzer, 1923)
Magnitude
Average Discharge
First
Second
Third
Fourth
Fifth
Sixth
Seventh
Eighth
>170,000 l/min (> 100.0 fP/s)
17,000 l/min -170,000 l/min (10.0 fP/s -100.0 fP/s)
1700 l/min -17,000 l/min (1.0 fP/s - 10.0 fP/s)
380 l/min -1700 l/min (100 gal/min -1.0 fP/s)
38 l/min - 380 l/min (10 gal/min -100 gal/min)
3.8 l/min - 38 l/min (1 gal/min -10 gal/min)
0.5 l/min - 3.8 l/min (1 pint/min -1 gal/min)
< 0.5 l/min (< 1 pint/min)
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The discharge rate of a spring is a function of three main variables - the hydraulic conductivity of the
aquifer, the area contributing recharge to the aquifer, and the quantity of recharge (Davis and De Wiest, 1966).
Based on the above classification there are only a few hundred springs of first magnitude in the world and about
fifty in the United States. Most large magnitude springs discharge from volcanic rocks, limestones, or boulder/
gravel aquifers.
Springs are also commonly classified based on geologic characteristics. The Water Resources Division of
the United States Geological Survey recognizes eight principal types of springs based on hydrogeologic charac-
teristics (Figure 1):
1. Artesian spring - release of pressurized water from a confined aquifer at the aquifer outcrop or
through an opening in the confining unit.
2. Contact spring - occurs where a permeable water-bearing unit overlies a less permeable unit that
intersects the ground surface.
3. Depression spring - forms where the ground surface intersects a water table.
4. Fracture spring - occurs where ground water flows predominantly through fractures that intercept the
ground surface.
5. Geyser spring - periodic thermal spring resulting from expansive force of superheated steam within
constricted subsurface channels (Todd, 1980).
6. Perched spring - occurs where infiltrating water discharges above the regional water table from a
permeable geologic unit that overlies a less permeable geologic unit.
7. Seep spring - discharges from numerous small openings in permeable material. These springs
typically have a very low discharge rate.
8. Tubular spring - discharges from rounded channels (karst solution openings, lava tubes).
The above classification is used by the U.S. Geological Survey for entering spring data into the National
Water Data Storage and Retrieval System (Baker and Foulk, 1975). Other spring data in this national water data
base include spring name, location, permanence, sphere of discharge, development or improvement data, and
flow variability. This type of hydrogeologic spring classification is important as an aid to hydrogeologic map-
ping for purposes of delineating ground-water protection zones around springs. These classifications form the
basis for a conceptual model of site hydrogeology. Each classification implies the potential existence of certain
physical and hydrologic controls on ground-water flow. Knowledge of potential controls allows identification
of appropriate site characterization techniques. The investigation may then be designed to test the assumptions
of the conceptual model.
OBJECTIVES AND METHODOLOGY
The objectives of this cooperative research were investigation of hydrogeologic mapping techniques to
determine the zone of contribution to springs and evaluation of this information for delineation of protection
zones. The project was designed as a field study conducted at two sites in different hydrogeologic settings.
Hydrogeologic mapping was chosen as the primary focus for this project because of its potential applicability to
many hydrogeologic settings that include strongly anisotropic aquifers, such as fractured bedrock (U.S. EPA,
1987). Additional considerations included the relatively low technological requirements and implementation
costs for many of these methods. Based on these studies, a general methodology for applying such techniques is
proposed. The report is designed to aid investigators involved in planning characterization studies leading to the
establishment of protection zones around springs. The project was limited in scope. Field studies were con-
ducted at only two sites in the arid southwest region of the United States. Although general conclusions regard-
4
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1. Artesian
5. Geyser
2. Contact
6. Perched
3. Depression
7. Seep or filtration
4. Fracture
7=d^Uy. ? \
8. Tubular
Figure 1. Principal types of springs based on hydrogeologic characteristics (Baker and Foulk, 1975).
Numbers are keyed to text descriptions. Arrows point to spring locations.
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ing utility of basic hydrogeologic characterization methods are applicable to delineation of protection zones in
most settings, the utility of specific methods will be highly dependent on site characteristics.
Hydrogeologic mapping, as applied in this project, refers predominantly to geologic, hydrologic, geochemi-
cal, and geophysical techniques for characterizing subsurface features using the surface expression and geo-
physical or geochemical signatures of such features. These techniques include:
• fracture-trace analysis,
• analysis of land surface topography,
• geologic mapping,
• potentiometric surface mapping,
• geophysical surveys,
• tracer studies,
• stable and radioactive isotope studies, and
• geochemical characterization.
Depending on site conditions, ground-water flow boundaries and pathways may be identified and mapped in
detail. A discussion of these methods and potential applicability is provided in Chapter 2.
The methodology applied in this study began with a review of site-specific literature to develop an initial
conceptual model for site hydrogeology. Based on this model, potential ground-water flow controls were identi-
fied. Hydrogeologic mapping methods were chosen to test assumptions of the conceptual model and, poten-
tially, delineate physical and hydrologic flow controls. Results were interpreted to refine the hydrogeologic
conceptual model and estimate the zone of contribution to the springs. These results were then evaluated for use
in delineating potential protection zones. Monitoring wells were installed at each site to provide subsurface
information and allow preliminary hydraulic testing. Results of subsurface investigations were compared with
indirect information inferred from the hydrogeologic mapping.
WELLHEAD PROTECTION CRITERIA AND THRESHOLDS
The delineation criteria and thresholds specified in the Drinking Water Source Protection (DWSP) Pro-
gram for the State of Utah were used in this investigation as an example of state requirements in drinking water
source protection programs. Approximately 785 springs are used in Utah as public water-supply sources. All
sources of ground water for public supplies, including springs and tunnel discharges, are subject to the provi-
sions of the wellhead protection program in Utah.
The Utah DWSP Program specifies three protection zones, with higher levels of management in the zones
closest to the ground-water source. Zone 1 (closest to each source) is based on arbitrary distance. Zones 2 and 3
are based on ground-water time-of-travel criteria or flow boundaries. The three drinking-water source protec-
tion zones defined in Utah are:
Zone 1 - the area within a 100-foot (30 m) fixed-radius zone around the well or spring;
Zone 2 - the capture zone area within the 250-day ground-water time-of-travel to the well or
spring, the boundary of the aquifer(s) that supplies water to the ground-water source, or
the ground-water divide, whichever is closer; and
Zone 3 - the capture zone area within the 15-year ground-water time-of-travel to the well or
spring, the boundary of the aquifer(s) that supplies water to the ground-water source, or
the ground-water divide, whichever is closer.
The applicability of the information inferred from mapping to practicable protection zone delineation was evaluated
at each study site using these management criteria.
6
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Chapter 2
HYDROGEOLOGIC MAPPING METHODS FOR
PROTECTION ZONE DELINEATION
The following is a general discussion of hydrogeologic mapping techniques and is not limited to those
methods chosen for use at the study sites. These methods may be used to delineate ground-water flow bound-
aries, identify geologic features that affect ground-water velocity in many settings, and assess vulnerability of an
area to ground-water contamination from surface sources. Potential physical and hydrologic controls on ground-
water flow to springs include:
1. geologic units that act as flow boundaries due to low hydraulic conductivity,
2. physical limits on geologic formations that serve as aquifers,
3. physical properties of the aquifer materials (e.g., hydraulic conductivity),
4. fractures or karst features that act as preferential flow paths,
5. faults that may act as flow paths or flow boundaries,
6. ground-water divides, and
7. recharge and discharge locations/characteristics, including well pumping and irrigation.
Techniques that are useful in characterizing the hydrogeologic setting and the zone of contribution to springs are
described below.
SURFACE MAPPING OF HYDROGEOLOGIC FEATURES
Geologic Mapping
Geologic mapping can be used to identify surface and subsurface lithologic units and their stratigraphic
relationship, map geologic structure (faults, folds, and fractures), and map the spatial orientation of these fea-
tures. Geologic mapping techniques are explained in more detail in Compton (1962) and Barnes (1981). In
hydrogeologic studies, geologic mapping may often be used to:
1) infer the extent of lithologic units that provide water to springs,
2) identify low-permeability lithologic units that may form ground-water flow boundaries,
3) identify faults that may act as either barriers or preferential paths for ground-water flow,
4) characterize fracture orientations or karst features that may control ground-water flow, and
5) identify potential recharge areas.
Geologic mapping may yield sufficient information to indicate the zone of contribution to a spring through
identification and delineation of formations that behave as aquifers and aquitards. The potential utility of such
mapping depends, in part, on the degree to which the lithologic units of interest are exposed and the hydrogeologic
setting.
Fracture-Trace Analysis
Fracture traces and lineaments are natural linear features expressed on the surface of the earth. These
features are commonly caused by steeply dipping faults or joint systems (Billings, 1972). The terms fracture
trace and lineament have been differentiated on the basis of length (Fetter, 1988; Lattman, 1958). Fracture
traces are between 300 m and approximately 1300 m (1000 ft to 4300 ft) in length and lineaments are longer than
1300 m (4300 ft). Fracture traces may be identified on aerial photographs or topographic maps by straight
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stream segments, aligned erosional features, low spots in ridges, vegetation patterns, or other visually linear
features (Lattman, 1958). Major fracture systems or faults, expressed as fracture traces and lineaments, may
control ground-water movement, acting as preferential flow paths or flow boundaries. Fracture-trace analysis
may be used to locate features for more detailed study during geologic mapping and other investigations. Such
analysis generally is conducted as a preliminary step in characterization of fractured or faulted terrains.
Topographic and Geomorphic Analysis
Analyses of topographic maps may be used to locate the boundaries of surface drainage basins surround-
ing springs. In shallow, water-table aquifers, ground-water divides often coincide with drainage-basin divides
(Fetter, 1988). Therefore, delineation of topographic divides surrounding springs is useful in developing the
conceptual model for potential zones of contribution. Geomorphic analysis uses techniques of both topographic
analysis and geologic mapping. Geomorphic principles (Bloom, 1978) may be applied to infer subsurface
structure from present landforms. For example, the potential presence of subsurface karst features may be
inferred from landforms such as depressions indicative of sinkholes.
CATCHMENT AREA ESTIMATION
In most hydrogeologic settings, the rate of discharge of a spring is related to the area contributing recharge
to the aquifer and recharge rate. Using estimates of recharge rate and the discharge rate from the spring, the area
of the topographic catchment basin required to provide the spring discharge can be estimated. This information
may aid in determining if the zone of contribution to a spring is potentially larger than the topographic basin
surrounding the spring. It should be recognized that detailed recharge information usually is difficult to acquire
and the recharge estimate will generally incorporate a significant degree of uncertainty. Thus, estimates of
required catchment areas may also include significant uncertainty. However, such techniques have been suc-
cessfully used in many settings, including hydrologically ungauged catchments (Bonacci and Magdalenic, 1993).
TRACER STUDIES
Tracer studies are often used to identify specific ground-water flow paths and estimate ground-water
velocity to discharge. Such investigations usually are performed by introducing a tracer into an aquifer through
a well or sinkhole. Many inorganic and organic compounds have been used as tracers in previous studies.
Applicable compounds depend on many site-specific conditions such as ambient ground-water quality, study
objectives, and local regulations. Selected wells and/or springs are monitored to determine the discharge point(s),
time of travel, and recoverable quantity of the tracer. In karst settings and other areas where ground-water flow
occurs predominantly through a few discrete pathways, tracer studies may be the only mean's for reliable estima-
tion of the zone of contribution to springs. Additional discussion regarding tracer studies is provided in numer-
ous texts and journals (e.g., Davis and others, 1985; and Mull and others, 1988).
GEOCHEMICAL CHARACTERIZATION
The chemical composition of ground water is influenced by many factors as the water is recharged and
flows through an aquifer. These factors include water-rock interaction along the flow path, composition of the
recharging precipitation, water-soil-air interaction in the unsaturated zone, and biologic activity (Hem, 1985).
-------
Analyses that may be useful indicators of aquifer processes include concentrations of major ions. Classification
using such systems as the hydrochemical facies of Back (1966) and the calculation of pertinent mineral satura-
tion indices (e.g., calcite, dolomite, and aragonite) may be valuable tools for inferring aquifer lithology and
relative ground-water residence times. Temporal fluctuations in these parameters may indicate variations in
ground-water flow paths or flow rates and residence time (Shuster and White, 1971; Quinlan and others, 1991).
Temporal changes in physical/chemical parameters (e.g., temperature, pH, specific conductance, turbidity, and
discharge rate) measured in the field may also indicate changes in ground-water flow path or residence time and
may aid in defining the zone of contribution by indicating connections between ground water and surface water
(e.g., Quinlan and others, 1991). Daily or weekly fluctuations in water temperature or other parameters .may
indicate a shallow aquifer, rapid recharge, recharge in proximity to the spring, or conduit flow conditions.
Fluctuations in pH, specific conductance, turbidity, or flow may also be indicative of changes in recharge vol-
ume, area, or rates.
ISOTOPE STUDIES
Elemental isotopes may be used to estimate ground-water residence time, characterize recharge distribu-
tion, and trace ground-water flow from recharge to discharge areas. Several radioactive and stable isotopes have
been used in such investigations (e.g., Hendry, 1988; Mazor and Nativ, 1992; Solomon and Sudicky, 1991;
Taylor and others, 1992). Two common radioactive isotopes used to date ground water are tritium (3H) and
carbon-14 (14C).
Tritium occurs naturally in the atmosphere, has a half-life of approximately 12.4 years, and is carried into
ground water by recharging precipitation. Tritium was introduced into the atmosphere, and, thus, into ground
water, at increased levels by atmospheric testing of nuclear weapons between 1952 and 1969 (Drever, 1988). As
the timing of this significant increase in atmospheric tritium is known, tritium can be used in a qualitative way
to estimate the average residence time of ground water.
Before atmospheric testing of nuclear weapons began in 1952, the natural tritium activity in rainwater was
approximately 10 tritium units (TU). One TU equals one tritium atom per 1018 hydrogen atoms. Ground water
containing less than about 2 TU to 4 TU entered the ground-water system prior to 1953. If the tritium content is
significantly greater than 10 TU to 20 TU, the ground water has been recharged or in contact with the atmo-
sphere since 1953 (e.g., Drever, 1988; Fetter, 1988). A more detailed analysis of average ground-water age
based on tritium concentrations is also presented by Hendry (1988). However, site-specific data on tritium
activity in precipitation since nuclear testing began is necessary for a more quantitative age assessment. Recent
studies (e.g., Solomon and others, 1992) have used analyses of tritium and tritiogenic helium-3 (3He) to over-
come this limitation in certain situations and accurately date ground water with ages between 0 and 50 years.
Ground water of similar age may also be dated using concentrations of anthropogenic chlorofluorocarbon com-
pounds (Busenberg and Plummer, 1992; Szabo and others, 1996).
Carbon-14 also occurs naturally in the atmosphere and was introduced in elevated concentrations by
atmospheric nuclear testing. Carbon-14 has a much longer half-life (about 5730 years) than tritium, so it can be
used to date waters ranging in age from several thousand years to a few tens of thousands of years (Drever, 1988;
Hendry, 1988). Difficulties in determining the initial carbon-14 concentration and separating the influence of
carbon contributed from aquifer material may limit reliability of this technique. Other isotopes that have been
-------
used to assess the residence time of ground water and trace flow paths from recharge to discharge include the
isotopic ratios of oxygen (18O/16O), hydrogen (WH), chlorine-36 (36C1), and helium-4 (4He) (Hendry, 1988;
Marty and others, 1993; Mazor and Nativ, 1992).
Information regarding relative ground-water residence time and flow paths derived from isotopic studies
may be useful in evaluating potential protection area delineations (Wisconsin Geological and Natural History
Survey, 1991). This information may also be used to support or refute relative ages determined from hydraulic
calculations (Mazor and Nativ, 1992). However, such determinations yield the average age of dissolved species.
Mixing of ground waters of different ages may result in an isotopic age that differs significantly from the
average ground-water travel time (Solomon and others, 1992). The applicability of these methods depends on
the hydrogeologic setting and the results may be qualitative or ambiguous (Wisconsin Geological and Natural
History Survey, 1991).
POTENTIOMETRIC SURFACE MAPPING
Potentiometric surface mapping is one of the most powerful tools in aquifer studies. Information regard-
ing the hydraulic gradient, potential ground-water flow directions, and ground-water flow boundaries may be
obtained using this tool. Temporal fluctuations in these parameters may also be assessed. However, care must
be exercised in construction of potentiometric maps to ensure data representative of the aquifer in question are
used. Equal care must be exercised in interpretation of this information, particularly in fractured media (e.g.,
Wisconsin Geological and Natural History Survey, 1991). This technique could not be applied in this investiga-
tion due to a lack of monitoring points (e.g., wells, springs, or seeps).
GEOPHYSICAL TECHNIQUES
Many surface geophysical methods have been applied to identify signatures of subsurface features that
may control ground-water flow (e.g., faults, changes in lithology). Potential uses include delineation of subsur-
face lithologic and stratigraphic boundaries, depth to ground water or bedrock, and fracture/solution channel
distribution/orientations. These techniques often complement results of geologic mapping by providing con-
trols on the positions of aquifer and aquitard formations at depth. Applications of such methods are discussed in
numerous texts (e.g., Telford and others, 1981). Surface geophysical techniques and their application to protec-
tion zone delineation are explained in greater detail in Violette (1987). Techniques include seismic methods
(refraction and reflection surveys), electrical methods, ground penetrating radar, and potential field surveys
(e.g., gravity and magnetic methods).
Borehole geophysical techniques may be useful if appropriate boreholes or wells are available. Informa-
tion regarding subsurface lithology, stratigraphy, fracture density and orientation, and other important param-
eters are obtainable using such tools (Keys, 1989). Information from these studies may be used to control
subsurface projections interpreted from mapping of surface features.
METHODS USED TO SUPPORT HYDROGEOLOGIC MAPPING
Hydraulic tests (e.g., pumping and slug tests) were performed and ground-water time-of-travel was esti-
mated to support the hydrogeologic mapping conducted in this investigation. Hydraulic testing provided an
independent estimate of aquifer parameters for evaluating zones of contribution, potential protection zones, and
10
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parameters qualitatively characterized using hydrogeologic mapping. The objective of these tests was to pro-
vide preliminary estimates of hydraulic parameters. The tests were limited in scope and were not designed to
define heterogeneity or anisotropy within the potential protection zones at these sites.
As with many investigations in fractured media (U.S. EPA, 1991), the hydraulic tests were analyzed using
standard solutions and methods based on porous media assumptions. An underlying assumption governing the
use of standard ground-water flow equations to describe fractured bedrock settings is that the aquifer behaves as
a porous medium at the problem scale (Wisconsin Geological and Natural History Survey, 1991). A second
implicit assumption is that aquifer properties can be characterized using methods developed with the porous-
media assumption. Although the spring sites studied in this project may better satisfy the porous-media assump-
tions at the scale of the zone of contribution, it is unlikely that either site behaved as a porous medium at the
scale of the aquifer tests. Thus, results from these tests are considered preliminary estimates and subject to a
relatively high degree of uncertainty. However, these results appear to be sufficient for the qualitative compari-
sons made in this study.
11
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Chapter 3
RESULTS OF HYDROGEOLOGIC MAPPING
Two public-supply springs in Utah were selected as study sites in different hydrogeologic settings: a
fractured carbonate aquifer (Olsens Spring) and a fractured clastic-rock aquifer (Sheep Spring). Both settings
are common in this region. In each case, site-specific literature, including investigations by state and federal
agencies and universities, was first reviewed to develop an initial conceptual model for site hydrogeology.
Based on the conceptual model, potential ground-water flow controls (e.g., dominant fracture systems and low
permeability boundaries) were identified and hydrogeologic mapping methods were chosen for investigation.
OLSENS SPRING
Location and Description
Olsens Spring is located in the Wasatch Range of the Middle Rocky Mountains physiographic province
(Stokes, 1986) of northern Utah (Figure 2 and Table 3). The spring is about 5 km (3 mi) east of the Wasatch
Fault, the eastern boundary of the Basin and Range province. It is located along the northwest margin of Mantua
Valley, at an elevation of approximately 1579 m (5180 ft). It is one of several springs located around the margin
of this small valley. West Ballings Spring is located approximately 275 m (900 ft) southwest of Olsens Spring
within the same surface drainage basin. This study focused on Olsens Spring. However, West Mailings Spring
and other springs in Mantua Valley were considered during these investigations, as appropriate.
Brigham City, which has a population of approximately 17,000, uses Olsens Spring and other springs as
sources of potable water. Olsens Spring was developed as a water supply approximately 70 years ago, and no
OLSENS SPRING
STUDY AREA
BOX ELDER
BRIGHAM
CITY
0
i
50
100 Kilometers
Figure 2. Location of Olsens Spring study area.
12
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Table 3. Location and Elevation of Springs and Associated Monitoring Wells Interpolated from 1:24,000 •
Scale Topographic Maps
Olsens Spring
West Mailings Spring
Mud Spring
Well 1 (Top of casing)
Well 2 (Top of casing)
Latitude
41°30'52"N
41°30'44"N
41°31'08"N
41°30'54"N
41°30'55"N
Longitude
111°56'30"W
111°56'36"W
111°56'05"W
111°56'34"W
111°56'32"W
Elevation
1579m (5180ft)
1579m (5180ft)
1580m (5185ft)
1609m (5280ft)
1608m (5275ft)
plans are available to show how it was developed. The spring appears to have been developed by laying collec-
tion pipes in gravel. The pipes feed concrete spring boxes connected to the city water-supply system.
The collection area is in alluvial fan deposits, within several feet of the contact between the dolomite
bedrock and the fan deposits. The recharge area is mostly hilly and mountainous topography, with several
surface drainages that potentially contribute recharge. Estimated yield of this fourth magnitude spring (Table 2)
is about 17001/min (1.0 ftVs). West Mailings Spring is a third magnitude spring that discharges about 64001/min
(3.8 ft3/s) (Brigham City Corporation, personal communication, 1987) from faulted limestone stratigraphically
lower, but, in the same hill as Olsens Spring. Mud Spring is 790 m (2600 ft) northeast of Olsens Spring, and is
a third magnitude spring discharging a maximum of about 5940 1/min (3.5 ftVs) (Bjorklund and McGreevy,
1973). A canyon separates Mud Spring from Olsens Spring.
Climate
The nearest weather observation station (Table 4) is in Brigham City, about 6.5 km (4 mi) west of Olsens
Spring. Brigham City is 290 m (950 ft) lower in elevation than Olsens Spring and is located in a very large
valley west of the Wellsville Mountains. The weather station at Pineview Dam is in a geographic setting similar
to Olsens Spring and is only 73 m (240 ft) lower in elevation than the spring. Pineview Dam is located about
30 km (19 mi) southeast of Olsens Spring.
Table 4. Climatic Data from Weather Stations Near Olsens Spring (Ashcroft and others, 1992)
Brigham City
Waste Plant
Pineview Dam
Elevation
Average mean temperature
A verage precipitation
Average snowfall
Maximum evapotranspiration
1290 m (4230 ft)
9.7 °C (49.5 °F)
48 cm (18.8 in)
75 cm (29.7 in)
118 cm (46.6 in)
1506m (4940 ft)
6.5 °C (43.7 °F)
79 cm (30.9 in)
301 cm (118.4 in)
110 cm (43.2 in)
13
-------
Average annual precipitation at Brigham City and Pineview Dam is 48 cm (18.8 in) and 79 cm (30.9 in),
respectively. Average annual evapotranspiration of rangeland and mountains at the altitude of Olsens Spring is
estimated to be about 35 cm (13.6 in) (Bjorklund and McGreevy, 1974).
Previous Hydrogeologic Investigations
A thesis describing the hydrogeology and hydrochemistry of several springs in Mantua Valley was com-
pleted by Rice (1987). Olsens Spring was not directly included in that study, but the "West Railings Spring"
samples collected during that investigation were actually composite samples of West Mailings Spring and Olsens
Spring discharges (K. Rice, personal communication to M. Jensen, 1992). Rice (1987) mapped the general
geology of the entire Mantua Valley area, but did not differentiate the individual stratigraphic units. Therefore,
it was necessary to re-map the Olsens Spring area as part of this investigation. The Mantua Valley area is also
included in a report on ground-water resources of the lower Bear River by Bjorklund and McGreevy (1974).
The Olsens Spring study area is included on the Box Elder County geologic map produced by Doelling
(1981). The stratigraphy and structure of Mantua Valley are discussed in a regional geologic study by Williams
(1948). The Mantua Valley area is also included on a surficial geologic map by Personius (1990) and in a soil
survey of eastern Box Elder County conducted by Chadwick and others (1975).
Geologic Setting
Olsens Spring is situated at the base of the Wellsville Mountains. The area around the spring is underlain
by a Paleozoic stratigraphic section of interbedded limestone, dolomite, shale, and quartzite (Hintze, 1988).
These rocks were faulted and fractured by east-directed thrust faulting during the Cretaceous Sevier orogeny.
Eastward tilting and further faulting and fracturing are a result of normal faulting during the last 10 million years
along the Wasatch and related faults (Hintze, 1988).
Surface Mapping of Hydrogeologic Features
Prior to geologic mapping, topographic maps (1:24,000 scale) were studied to identify the topographic
divides of the surface drainage basin surrounding Olsens Spring and West Hallings Spring. Ground-water
divides, which represent hydrologic boundaries to shallow ground-water flow, may be coincident with the topo-
graphic divides. For this reason, the surface drainage basin was defined as the area for detailed study. No
perennial streams or other surface water bodies that may infer the presence of hydrologic boundaries are located
within the basin.
Fracture traces and lineaments were identified on l:20,000-scale aerial photographs in conjunction with
the geologic mapping. The most extensive feature mapped on the aerial photographs was a lineament found to
coincide with a strike-slip fault located about 3 km (2 mi) northwest of Olsens Spring. No fracture traces were
mapped through Olsens Spring or West Hallings Spring. Geologic field mapping was conducted in the surface
drainage area above Olsens Spring. Lithologic contacts, bedding attitude, and structural features (e.g., joint
location and attitude) mapped in the field were plotted on l:20,000-scale black-and-white stereo-pair aerial
photographs. These data were then transferred to a 1:24,000-scale topographic base map.
Results of this mapping (Figure 3) indicate that Olsens Spring discharges from dolomite of the Cambrian-
age Nounan Formation. The Nounan consists of medium to thick-bedded dolomite and sandy dolomite. The
14
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111°58.25'W
I
111°56'W
I
— 41°33'N
\
Geologic
Period
Geologic
Unit
2000 feet
1
500 meters
41°31.5'N
Mud
ipring
Description
Alluvial and colluvial deposits
Alluvial fan deposits
Lake Bonneville deposits -
silt, sand, day, and gravel
Fish Haven and Laketown Dolomites,
undifferentiated dolomite
Swan Peak Formation - quartette
Garden City Formation - lirriestone
St. Charles Formation - dolomite
Nounan Formation - dolomite
Calls Fort Shale Member of the Bloomington Formation
Middle limestone Member of the Bloomington Formation
Hodges Shale Member of the Bloomington Formation
Blacksmith Dolomite
Langston and Ute Formations, undifferentiated -
interbedded limestone and shale
Geertsen Canyon Quartzite
Contact, dashed where approximate
and dotted where inferred
Fault, dashed where approximate,
arrows show apparent, relative movement
Strike and dip of bedding
Boundary of study area
Figure 3. Geologic map of the Olsens Spring study area. Field mapping was performed by M. Jensen and
M. Lowe, with modifications after Williams (1948) and Gelnett (1958). The bedding orientations
presented on this figure are provided for illustration. Some measurements were not posted for
clarity.
15
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Nounan Formation is about 250 m (825 ft) thick in this area. The spring collection area is located approximately
9 m (30 ft) stratigraphically above the contact between the Nounan and the underlying Calls Fort Shale Member
of the Bloomington Formation. The Calls Fort Shale Member consists of well-indurated shale with interbedded
pods of limestone. The contact between the Calls Fort Shale and the Nounan Formation dolomite is a grada-
tional or interbedded contact. Stratigraphically below the Calls Fort Shale is an unnamed middle limestone
member from which West Hallings Spring discharges and a basal shale member similar to the Calls Fort.
Mud Spring is located approximately 800 m (2600 ft) northeast of Olsens Spring, at approximately the
same elevation as Olsens and West Hallings Springs. Mud Spring discharges at the base of Round Hill from
limestone of the Ordovician-age Garden City Formation at the contact with Quaternary-age valley-filling sedi-
ments. The Garden City Formation consists of medium dark gray, finely crystalline limestone, which weathers
medium light gray and forms slopes and ledges. The Garden City Formation in this area contains intraformational
conglomerate, which is an identifying characteristic. This unit is approximately 460 m (1500 ft) thick beneath
Round Hill and contains pods and layers of black chert in the upper portion. The Garden City Formation is
separated from the Nounan Formation at Olsens Spring by the Cambrian-age St. Charles Formation. No fracture
traces or faults were mapped through Mud Spring.
The stratigraphic units near the springs strike about N30°W and dip 45° to 60° to the northeast. Section
A-A' (Figure 4) was chosen to show the eastward-dipping Paleozoic units and minor faults in the potential zone
of contribution. The location of this cross section is shown on Figure 3. Dominant joints near Olsens Spring
strike N28°E to N86°E (Figure 5). Joints throughout the surface drainage area strike predominantly north to
northeast (Figure 6). Joint spacing mapped in outcrops at most locations in this area is generally 0.3 m to 1.2 m
(1 ft to 4 ft). However, joint density near Olsens Spring is higher, with a spacing of about 0.15 m to 0.20 m (0.5
ft to 0.7 ft) in outcrops. Joints were generally traceable across an individual outcrop. However, outcrops
generally were limited to 1 m to 3 m (4 ft to 10 ft) wide exposures and surface exposures were often weathered.
Thus, joint lengths could not be reliably mapped. Joint apertures appeared to range from about 0 to 2 cm (0 to
0.75 in) in diameter. Thin calcite fillings were found in many joints with localized deposits of coarsely crystal-
line calcite.
Several small faults were mapped in the drainage basin (Figure 3). Many of these faults appeared to have
a transverse component of motion. Some of this apparent movement may be the result of dip-slip motion across
dipping beds. Mapped fault lengths within 1000 m (3300 ft) of the springs ranged from approximately 600 m to
1200 m (2000 ft to 4000 ft). The closest fault appears to intersect West Hallings Spring. Faulting was not
identified in the immediate vicinity of Olsens Spring. In addition, no surface expressions of karst features were
observed during the field mapping.
Based on the geologic mapping, it appears that the area surrounding Olsens Spring and West Hallings
Spring is extensively fractured. From these results, several hypotheses regarding the hydrogeologic system may
be formulated:
1. Extensive fracturing in the exposed bedrock may indicate a relatively high degree of vulnerability of the
ground water to'surface contamination sources.
2. The hydrologic system may behave as a porous media on the scale of the surface drainage basin.
However, sufficient data are not available to test this hypothesis.
3. Hydraulic conductivity and, potentially, ground-water velocity may be relatively high in this system.
16
-------
Feet
NW
SE
3000
Meters
— 2000
— 1500
— 1000
0 500 meters
Figure 4. Geologic cross section through the Olsens Spring area. Projections of mapped faults are dashed
Elevations are referenced to mean sea level.
111°58.25'W
111*56W
i \ -41°33'N
x
Contour Interval 200 feet
Datum is sea level.
EXPLANATION
.Strike and dip of
dominant joint set
.Strike and dip of
secondary joint set
itrike and dip of bedding
-Strike of vertical
joint set
./_Strike of dominant joint set -
\X selected from Rice (1987)
Strike of secondary joint set -
selected from Rice (1987)
-Lineament from air photo
"\^ ^_Dot marks location of
•«| A'* joint set
-_. 4 Boundary of study area
Figure 5. Joint pattern map of the Olsens Spring study area. Field mapping was performed by M. Jensen
and M. Lowe, with selected joint measurements from Rice (1987).
17
-------
N
Figure 6. Rose diagram depicting strike of joints in the surface drainage area surrounding Olsens Spring.
Unshaded sectors represent strike of bedding. Number of joint measurements - 12, number of
bedding plane measurements = 34.
4. The system may be anisotropic with an increased potential for ground-water movement to the springs
from the north to east-northeast based on joint orientations.
5. The Calls Fort Shale Member separating Olsens Spring and West Hallings Spring does not appear to be
a boundary to ground-water flow based on the jointing mapped in outcrops of this unit.
6. Ground-water flow to West Hallings Spring may be controlled to some degree by the fault on which it
is located. However, no evidence of significant fault control of ground-water flow to Olsens Spring
could be inferred from geologic mapping. It is also conceivable that these springs are controlled by a
single structural feature not identified in this study.
Borings and Well Installations
Two monitoring wells were installed topographically upgradient from Olsens Spring (Figure 7) to provide
additional subsurface lithologic information and for use in estimating hydraulic parameters. Potential well
locations in this area were limited by the relatively steep topography. Of the locations accessible to the drilling
rig, these sites were chosen on the premise that hydraulic gradients may mimic topographic gradients and that
the formations may be sufficiently fractured to provide flow paths to Olsens Spring or West Hallings Spring.
Well 1 was located approximately 114 m (373 ft) topographically upgradient from Olsens Spring. Well 2 was
located approximately 35.4 m (116 ft) northeast of Well 1. The wells were drilled using air rotary technology.
Well 1 is 44.2 m (145 ft) deep (Figure 8). The upper 33.5 m (110 ft) of the rock strata consists of dolomite
18
-------
(Nounan Formation), with interbedded shale and some limestone in the bottom 10.7 m (35 ft) of the well.
Intervals of fine sand-sized fractions of dolomite were encountered above the water table in both holes, which
may indicate solutional enlargement of fractures. However, it is not known if the dolomitic sand was formed in
place or was deposited in fractures. Well 2 is 38.1 m (125 ft) deep, all in dolomite of the Nounan Formation
(Figure 8). Both borings indicate highly jointed dolomite through most of the drilled intervals. Geophysical
logging of the boreholes, which was originally proposed, was not conducted. Open-hole logging potentially
could have provided information regarding solution channeling and fracture density at depth. However, such
logging would not have been feasible due to a lack of competency of some materials and restrictions on drilling
methods that could be used upgradient of the water-supply springs. Based on the results of geologic mapping,
the stratigraphic information obtainable from the proposed logs was not considered essential at this location.
Sufficient stratigraphic control was available from the samples obtained during drilling.
The wells were constructed using 10.2 cm (4 in) diameter schedule 40 PVC casing and screened immedi-
ately below the water table. An appropriately graded filter pack was installed within the screened zone, a
bentonite seal was placed above the sand pack, and the annulus was grouted to the surface. The location and
elevation of the top of each well casing and the top of the north collection box at Olsens Spring were surveyed
with respect to the top of the south collection box.
Ground-water elevations in the wells and spring boxes were measured to 0.3 cm (0.01 ft) on three occa-
sions following well installation. Continuous monitoring may have provided much valuable information re-
garding site hydrology. However, such monitoring was not feasible at this site due to a lack of security and
Well #2
Well #1
o
100 feet
30 meters
North box
Olsens Spring
South box
West Mailings Spring
275m
Figure 7. Monitoring well locations relative to Olsens Spring.
19
-------
WELL#1 Depth (below
land surface)
WELL #2 Depth (below
land surface)
n
^^z
=
in
Construction
og-symbols
cased
screened
cement
grout
bentonite
silica sand
size 6/9
.;
S;
Wj
1
V, f'.
\ :
Si
s |
gg
— —
—
:
i
V
"i
<1"l
v^.
S :
•i
•\ :
P
20-
30-
-^-
40-
i — n
-7-
"^s3M
^•vT
wVi
1
jjg?i
/
^
/
J=L
I
i_l
i
J L
- 0
dolomite
no cuttings
- 25
rounded pebbles
in dolomitic sand
~ 50 no cuttings
cement
grout
dolomitic sand
rare rounded
- 75 pebbles
do lorn sts
— 1 00 bentonite
silica sand
- 125
intflrhflririfiH limftQtnnpi
and shale
- 14R
\ V
'.Vs
o*
3
,"• s
*?
\
'i
'*
*
<"
2
^
s<
^
11 %'
1
1
^
^
8
'
m
10-
20-
30-
\7
/
/
4
c
/
/
*T
t
t
f
/
l
WC^
^2
^
ft
-25
-50
dolomite
-75
-100
.,OE dolomite
Figure 8. Geologic and construction logs of monitoring wells at Olsens Spring site.
concerns about vadalism of equipment. The hydraulic gradient near Olsens Spring was estimated using ground-
water elevations in the two monitoring wells relative to that in the north spring collection box (Table 5). Based
on the construction of this collection box, it appears that water elevation measured in this box may be represen-
tative of ground-water elevation at the discharge point. However, construction plans for the collection struc-
tures at the spring were not available. Therefore, gradient estimates using this monitoring point may not be
reliable. Calculated maximum gradients ranged from 0.007 to 0.014 for the three sets of measurements that
were obtained. The direction of maximum gradient was toward the south-southeast (152°) and did not vary
significantly. The monitoring wells were installed during the final phase of field studies to allow use of informa-
tion acquired during previous phases in siting the wells. Thus, limited information on temporal fluctuations in
ground-water elevations was obtained.
Physical and Chemical Properties of Discharge
Field Parameters
Temperature, pH, specific conductance, and turbidity of water from Olsens Spring were monitored in the
field on a monthly basis from November 1990 through October 1991 (Table 6). Due to the manner in which
Olsens Spring was developed, it was not possible to monitor the discharge rate at the spring. Records of total
discharge from seven springs in Mantua Valley, including Olsens and West Hallings Springs, were maintained
by Brigham City personnel (Figure 9). These data indicate significant seasonal variation in discharge rate
related to differences in recharge. Although data regarding daily fluctuations in discharge are not maintained,
20
-------
Table 5. Ground-Water Elevations in Wells Relative to Ground-Water Elevations in North Spring Collection
Box and Estimated Hydraulic Gradients Near Olsens Spring
Date Monitoring Point
22-Oct-92
4-Nov-92
17-May-93
Welll
Well 2
Welll
Well 2
Welll
We II 2
Relative Elevation
(m)
0.59
0.65
0.66
0.73
1.24
1.35
(ft)
1.92
2.12
2.17
2.40
4.06
4.42
Gradient
(m/m)
0.007
0.007
0.014
Azimuth (°)
152
152
151
discharge does not appear to noticeably fluctuate daily or following precipitation events (L. Pebley, Brigham
City Corporation, personal communication to M. Jensen, 1993).
The temperature of the water varied from 9.5 °C to 10.6 °C (49.1 °F to 51.1 °F) during the 12-month
sampling period (Figure 10). This small variation did not appear to be significant. Specific conductance did not
change significantly during the year, ranging from 498 (imhos/cm to 528 ^mhos/cm (Figure 11). The turbidity
of the spring water was low, ranging from 0.04 NTU to 0.08 NTU (nephelometric turbidity units). The range of
observed pH values was between 7.54 and 7.82, potentially displaying a slight seasonal trend. However, addi-
tional data would be required to evaluate the existence or significance of such a trend. The lack of distinct
seasonal variations in water temperature, specific conductance, and turbidity indicate that residence time of the
discharge from Olsens Spring is sufficiently long to mask the changes in these parameters due to seasonal
variation in recharge. More frequent measurements (e.g., hourly or daily) of these parameters and discharge
from individual springs would be required to determine whether some component of conduit flow results in
rapid transport of surface water to the spring following precipitation.
Geochemical Characterization
Water samples from Olsens Spring were collected quarterly for one year and analyzed by standard labora-
tory techniques at the Utah State Health Laboratory for major ions, selected trace elements, radiological con-
stituents, and total dissolved solids (Table 7). Water samples from West Hallings Spring and the two monitoring
wells near Olsens Spring were also analyzed (Table 8). The laboratory methods that were used are listed in
Table 9. Samples were collected from the wells for comparison with results from Olsens Spring and West
Hailing Spring in an attempt to identify differences in flow paths to these springs.
Ion balances for the first two samples from Olsens Spring have calculated errors of approximately 0% and
2%, respectively, indicating reliable analyses. Ion balances for the third and fourth analyses display differences
between cations and anions of approximately 16% and 11%, respectively. Calcium and sulfate concentrations in
the third and fourth analyses, respectively, were significantly different from previous values. These anomalous
values may be the result of laboratory error. If the two values are changed to match values of the same constitu-
ents in the other samples, the calculated ion balances for the third and fourth samples would be in error by less
than 5%. Based on the ion balances, and because there are no significant variations in other parameters, the two
indicated values appear to be in error and may be laboratory errors. The total dissolved solids (TDS) concentra-
21
-------
Table 6. Field Parameters for Water Samples from Olsens Spring
Temperature
Date (° C)
29-Nov-90
17-Dec-90
22-Jan-91
12-Feb-91
14-Mar-91
22-Apr-91
22-May-91
13-Jun-91
W-Jul-91
22-Aug-91
18-Sep-91
21-Oct-91
9.5
—
9.6
10.1
10.2
10.1
10.3
10.4
10.3
10.4
10.2
10.6
Specific
Conductance
famhos/cm)
498
499
517
506
504
523
505
507
500
528
521
523
Total Dissolved
Solids
(mg/1)
258
249
259
254
252
261
253
254
250
265
261
262
pH
7.76
7.82
7.70
7.66
7.66
7.61
7.57
7.54
7.57
7.57
7.71
7.72
Turbidity
(NTU)
0.07
0.04
0.06
0.04
0.05
0.08
0.04
0.04
0.04
0.04
0.04
1400 r
1
o
I
CD
.g
1,
a
en
£2
Q
1200
1000
800
600
I I I I I I I I I I I I I I I I I I I I I I I I I I I I I I I I I I
AJO JAJ OJAJ O
AP U C|AP U C
R L T
IA
IN
1990
1991
1992
Figure 9. Combined monthly flow of seven springs in Mantua Valley. E = estimated flow, due to malfunction
of flow recorder. (B. Bingham and L. Pebley, Brigham City Corporation, personal communication
to M.Jensen, 1993.)
22
-------
Temperature (°C)
— *• i — 1. — i. M N_
0010 OIO 01 O O
^"®,
>' \
/
*
• *
X
1 • I 1 1 I lift
' ' -«'
\ /
\ /
» ^ «___„=_=
» * Spring Temp. -^ Mean Air Temp.
I I i i I 1 i 1 I I I i
Nov Dec Jan Feb Mar Apr May Jun Jul Aug Sep Oct
1990 199-1
Figure 10. Temperature of water samples from Olsens Spring and mean monthly air temperature in Brigham
City (various personnel, Utah Climate Center, personal communication to M. Jensen, 1993).
1000
900
8
o
800
700
600
400
I I I
Specific Conductance ^pH
I i
I i
10
Q.
4
Nov Dec Jan Feb Mar Apr May Jun Jul Aug Sep Oct
1990 1991
Figure 11. Specific conductance andpH of water samples from Olsens Spring.
23
-------
Table 7. Analytical Results for Water Samples from Olsens Spring
Sample Date
17-Dec-90
14-Mar-91
13-Jun-91
18-Sep-91
Calcium (mg/l)
Magnesium (mg/l)
Potassium (mg/l)
Sodium (mg/l)
Bicarbonate(mg/l)
Carbonate (mg/l)
Chloride (mg/l)
Silica (mg/l)
Sulfate (mg/l)
N!trate+Nitrite(mg/l)
Phosphate(mg/l)
Iron (mg/l)
Alkalinity(mg/l)
Specific Conductancefamhos/cm)
Total Dissolved Solids (mg/l)
Gross Alpha (pc/l)
Gross Beta (pc/l)
Cations (meq/1)
Anions (meq/1)
51
26
<1
15
268
0
25
9
10
1.25
0.02
<0.02
220
492
370'
<1 •
<5
5.4
5.4
53
27
<1
15
269
0
26
9.8
10
1.40
0.34
<0.02
221
486
268
3
<5
5.5
5.4
33'
26
<1
16
257
0
28
9.4
9.6
1.39
0.02
<0.02
211
492
274
<1-
<5
4.5
5.2
50
26
<1
17
267
0
29
9.4
37.4"
1.31
0.01
<0.02
219
509
282
<1
<5
5.4
6.0
' These values appear to be errors.
Table 8. Analytical Results for Water Samples from West Mailings Spring and Monitoring Wells Near Olsens
Spring
West Mailings Spring
Sample Date 28-May-92
Calcium (mg/l)
Magnesium (mg/l)
Potassium (mg/l)
Sodium (mg/l)
Bicarbonate(mg/l)
Carbonate (mg/l)
Chloride (mg/l)
Silica (mg/l)
Sulfate (mg/l)
Nitrate+Nitrite (mg/l)
Phosphate (mg/l)
Iron (mg/l)
Alkalinity(mg/l)
Specific Conductancefamhos/cm)
Total Dissolved Solids (mg/l)
Gross Alpha (pc/l)
Gross Beta (pc/l)
Cations (meq/1)
Anions (meq/1)
44
22
1.2
8.7
242
0
'11
11.7
15.9
0.9
<0.01
<0.02
198
402
222
<2
<10
4.4
4.6
WelH
4-Nov-92
51
26.2
<1
14.7
271
0
24
9.5
10.2
1.5
0.02
0.02
222
504
268
<2
<10
5.4
5.3
Well 2
4-Nov-92
51.5
26.3
<1
15.1
272
0
25
9.1
10.2
1.47
0.02
<0.02
223
506
268
<2
<10
5.4
5.3
24
-------
Table 9. Analytical Parameters and Laboratory Methods
Analyte
Laboratory Method
Calcium
Magnesium
Potassium
Sodium
Bicarbonate
Carbonate
Chloride
Silica
Sulfate
Nitrate+Nitrite
Phosphate
Iron
Alkalinity
Specific Conductance
Total Dissolved Solids
Gross Alpha
Gross Beta
EPA 200.7
EPA 200.7
EPA 200.7
EPA 200.7
Calculated
Calculated
Standard Method 407A
EPA 370.1
EPA 375.2
EPA 353.2
EPA 365.1
EPA 200.7
EPA 310.1
EPA 120.1
EPA 160.1
EPA 900.0
EPA 900.0
tion of the sample obtained in December 1990 also appears to be in error based on the lack of correlation with
other parameters. Ion balances for samples from West Hallings Spring and the monitoring wells have balance
errors between 2% and 5% (Table 8).
Major ion concentrations from these analyses were plotted on a trilinear diagram as developed by Piper
(1944) (Figure 12). Based on the hydrochemical facies of Back (1966), the trilinear diagram indicates a calcium
magnesium bicarbonate type ground water. Bicarbonate is the dominant anion in all of the analyses. The
predominance of bicarbonate and calcium is indicative of the carbonate aquifer through which the water flows
(Hem, 1985). No significant temporal variations in water type were noted. The small shifts in relative ion
concentrations observed in Figure 12 are due to the anomalous values which are believed to be errors. Analyses
of water from the two monitoring wells correlate closely with samples from Olsens Spring. Several of the
constituents in the sample from West Hallings Spring are present at relatively lower concentrations than in
samples from Olsens Spring or the monitoring wells. However, the samples were not obtained synchronously
and temporal variations in constituent concentrations may occur.
Saturation Indices and Calcium/Magnesium Molar Ratios
Analytical results of discharge from Olsens Spring were used by L. Spangler (U.S. Geological Survey,
Water Resources Division, Salt Lake City, Utah) to determine the relative states of saturation with respect to
calcite and dolomite and calcium/magnesium (Ca/Mg) molar ratios. Saturation indices for calcite and dolomite
(Table 10) were calculated using the U.S. Geological Survey WATEQ4F model (Ball and Nordstrom, 1991). As
analytical results for calcium (June 1991), sulfate (September 1991), and total dissolved solids (December
1990) are believed to be errors, alternate values were used for these calculations. Average concentrations for
calcium, sulfate, and total dissolved solids in the remaining three samples were used to calculate mineral satura-
25
-------
80 A A 80
60
60
S04+C1
Ca+Mg
\ A A A A A A A 7
^ 7\ A. A/ \ A A A A/
80 60 40 20
Ca
20 40 60 80
Cl
Figure 12. Trilinear diagram depicting geochemical analyses of water samples from OlsensSpring (+),
monitoring wells (O), and West Mailings Spring (0).
tion indices. Although there is uncertainty in using alternate values, the alternate values are believed to be closer
than the reported laboratory values to the actual values and resulted in better ion balances. It is noted that no
errors are suspected for the samples collected on March 14. The results of calculations using these data are
assumed to be reliable. The results from this sample indicate calculations using the alternate values are reason-
able and may be representative of site conditions. However, results of calculations using these data should be
considered preliminary. Additional data would be required to test significant conclusions.
Saturation indices for calcite ranged from -0.008 to +0.272, indicating that most samples were saturated or
near equilibrium with respect to calcite (Table 10). This implies a relatively long residence time indicative of
flow through a system of fractures that may have undergone relatively minimal amounts of solutional enlarge-
26
-------
Table 10. Calcite and Dolomite Saturation Indices and Ca/Mg Molar Ratios Calculated for Water Samples
from Olsens Spring
Date
17-Dec-90
14-Mar-91
13-Jun-91
18-Sep-91
Calcite
0.272
0.133
-0.008
0.156
Dolomite
-0.242
-0.515
-0.801
-0.461
Ca/Mg
Molar Ratio
1.19
1.19
1.21
1.17
ment. Ground water in systems primarily characterized by conduit flow (solutionally enlarged fractures) gener-
ally has a relatively short residence time and mineral saturation indices indicating undersaturation (more nega-
tive values) (White, 1988). Dolomite saturation indices ranged from -0.242 to -0.801, indicating undersaturation.
However, hydrologic (e.g., residence times) and geochemical factors inhibit dolomite from reaching saturation,
even in stratigraphic sections composed largely of dolomite (White, 1988).
The calcite saturation index (Figure 13) shows a trend suggestive of seasonal variations in saturation that
may be related to discharge. As discharge increases in the spring due to increased recharge, saturation of the
ground water with respect to all minerals would be expected to decrease as the result of decreased residence
time. In support of this interpretation, it is noted that discharge from seven springs in Mantua Valley (Figure 10)
peaked during June 1991. However, longer term monitoring of water chemistry and accurate measurements of
discharge from only Olsens Spring are required to verify this trend.
Calcium/magnesium molar ratios can be used to help understand flow paths and source areas for hydro-
logic systems in carbonate settings (Langmuir, 1971). Ground water that has moved through or been in contact
with formations composed predominantly of dolomite has a Ca/Mg molar ratio of approximately 1.00. Water
with a Ca/Mg molar ratio exceeding 3.00 has been in contact with, or flowed through rocks composed mostly of
limestone. Ratios between 1.00 and 3.00 indicate grouridwater flow through interbedded dolomite and lime-
stone, calcareous dolomite, or dolomitic limestone. Ratios of water samples obtained from Olsens Spring
(Table 10) ranged from 1.17 to 1.21, indicating that the ground water flowed through formations composed
largely of dolomite. Calcium/magnesium ratios calculated for West Hallings Spring discharge and the monitor-
ing wells (Table 11) are in the same range as values for water from Olsens Spring, also implying a ground-water
flow system that has been in contact with dolomite or dolomitic lithologies (L. Spangler, U.S. Geological Sur-
vey, Water Resources Division, written communication to M. Jensen, 1993). Differences in flow paths to the
springs could not be inferred from these data.
Two water samples were collected from Olsens Spring for tritium analysis, one each in June and October
1991. The tritium activities in these samples were 16.8 (±2.8) TU and 22.4 (±2.8) TU. These values indicate
that the average age of ground water in these samples is less than approximately 40 years and, possibly, less than
20 years (Fetter, 1988; Hendry, 1988).
Geochemical characterization of water from Olsens Spring, West Hallings Spring, and the monitoring
wells supports several hypotheses and conclusions concerning ground-water residence time and aquifer lithol-
ogy:
27
-------
0.3
0.2
CO
.•i
;o 0.1
CO
O
-0.1
Dec
Mar Jun
1991
Sep
Figure 13. Seasonal variation ofcalcite saturation index in water samples from Olsens Spring for the
sampling period of December 1990 to September 1991.
1. Lack of consistent seasonal variation in physical/chemical parameters, consistently low turbidity, and
saturation with respect to calcite indicate that the ground water discharging at Olsens Spring may be
migrating predominantly through fractures that have not undergone significant solutional enlargement.
However, these observations may also be explained by other conditions such as a lack of recharge from
only a few discrete points or by existence of substantial storage in a system with some component of
conduit flow (Quinlan and others, 1991). Additional studies, including more frequent monitoring of
parameters such as specific conductance, discharge, and turbidity, would be required to better define
potential conduit-flow contributions.
2. Potential seasonal fluctuation in the mineral saturation indices may indicate seasonal decreases in resi-
dence time related to increased recharge. This is also supported by the increased hydraulic gradient
estimated from ground-water elevation data obtained during May 1993.
3. The formations through which ground water migrates are composed largely of dolomite.
4. The average ground-water age is relatively young, less than 40 years. Additional studies (e.g., Busenberg
and Plummer, 1992; Solomon and Sudicky, 1991) may provide more information regarding the average
time period since recharge.
Table 11. Calcium/Magnesium Molar Ratios in Water Samples from West Mailings Spring and Monitoring
Wells Near Olsens Spring
Site
Ca/Mg Ratio
West Mailings Spring
Welll
Well2
1.21
1.19
1.18
28
-------
Catchment Area Calculations
A value of the average annual precipitation of 79 cm (30.9 in) was used (Table 4) for purposes of catch-
ment area estimation, assuming that the climatic setting is similar to the location at Pineview Dam. However, if
is noted that these recharge area calculations are sensitive to the recharge estimate, which is not well defined.
Thus, the results are subject to significant uncertainty. Data from Bjorklund and McGreevy (1974) indicate that
evapotranspiration on rangeland and mountainous areas at the elevation of Olsens Spring is about 35 cm (13.6 in)
per year. If it is assumed that the difference between precipitation and evapotranspiration is recharge to ground
water within the catchment area, then recharge would be about 44 cm (17.3 in) per year. This represents about
56% of total precipitation and is considered to be a significant overestimate of recharge for the assumed precipi-
tation. Insufficient data were available to conduct a detailed water balance. In order to bracket the catchment
area calculations, a recharge value of 10% of total precipitation was used as a lower bound. Based on the
fractured nature of the site, it is believed that 10% of available precipitation may represent the minimum re-
charge at this site. Use of the minimum recharge estimate will result in conservative calculation of maximum
catchment areas. The discharge of Olsens Spring is about 1700 1/min (1.0 ftVs). Using the method of Todd
(1980), a plot of catchment area as a function of estimated recharge rate and discharge for Olsens Spring indi-
cates that a catchment area ranging from about 2 km2 (0.8 mi2) to 10 km2 (4 mi2) would be necessary to supply
enough water for the spring to yield 1,700 1/min (450 gal/min) (Figure 14). West Hallings Spring, which is
located 270 m (900 ft) southwest of Olsens Spring, discharges from a limestone unit that is stratigraphically
below the discharge point of Olsens Spring. However, this spring is within the same surface drainage basin.
Total discharge from both springs is about 81001/min (4.8 ftVs). A catchment area ranging from about 10 km2
(4 mi2) to 54 km2 (21 mi2) would be required to provide enough water for the total discharge of these two springs.
This indicates that the zone of contribution for the two springs may be significantly larger than the topographic
basin surrounding the springs, which has a surface area of about 17 km2 (6.5 mi2). Mud Spring also may receive
at least a part of its recharge from the same surface drainage basin, but there are no data available to determine
to what degree Mud Spring affects Olsens Spring. If discharge from Mud Spring was included in the catchment
area calculations, the catchment area would be even larger than that calculated for Olsens and West Hallings
Springs. In addition, if the actual precipitation and, thus, recharge rates were less than the assumed values, the
required catchment area would increase.
Tracer Study
A tracer study was conducted to estimate ground-water velocity in the vicinity of the springs. The springs
were monitored for background fluorescence for ten consecutive days before injecting 500 ml of rhodamine WT
dye into Well 1, which is located 114 m (373 ft) from the Olsens Spring collection area (Figure 7). Approxi-
mately 1900 liters (500 gal) of water were injected to force the dye out of the well and into the aquifer.
Dye detection was by adsorption on activated charcoal contained in nylon screen packets that were placed
in Olsens Spring, the combined flow of Olsens Spring and West Hallings Spring, and in Mud Spring, which is
northeast of Olsens Spring (Figure 3). Due to the design of the water system, waters from Olsens Spring and
West Hallings Spring mix before water from West Hallings Spring can be sampled under normal conditions.
The charcoal detectors were initially changed as frequently as every 4 days, and were left in place for periods as
long as 40 days as the test continued. After removal from the springs, the dye detectors were eluted using a
mixture of ammonium hydroxide, 1-propanol, and distilled water and the solution was analyzed on a Turner
29
-------
10000
£2
M 1000
100
O
10
100
CO
&
cu
•3
(U
//
0.1 1 10
Liters/second
I
I
100
Spring Discharge
1 10
Meters^/second
Figure 14. Estimation of catchment areas for Olsens Spring and combined Olsens Spring and West Mailings
Spring. (Modified from Groundwater Hydrology, O.K. Todd, copyright © 1980, John Wiley & Sons,
Inc., reprinted by permission of John Wiley & Sons, Inc.)
filter fluorometer for the presence of dye (L. Spangler, U.S. Geological Survey, Water Resources Division,
written communication to M. Jensen, 1992). Analysis of dye recovery hydrographs (Figure 15) indicates that
dye was present on charcoal samples from the combined flow of Olsens Spring and West Hallings Spring. Dye
was not detected in discharge from only Olsens Spring or from Mud Spring during the test. This indicates that
dye only appeared in the discharge from West Hallings Spring. First arrival of the dye in the combined flow of
the two springs appeared to be between 12 days and 18 days after dye injection. Maximum response in the
combined flow of the two springs was approximately two months after dye injection. The quantity of dye in the
charcoal detectors in the combined discharge appears to have gradually decreased since that time.
Fluctuations in instrumental intensity producing secondary peaks during the test are related to length of
time that detectors were in the springs and are not the results of dye pulses. Peaks are also due, in part, to
adsorption of organic material. More consistent and accurate results would be obtained using an automatic
water sampler and fluorometer.
West Hallings Spring is approximately 350 m (1150 ft) southwest of Well 1, where the dye was injected.
Based on first inferred arrival of the dye and the straight-line distance from the spring to the well, ground-water
velocity potentially ranged from approximately 29 m/d to 19 m/d (96 ft/d to 64 ft/d) (i.e., dye arrival between
12 days and 18 days). Assuming the first arrival of the dye was in 15 days, the velocity would be 23 m/d
(77 ft/d). This calculated velocity range is subject to uncertainly resulting from assumptions involved in the
calculations (e.g., actual path length and arrival time). In addition, this velocity represents conditions at only
30
-------
2.5 •—
CO
UJ
LLJ
OLSENS/WEST MAILINGS SPRINGS
OLSENS SPRING
MUD SPRING
Q.
0)
CO
t>
o
o
o
o>
Q
CO
fc.
<
Figure 15. Relative instrumental intensity of eluted samples from Olsens Spring, Mud Spring, and the
combined discharge of Otsens Spring and West Hallings Spring.
one point in time. Seasonal fluctuations in hydraulic gradients may result in different ground-water velocities.
The azimuth from Well 1 to West Hallings Spring is approximately 190° and may indicate preferential flow
along north to northeast striking joints. This general flow direction appears to be consistent with the direction of
maximum gradient indicated from water levels in the monitoring wells and spring collection area (Figure 7)
within the limits of the uncertainty involved in each of these determinations. Results of this test indicate that the
ground-water velocity is relatively high in the vicinity of the springs and that Well 1 appears to be within the
zone of contribution to West Hallings Spring.
Hydrogeologic Conceptual Model
The conceptual model for site hydrology is that of a fracture-type spring (Figure 1) dominated by flow
through a fracture network and receiving recharge from a relatively large area. The potential recharge area may
be larger than the surface drainage area in which the spring is situated, indicating that these springs may be more
regional discharge points. This implies that the zone of contribution may extend beyond the boundaries of the
topographic basin. Additional studies would be required to confirm these hypotheses.
The evidence for this model includes highly fractured bedrock units, stable field parameters, relatively
stable water chemistry, very low turbidity, estimated size of required catchment basin, discharge that does not
appear to fluctuate daily or after storm events, and lack of observed surface karst features. In addition, the
carbonate units at this site are primarily dolomitic and diffuse-flow conditions (i.e., flow through closely-spaced
fractures interconnected, potentially, on the scale of centimeters) are common in dolomitic bedrock (Fetter,
1988).
31
-------
The only surface water body in the vicinity of the springs is Mantua Reservoir, which is located approxi-
mately 610 m (2000 ft) southeast of Olsens Spring and is outside the surface drainage basin. Surface water
elevations in Mantua Reservoir were several feet lower than ground-water elevations in the springs and monitor-
ing wells based on records maintained by Brigham City Corporation (Alan Wright, Brigham City Corporation,
personal communication to M. Jensen, 1997). The reservoir is not considered to have significant influence on
the zone of contribution to the springs in this study.
Based on the relatively high velocity, ground-water movement through the aquifer appears to be predomi-
nantly through fractures. The system appears to be anisotropic due to the existence of preferred joint orienta-
tions. More extensive aquifer testing may aid in determining hydraulic conductivity tensors. Rock units throughout
the surface drainage area may be sufficiently fractured to potentially allow the aquifer to behave as a porous
medium at the scale of the zone of contribution. However, this hypothesis is difficult to assess. A study
conducted by the Wisconsin Geological and Natural History Survey (1991) describes subjective criteria for
determining whether fractured-rock aquifers behave as porous media. These criteria are based on fracture
density, aquifer testing, spatial variations in water quality, and potentiometric information. Sufficient data from
subsurface investigations are not available, however, to make this determination at this site.
Aquifer Tests
A short-term, constant-rate pumping test was performed in each monitoring well at Olsens Spring. A
step-drawdown test was performed to determine sustainable discharge rates prior to the tests. Each well was
pumped for two hours and drawdown was monitored in the pumping well and in the observation well. Maxi-
mum drawdown in the pumping wells was approximately 0.3 m (1 ft) during each test. Significant drawdowns
were not noted in the observation wells. This probably indicates that the aquifer was not stressed sufficiently to
expand the cone of depression to the monitoring well during the short test. However, the tests also may have
been complicated by a lack of hydraulic communication between wells due to a lack of interconnecting frac-
tures.
Complications with pumping and metering equipment were encountered during the pumping portion of
the tests. Although the discharge rate appeared relatively stable by visual observation, the flowmeter used to
monitor discharge rates fluctuated significantly during the first 10 minutes of the test conducted in Well 2. The
discharge rate indicated by the meter during the remainder of the test varied approximately 8% with an average
rate of 132 1/min (35 gal/min). The average discharge rate for the test performed in Well 1 was 285 1/min
(70 gal/min) with a variation of approximately 11%. In addition, it appeared that Well 1 may have been poorly
developed prior to the test as well efficiency appeared to. increase during the test. Based on these complications,
only recovery data from the pumping wells (Figure 16) and time-averaged pumping rates were analyzed to
obtain a preliminary estimate of transmissivity. As it did not appear that actual discharge fluctuated signifi-
cantly during these tests, methods for analyzing recovery from stepped discharge tests (e.g., Kawecki, 1993;
Kruseman and de Ridder, 1990) were not applied.
The recovery data were interpreted using several solutions applicable to porous media. It is recognized
that the aquifer may not behave as a porous medium at the scale of these tests. Therefore, the results of these
tests are viewed as preliminary due to uncertainties associated with well losses, equipment complications during
the tests, subsurface anisotropy and heterogeneity, and application of porous-media assumptions. Extensive
32
-------
0.40'
^0.30
|
1 0.20 •
Q
15
lo.io-
0
a:
0.00'
Well!
1 1 llll
1.0
""I ' ' r^rr
10.0 100.0 1000.0
Dimensionless Time Ratio (t/f)
0.20
0.00
Well 2
1 ' I ' ' ' ' "" I ^ r
1.0 10.0 100.0
Dimensionless Time Ratio (t/t1)
1000.0
Figure 16. Residual drawdown plotted as function of the ratio of the elapsed time since pumping began (t) to
the elapsed time since pumping ended (f).
aquifer testing beyond the scope of this investigation would be required to test the porous-media assumptions
and better define hydraulic parameters. The interpretation methods used and transmissivity values calculated
from these tests are presented in Table 12.
Calculated transmissivity ranged from 270 m2/d (2900 fF/d) to 550 m2/d (5900 ft2/d). Based on these
results, the aquifer in the area of these wells appears to be relatively transmissive, in support of qualitative
hypotheses derived from hydrogeologic mapping. Calculation of hydraulic conductivity from these data in-
volves additional assumptions, including aquifer thickness influenced during pumping. In a porous medium
with fully penetrating wells, this thickness would generally be the saturated thickness within the well screen.
However, this thickness is unknown in a fractured setting with no definable confining units and may be consider-
ably greater than the screened interval of the well.
33
-------
Table 12. Interpretation Methods and Estimated Transmissivity from Aquifer Tests in Monitoring Wells at
Olsens Spring
Solution
Welll
Well 2
Theis (1935)
Theis Recovery Method
(Theis, 1935)
Neuman (1975)
Jacob (1950)
550 rtf/d (5900 fp/d)
510 rrf/d (5500 ff/d)
440 rrf/d (4800 ff/d)
540 rrf/d (5800 ff/d)
270 rrf/d (2900 ff/d)
440 rrf/d (4700 tf/d)
Not Applied
330 rrf/d (3600 tf/d)
Zone of Contribution and Delineation of Spring Protection Zones
General site hydrogeology and certain geologic and hydrologic controls on ground-water flow to the
springs were qualitatively defined using the mapping techniques applied in this study. However, definitive
ground-water flow boundaries were not delineated. In addition, the zones of contribution to Olsens Spring,
West Hallings Spring, and Mud Spring could not be differentiated using these techniques. At a minimum, it
appears that the zone of contribution to these springs may encompass the entire surface drainage basin (Fig-
ure 17), an area of approximately 17 km2 (6.5 mi2), and that ground-water velocities may be relatively high near
the springs. The results also indicate that the zone of contribution may be significantly larger than this area.
Extrapolating the velocity calculated from the tracer study, the time-of-travel from the northern extent of the
topographic basin to the area of the springs would be less than 200 days. Similar results are obtained using the
limited information available from the pumping tests and estimates of hydraulic gradient in the area of the
springs. A high degree of uncertainty exists in applying these velocity estimates over this distance as the
average velocity is probably significantly lower. However, this analysis indicates that the topographic basin
may not be sufficient for use as the 15-year time-of-travel boundary specified in the regulations promulgated by
the State of Utah. Additional studies, potentially including installation of a piezometric network, would be
required to better determine the zone of contribution and protection zones outside the surface drainage basin.
SHEEP SPRING
Location and Description
Sheep Spring is located in southwestern Utah, approximately 2.7 km (1.7 mi) northwest of St. George and
2.0 km (1.25 mi) east of Santa Clara in Washington County (Figure 18 and Table 13). This area of Utah is near
the western edge of the Colorado Plateau physiographic province (Stokes, 1986). The Santa Clara city water
system serves about 1,500 people, and uses Sheep Spring as a water source. The spring is at an elevation of
about 890 m (2920 ft). Sheep Spring discharges approximately 7.61/min (2 gal/min), making it a sixth magni-
tude spring (Table 2).
Sheep Spring was developed by digging a collection tunnel about 10.7 m (35 ft) eastward into a low
bench. The collection tunnel is roughly 1 m (3 ft) wide and 1.8 m to 2.1 m (6 ft to 7 ft) high. Ground water seeps
and flows from the north wall of the collection tunnel, ponds on the floor of the tunnel, and exits through a
discharge pipe to the water system.
34
-------
111°58.25'W
1
Estimated minimum
zone of contribution
Olsens Spring
West Railings
Spring
k-41°33'N
4000 feet
1000 meters
Contour Interval 200 feet
Datum is sea level
Mantua Reservoir
— 41°30'N
Figure 17. Surface drainage basin and estimated minimum zone of contribution to springs in this basin.
WASHINGTON
TUDYAREA
it. George
Miles
50
I
I
Kilometers
Figure 18. Location of Sheep Spring study area.
35
-------
Table 13. Location and Elevation of Sheep Spring and Associated Monitoring Wells Interpolated from
1:24,000-Scale Topographic Map
Sheep Spring
Monitoring Wells 1,2,3
Latitude
37°08'17"N
37°08'18"N
Longitude
113>37'06"W
113°37'06"W
Elevation
890 m (2920 ft)
893 m (2930 ft)
Climate
Average annual precipitation in the St. George area is about 21 cm (8.1 in), of which, 6.1 cm (2.4 in) is
snow. Average annual mean temperature is 16.8 °C (62.3 °F) and maximum annual evapotranspiration is ap-
proximately 160 cm (63 in) (Ashcroft and others, 1992). Vegetation in the Sheep Spring area is sparse desert
vegetation. The area is underlain by thin eolian and alluvial deposits, and exposed bedrock.
Previous Hydrogeologic Investigations
The first ground-water investigation that included the Sheep Spring study area was that of Cordova and
others (1972). This regional study presented information on the principal aquifers and on the recharge, move-
ment, discharge, chemical quality, storage, and utilization of ground water. A more recent study by Cordova
(1978) presents hydrologic data and a hydrologic evaluation of the Navajo Sandstone, a major aquifer in the St.
George area. A preliminary evaluation of the hydrogeology of Sheep Spring was completed by Jensen and
Lowe (1992). Clyde (1987) compiled a regional report which covers the St. George area, including the geology,
recharge, occurrence, movement, use, quality, and potential development of the ground-water resources. An
assessment of the low-temperature geothermal potential of the Santa Clara area, which also includes a geologic
map, was completed by Budding and Sommer (1986). Geologic maps that cover the Sheep Spring area include
a 1:125,000-scale map by Cook (1960) and a 1:31,250-scale engineering geology map by Christenson and Deen
(1983).
Geologic and Hydrologic Setting
Sheep Spring flows from the lower member of the Kayenta Formation, which is interbedded siltstone,
sandstone, and shale. The complete Kayenta Formation is approximately 655 m (2150 ft) thick in this area and
consists of three members. The Kayenta Formation is underlain by the Moenave Formation and overlain by the
Navajo Sandstone. All three formations are Jurassic in age (Hintze, 1988).
The Navajo Sandstone, overlying the Kayenta Formation (Figure 19), is a regional aquifer. The regional
ground-water flow direction in the Navajo Sandstone in this area is from north to south, from the Pine Valley
Mountains toward discharge at the Santa Clara River and Virgin River (Figure 20). The Pine Valley Mountains
are the regional recharge area for much of the St. George and Santa Clara area (Cordova, 1978).
Minor seeps are present along the topographic bench northwest of Sheep Spring, and springs are present
0.8 km (0.5 mi) and farther away, which discharge from a similar stratigraphic interval as Sheep Spring. As the
topographic basin in which Sheep Spring is located is large (130 km2, 50 mi2), and because other springs dis-
charge from the Kayenta Formation within the basin, Sheep Spring appears to represent only one of many local
discharge points for this aquifer system.
36
-------
Jms_
Scale 1:24,000
2000 feet
500 meters
37°7.5' N
A1
113°37.5'W
Contour Interval 200 feet
Datum is sea level
Qal
Qes
Qb
Qst
Jn
Jku
Jkm
Jkl
younger Holocene alluvium
eolian sand
Quaternary basalt
older Holocene spring tufa
Navajo Sandstone
Upper member of Kayenta
Formation
Middle member of Kayenta
Formation
Lower member of Kayenta
Formation
Explanation
Jms
X
\8
\87
Springdale Sandstone
Member of Moenave Formation
Spring
Strike and dip of bedding
Strike and dip of joints
Figure 19. Geologic map of the Sheep Spring study area.
37
-------
km
0
Contour Interval 200 feet
MOHAVE COUNTY ARIZONA
Figure 20. Potential zone of contribution to Sheep Spring and other springs in this area. (Base map from
U.S. Geological Survey 1:250,000 -scale topographic map.)
38
-------
Surface Mapping of Hydrogeologic Features
Prior to conducting geologic mapping, topographic maps of the area (1:24,000 scale) were studied to
identify the surface drainage divide upgradient of Sheep Spring. The divide, which appears to be the upgradient
extent of the potential zone of contribution to Sheep Spring, is approximately 24 km (15 mi) northeast of the
spring. Due to the significant distance to this divide, an area of approximately 4.5 km2 (1.75 mi2) surrounding
the spring was initially designated for detailed field study. Aerial photographs (1:20,000 scale) were also stud-
ied to identify fracture traces and lineaments in this area (Figure 21).The photographs displayed a pervasive
north-trending grain resulting from the dominant joint orientation (Figure 22). Only one individual feature,
approximately 1100 m (3600 ft) in length, was identified within the study area. This feature was located 1160 m
(3850 ft) from Sheep Spring.
The Kayenta Formation in this area consists of three members (Figures 19 and 23). Sheep Spring dis-
charges near the top of the lower, slope-forming member. This lower member consists of reddish-brown,
interbedded siltstone, sandstone, and shale, and is 90 m to 120 m (300 ft to 400 ft) thick. The middle member is
interbedded, moderate-reddish-orange and moderate-reddish-brown, fine- to medium-grained sandstone which
forms blocky ledges and slopes. The upper member consists of sandstone and crops out as massive, rounded
ledges. The Kayenta Formation dips approximately 8° to 12° east in this area (Figure 19).
The lower member of the Kayenta Formation forms a northwest-trending bench approximately 6 m (20 ft)
high. The spring collection tunnel is excavated into this bench. Several minor seeps are scattered along this
small bench northwest of Sheep Spring (Figure 19). These small seeps and accompanying vegetation appear to
be localized near outcrops of a light-colored sandstone layer. This layer is present in the top of the spring
collection tunnel, where it is 0.23 m (0.75 ft) thick, and was also encountered in a detailed measured section
approximately 30 m (100 ft) northwest of the collection tunnel, where it is 0.5 m (1.5 ft) thick. In outcrop, this
layer consists of moderately well-sorted silty sandstone which is medium-gray with pale bluish-green streaks,
has calcareous cement, and appears to have high porosity. This unit, or a similar unit in approximately the same
stratigraphic position, can be traced to the northwest of Sheep Spring about 670 m (2200 ft) and southeast about
21 m (70 ft).
Calcareous tufa deposits are associated with outcrops of the sandstone layer along the bench. The tufa is
light gray and porous. It is thickest (approximately 1.2 m (4 ft)) in an exposure about 14 m (45 ft) southeast of
Sheep Spring and directly overlies the thin sandstone layer. In other outcrops, the tufa is located directly on or
up to several feet above the sandstone layer. These tufa deposits, based on porosity and appearance, were
apparently deposited at land surface by the evaporation of spring water, then buried by eolian and alluvial
sediments.
Older tufa deposits are exposed about 1.9 km (1.2 mi) northwest of Sheep Spring. These deposits are
light to very light gray, friable to poorly indurated with blocky weathering, and up to about 1.2 m (4 ft) thick.
The tufa forms tabular deposits which cap the tops of sloping hills. The tabular deposits strike N28°W and dip
14° west. The deposits have been eroded, and appear to be remnants of a tufa sheet deposited on the land surface
before the slope retreated to its present position. These two levels of tufa deposits indicate that ground water has
been discharging from this area for thousands of years.
39
-------
She«p
Spring ""*""
Strike and dip of dominant joint set
Strike and dip of secondary joint set
Strike and dip of bedding
Strike of vertical joint set
Lineament from air photo
Scale 1:24,000
0 2000 feet
t— P-«
0 500 meters
Contour Interval 200 feet
Datum is sea level.
t
N
- 37° 7.5'N
Figure 21. Joint orientations in the Sheep Spring study area.
Figure 22. Hose diagram depicting strike of joints in the Sheep Spring study area. Unshaded sectors
represent strike of bedding. Number of joint measurments = 28, number of bedding
measurements = 5.
40
-------
Navaho
Sandstone
yenta Fm.
(Upper Member)
yenta Fm.
(Middle Member]
Sheep
Spring
renta Fm.
(Lower Member)
Moenave Fm
o 4000-
15
uJ 3000H
2000
2000 feet
I
0 500 meters
scale
Figure 23. Generalized geologic cross section through Sheep Spring study area.
The Kayenta Formation and Navajo Sandstone are highly jointed at Sheep Spring and north of the spring
(Figure 21). The dominant joints strike approximately north (Figure 22), and are nearly vertical, dipping steeply
to the east and west. The joints are clearly evident in outcrops and in the spring collection tunnel. Spacing of the
joints in the collection runnel ranges from 0.1 m to about 0.9 m (0.3 ft to about 3 ft), with areas of more extensive
jointing where the spacing is generally 0.2 m to 0.3 m (0.7 ft to 1 ft). Ground water discharging into the spring
collection tunnel appears to flow both from the joints (secondary porosity) and from pore spaces (primary
porosity) in the rock matrix on the north wall of the collection tunnel. Apertures of joints in the tunnel are very
small, approximately 0.15 cm (0.06 in) and less.
Joint spacing in outcrops in the study area varies from about 0.1 m to about 9 m (several inches to about
30 ft). In outcrop, there are localized areas displaying abundant joints with only 0.1 m to 0.2 m (several inches)
of separation between the joints. Joint lengths mapped in outcrops ranged up to the size of the outcrop, approxi-
mately 60 m (200 ft). Joint apertures measured in outcrop ranged up to about 1 cm (0.4 in). Thin calcite
coatings (0.25 cm (0.1 in)) were observed in some exposed joints. Joints appear to be more abundant in the
ledge-forming units, which are usually better indurated sandstone. The jointing extends stratigraphically up-
ward, and is well developed in the overlying Navajo Sandstone.
Based on the geologic mapping, several hypotheses regarding the hydrogeologic system may be formu-
lated:
1. Fracturing in the exposed bedrock may indicate a relatively high degree of vulnerability of the ground
water to surface contamination sources.
2. The hydrologic system may be highly anisotropic due to the pervasive north-trending joints. However,
as at Olsens Spring, sufficient data are not available to test this hypothesis.
3. Hydraulic conductivity may be relatively low in this system.
4. The potential zone of contribution to springs in this area may be relatively large based on surface
drainage divides present in this area.
5. Ground water discharging at Sheep Spring appears to be transmitted mainly, though not entirely,
41
-------
through a relatively thin sandstone unit in the vicinity of the spring. Flow appears to be through pore
spaces and joints.
6. Ground-water flow boundaries could not be inferred from surface mapping of hydrogeologic features
within the mapped area.
Borings and Well Installations
Three monitoring wells were installed near Sheep Spring (Figure 24) to provide stratigraphic control and
allow performance of aquifer tests. The boreholes were drilled using air rotary technology in interbedded
siltstone, sandstone, and shale of the Kayenta Formation. Well 1 was cored from 4.3 m to 7.9 m (14 ft to 26 ft)
below land surface and Well 2 from 6.1 m to 7.6 m (20 ft to 25 ft) to provide stratigraphic control across the thin
sandstone unit (Figure 25). Wet core and cutting samples were noted in Well 2 between 6.4 m and 7.3 m (21 ft
and 24 ft). The material consists of very fine- to fine-grained sandstone and interbedded siltstone. This appears
to be the same light-colored layer present in the top of the collection tunnel. A fracture observed in a core
sample through the sandstone layer in Well 2 is approximately 0.5 mm (0.02 in) wide, but is mostly filled with
silty material. Geophysical logging to aid in determining lithology was not conducted as originally proposed.
Coring was chosen over use of borehole geophysics as a more appropriate means of obtaining tight stratigraphic
control and collecting samples for laboratory tests of physical properties.
Horizontal hydraulic conductivity was estimated to be 2 x 10"7 cm/s (6 x 10"4 ft/d) using U.S. EPA Method
9100 for laboratory permeameter analyses (U.S. EPA, 1986) by the falling head method. The estimated hydrau-
lic conductivity is quite low but does not account for ground-water movement through fractures. However, it
does indicate that significant heterogeneity may exist in the sandstone matrix based on observations of flow
through primary porosity in the spring collection tunnel. An effective porosity of 13% was also estimated for a
sample from the base of this sandstone unit using Boyle's Law (Brown, 1981).
The wells were constructed using 10.2 cm (4 in) diameter schedule 40 PVC casing and screened across the
water table. A graded filter pack was installed within the screened zone, a bentonite seal was placed above the
sand pack, and the annulus was grouted to the surface. Each well was developed by alternately pumping and
injecting water obtained from the spring. The well was then pumped to remove the injected water following
development. The top of each well casing was surveyed with respect to the top of the spring collection tunnel.
The monitoring wells were installed during the final phase of field studies to take advantage of informa-
tion acquired during previous phases. Thus, limited information on temporal fluctuations in ground-water
elevations was obtained. As at the Olsens Spring site, continuous ground-water elevation data may have been
useful in understanding site hydrology and should often be obtained, if feasible. However, concerns for the
security of monitoring equipment prevented acquisition of such data at this site. Ground-water elevations in
Well 2 and Well 3 rose approximately 0.13 m (0.4 ft) between May 1992 and May 1993. The ground-water
elevation in Well 1 rose 0.25 m (0.8 ft) during the same period (Figure 26).The similarity in water level fluctua-
tions between Wells 2 and 3 and the difference between these wells and Well 1 may indicate that Well 1 was
notat equilibrium at the time of the measurement in May 1992, immediately following well installation. How-
ever, it is possible that water levels in any of the wells may not have returned to static conditions at the time of
the May 1992 measurements. Therefore, these measurements are considered to be suspect. Reliable hydraulic
gradients could not be calculated from relative ground-water elevations due to the close well spacing and the
42
-------
3
o
WELLS
\Tunnel
29 feet
Sheep Spring
5 meters
Figure 24. Location of monitoring wells near Sheep Spring.
Depth (below
land surface)
Depth (below
land surface)
m
cement
grout
bentonite
silica sand
size 10/20
4-*»
ft
' °
m
eolian sand
siltstone, with
interbedded sandstone,
1"=L-10 dark reddish-brown
^"^ and mottled with
yellowish-gray
sandstone
siltstone
sandstone,
very fine-grained
fracture noted in core
30 siltstone, sandy
cement
grout
0 bentonite
silica sand
-15 size 10/20
25
WELL #2
WELL #3
WELL #1
Figure 25. Construction logs of mentoring wells near Sheep Spring and representative geologic log of Well 1.
43
-------
M JJ ASOND JF MAM
1992 I 1993
Figure 26. Depth to ground water below the top of casing in monitoring wells at Sheep Spring.
small differences in hydraulic head. The well spacing was chosen to allow performance of a multi-well pump-
ing test for evaluation of transmissivity and anisotropy. Greater well spacing would be required to estimate
hydraulic gradient in this area. Ground-water elevations in the spring collection tunnel could not be used in the
calculations due to the manner in which the spring was developed. It appeared that the roof of the collection
tunnel was below the water table.
Physical and Chemical Properties of Discharge
Field Parameters
Discharge rate, temperature, pH, specific conductance, and turbidity of the water from Sheep Spring were
monitored on a monthly basis from November 1990 through October 1991 (Table 14). The discharge rate was
relatively constant at approximately 7.61/min (2.0 gal/min) throughout the 12-month sampling period. The flow
rate was determined at a drain outside the collection tunnel using a 7.6 1 (2.0 gal) bucket and stopwatch. The
temperature of the water varied through the sample period from a low of 16.8 °C (62.2 °F) in April 1991 toahigh
of 21.3 °C (70.3 °F) in September and October 1992 (Figure 27). Based on regional studies and the climate in
this area of Utah, the relatively high ground-water temperature does not appear to be related to geothermal
activity. The observed fluctuation appeared to be related to the mean monthly air temperatures recorded at the
weather station in St. George and may indicate seasonal changes in near-surface ground temperatures. Specific
conductance fluctuated only about 10% and ranged from 698 (imhos/cm to 770 |imhos/cm (Figure 28). These
44
-------
Table 14. Field Parameters for Water Samples from Sheep Spring
Date Temperature
(°C)
28-Nov-90
20-Dec-90
17-Jan-91
U-Feb-91
19-Mar-91
16-Apr-91
16-May-91
12-Jun-91
9-Jul-91
21-Aug-91
17-Sep-91
29-Oct-91
19.7
19.5
17.9
17.2
16.9
16.8
17.3
18.0
19.1
20.7
21.3
21.3
Specific Total Dissolved
Conductance Solids
(nmhos/cm) (mg/1)
770
727
720
741
698
736
733
735
730
762
727
729
371
363
360
370
348
370
367
368
365
381
364
365
PH
8.12
8.10
8.02
7.85
7.83
7.92
7.81
7.65
7.76
7.74
7.85
8.08
Turbidity
(NTU)
0.11
0.13
0.05
0.07
0.04
0.05
0.07
0.04
0.08
0.16
0.07
Flow
(l/min) (gpm)
7.9
7.9
7.9
7.9
7.9
7.6
7.9
7.2
7.9
7.2
7.2
2.1
2.1
2.1
2.1 .
2.1
2.0
2.1
1.9
2.1
1.9
1.9
35
30
25
cT
O
"^20
1
8.15
E
\—
10
5
0
$._ .
'$•••••
;%
* /
» *
* *
* /
, V.
p-
.6 ts^®^
^ — ^ — $r?-<$r~J^
cf
t
-^Spring Temp. -e-Mean Air Temp.
-I — U 1 1 J l i i i
Nov Dec Jan Feb Mar Apr May Jun Jul Aug Sep Oct
1990 199-j
Figure 27. Temperature of water samples from Sheep Spring and mean air temperature at St. George.
45
-------
1000
Specific Conductance -A- pH
i i i i i i
CO
400
Nov Dec Jan Feb Mar Apr May Jun Jul Aug Sep Oct
1990 1991
Figure 28. Specific conductance and pH of water samples from Sheep Spring.
minor fluctuations do not appear to correlate with changes in temperature. The pH ranged from 7.65 to 8.12
with only minor seasonal trends potentially present. The potential trends were not considered to be significant.
Turbidity was low and ranged from 0.04 NTU to 0.16 NTU.
Geochemical Characterization
Water samples from Sheep Spring were collected quarterly for one year and analyzed by standard labora-
tory techniques (Table 9) for major ions, selected trace elements, radionuclides, and total dissolved solids
(Table 15). Analyses of ground-water samples from the wells were not performed at this site as at the Olsens
Spring study site. The decision not to obtain samples from these wells was based on a high degree of confidence
in discharge location (i.e., Sheep Spring) for ground water at the location of the wells. Water samples from
Sheep Spring provide results that are more representative of average conditions within the zone of contribution.
Major ion concentrations from these analyses were used to construct a trilinear diagram as developed by Piper
(1944) (Figure 29). Based on the hydrochemical facies of Back (1966), the trilinear diagram indicates calcium-
sulfate type ground water. The ion balances for the four analyses had 0% to 4% error. The cations plot in the
field between calcium and sodium water types. Sulfate and bicarbonate are the predominant anions in all of the
samples. The relatively high concentrations of dissolved sodium, calcium, and magnesium may be due to the
fine-grained sediments in the Kayenta Formation.
The high sulfate content may indicate the presence of gypsum in the Kayenta Formation. The nearest
lithologic unit that contains large amounts of gypsum is the Carmel Formation, which overlies the Navajo
Sandstone. However, the nearest outcrop is approximately 11 km (7 mi) to the north of Sheep Spring (Cordova
and others, 1972). High sulfate content may also be due to sulfate cement in the Kayenta Formation. Sulfate
cement has been encountered locally in the Navajo Sandstone in the St. George area (R. Rasely, U.S. Soil
46
-------
Table 15. Analytical Results of Water Samples from Sheep Spring
Sample Date
Calcium (mg/l)
Magnesium (mg/l)
Potassium (mg/l)
Sodium (mg/l)
Bicarbonate(mg/l)
Carbonate (mg/l)
Chloride (mg/l)
Silica (mg/l)
Sulfate (mg/l)
Nitrate+Nitrite(mg/l)
Phosphate (mg/l)
Iron (mg/l)
Alkalinity (mg/l)
Specific Conductancefamhos/cm)
Total Dissolved Solids (mg/l)
Gross Alpha (pc/l)
Gross Beta (pc/l)
Cations (meq/l)
Anions (meq/l)
20-Dec-90
58
24
6
63
174
0
27
18
200
0.72
<0.01
<0.02
143
727
520
3
12.2
7.8
7.8
19-Mar-91
59
24
6
64
177
0
28.5
18
190
0.08
<0.01
<0.02
145
737
498
<1
<5
7.9
7.7
12-Jun-91
56
23
5.8
62
166
0
28
16.8
190
0.77
<0.01
<0.02
136
715
480
3
11
7.5
7.5
17-Sep-91
54
23
6.4
62
172
0
27.5
19
172
0.79
<0.01
<0.02
141
731
464
<1
<5
7.5
7.2
Conservation Service, personal communication to M. Jensen, 1992). Significant seasonal variation in
hydrochemical facies was not observed. Thus, no information regarding potential seasonal fluctuation in ground-
water flow path can be inferred.
Two water samples from Sheep Spring were collected for tritium analyses, one each in June and October
1991. Tritium analyses yielded activities of 2.4 (±2.8) TU and 1.9 (±2.6) TU. These low activities indicate that
recharge to the aquifer probably took place before 1953 (Fetter, 1988). These results indicate that the average
residence time of water discharging at Sheep Spring may be relatively long. No additional information regard-
ing seasonal fluctuations in ground-water flow paths or residence time could be reliably inferred.
Catchment Area Calculations
Average annual precipitation at the St. George weather station is about 21 cm (8.1 in) (Ashcroft and
others, 1992). There are no data from the vicinity of Sheep Spring that may be used to reliably estimate
evapotranspiration or recharge. For the purposes of this investigation it is estimated that recharge is equal to
10% of average annual precipitation or about 2.0 cm (0.80 in). However, that value may be an overestimate of
recharge in this arid area. The discharge of Sheep Spring is about 7.6 1/min (2 gal/min). Using the method of
Todd (1980), a plot of catchment area as a function of estimated recharge and discharge rates for Sheep Spring
(Figure 30) indicates that a catchment area of about 0.2 km2 (0.08 mi2) would be necessary to provide enough
water for the spring to yield 7.6 1/min (2 gal/min). Use of a recharge estimate of 0.2 cm (0.008 in) results in a
calculated catchment area of approximately 2 km2 (0.8 mi2). Minor, scattered seeps are present along the topo-
graphic bench north of Sheep Spring. However,.no data are available to'estimate the discharge of these seeps. If
the seeps were included in the calculations, the catchment area would be larger. The topographic basin in which
47
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80
60
S04+C1
( V V V V ••• '•'
\AAAA A/
V •./ •./ •./ ••./ •./
80 60 40 20 20 40 60 80
Ca Cl
Figure 29. Trilinear diagram depicting geochemical analyses of water samples from Sheep Spring.
Sheep Spring is located is significantly larger (130 km2,50 mi2) (Figure 20) than the estimated catchment area.
Although the estimated catchment area required to support the discharge is relatively small, little quantitative
information concerning the zone of contribution and its location can be inferred from this calculation.
Hydrogeologic Conceptual Model
Sheep Spring appears to be a contact or seep-type spring (Figure 1). The conceptual model, developed
from hydrogeologic mapping, is that of a spring discharging from a fractured clastic-rock aquifer of relatively
low hydraulic conductivity with ground water moving through primary and secondary porosity. A layer of very
fine-grained sandstone within the lower member of the Kayenta Formation appears to be transmitting much of
the water in the vicinity of the spring. The presence of spring tufa deposits indicates that ground water has been
discharging from the area of Sheep Spring for thousands of years.
Regional ground-water flow is from north to south. However, local flow directions may be affected by the
location and discharge rates of other springs in this basin. The system may be anisotropic and may not behave
48
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10000
1 10
Liters/second
100
1 10
Meters^/second
Spring Discharge
Figure 30. Estimation of catchment area for Sheep Spring. (Modified from Groundwater Hydrology, D. K.
Todd, copyright© 1980, John Wiley & Sons, Inc., reprinted by permission of John Wiley & Sons,
Inc.)
as a porous medium on the scale of a potential protection zone because essentially all of the fractures strike in
the same direction (Figure 21). Sheep Spring appears to be a local discharge point with, potentially, a relatively
large zone of contribution based on the locations of topographic divides. Ground-water residence times in this
setting may be relatively long. Based on average ground-water age, significant recharge may not be occurring in
the vicinity of Sheep Spring. No ground-water flow boundaries near the spring were defined. There are no
perennial streams or other major hydrologic features in the potential zone of contribution that may act as hydro-
logic boundaries. ;
Aquifer Tests
A short-term constant rate pumping test was attempted in Well 2. However, the sustainable discharge rate
was very low, approximately 0.71/min (0.18 gal/min). No significant drawdown was detected in either observa-
tion well, located 3.0 m (10 ft) and 3.2 m (10.4 ft) from the pumping Well during the three-hour test. The data do
not appear to adequately fit potential solutions for homogeneous porous media. It appears from a qualitative
analysis that the test may have been complicated by subsurface heterogeneity, low hydraulic conductivity, and
well storage effects.
Rising-head slug tests were performed in each of the three monitoring wells at Sheep Spring to estimate
hydraulic conductivity for comparison with qualitative assessments based on hydrogeologic mapping. The tests
were conducted by inserting a submersible pump in the well and allowing the water level to return to its static
elevation. Water was then rapidly pumped to lower the water level in the well about 0.9 m to 1.2 m (3 ft to 4 ft).
49
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This took only 5 seconds to 10 seconds and is considered to be approximately instantaneous under the condi-
tions of this test. The rising water level was then monitored using an electric water level indicator until it was
within 95% of the original static level.
Data from these tests (Figure 31) were analyzed using the method of Bouwer and Rice (1976), as modified
by Bouwer (1989). Data obtained during the first 30 seconds of recovery indicate a much higher recovery rate
than subsequent data. This appears to be due to backflow of water from the discharge hose recharging the well
as a functioning check valve was not in place in the discharge line during the tests. Drainage from the sand pack,
which has a relatively high hydraulic conductivity in comparison to the geologic formation, may also have
contributed to this effect. However, the dominant cause appears to be drainage from the discharge line based on
comparison of the water volume in the line and the anomalous head rise. Therefore, data collected during the
first 30 seconds of each test were not used in the interpretations. An aquifer thickness of 3.1 m (10 ft), which is
approximately the saturated thickness of the screened interval, was assumed. However, results of analyses
assuming an aquifer thickness of 15.2 m (50 ft) and 30.5 m (100 ft) were not significantly different from those
using an assumed thickness of 3.1 m (10 ft). Estimated hydraulic conductivities from the tests (Table 16) ranged
from approximately 0.04 m/d to 0.2 m/d (0.1 ft/d to 0.6 ft/d). These results support the conceptual model of an
aquifer with a relatively low hydraulic conductivity.
A similar value of 0.3 m/d (1 ft/d) for hydraulic conductivity was estimated by Cordova and others (1972)
based on the specific capacity of a well located about 4 km (2.5 mi) from Sheep Spring and aquifer charac-
teristics of the Kayenta Formation. Studies have not been sufficient to adequately define subsurface heterogene-
ity and the potential range of hydraulic parameters. For example, wells used for hydraulic testing may not
intersect transmissive joints and results may not be representative of the most conductive areas or average
conditions at the site. More extensive studies would be required to better define hydraulic parameters within the
potential zone of contribution to Sheep Spring. However, it is questionable as to whether such studies would
provide significantly more reliable estimates in a cost-effective manner. A tracer test was not conducted at this
site as it was at Olsens Spring due to the increased confidence in the discharge point for ground water in the
vicinity of the wells and the difficulties involved in conducting representative tests in materials of low hydraulic
conductivity.
Zone of Contribution and Delineation of Spring Protection Zones
Hydrogeologic mapping indicated that the potential zone of contribution to springs in this area may ex-
tend as much as approximately 24 km (15 mi) north of Sheep Spring. This would be considered to be the
maximum plausible extent. The actual zone of contribution may be much smaller. Many factors, including
ground-water discharge from nearby springs, control local flow directions. Ground-water flow boundaries
closer to Sheep Spring could not be inferred from hydrogeologic mapping methods applied in this study. Con-
servative protection zones that cover the entire surface drainage area (i.e., more than 130 km2) may be consid-
ered too large to be effectively managed by the water supplier. However, a relatively low ground-water seepage
velocity may potentially limit the zone of contribution within the 15-year time period to a much smaller area.
Therefore, incorporation of additional delineation techniques (e.g., estimates of ground-water time-of-travel)
may be required to define more practicable protection zones.
50
-------
10.00
f" 1'°°"t
e 0.10- -
0
I 0.01--
0.00
10.00
i.1.00--
TB
m
o>
•E o.io-L
O)
I 0.01 J-
o.oo.
WelH
50 100 150 200
Time (minutes)
Well 2
10 20 30 40 50
Time (minutes)
Well 3
10.00
J1.00-
~ 0.10-
0 0.01-
0.00-
•
-^
H I I I I 1 1 1 I n 1 I I I I I linn i i i i i I i i i i i i i ulin i n i i
10 20 30 40
Time (minutes)
50
Figure 31. Recovery data from slug tests in wells near Sheep Spring.
Table 16. Hydraulic Conductivity Estimated from Slug Tests in Wells at Sheep Spring
Well
Hydraulic Conductivity
(m/d) (ft/d)
1
2
3
0.04
0.2
0.1
0.1
0.6
0.4
51
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The information regarding hydraulic parameters obtained during this study would generally not be suffi-
cient to define reliable protection zones based solely on time-of-travel criteria. However, ground-water seepage
velocity was estimated from these data to illustrate the potential use of such criteria in refining the protection
zone delineation. For purposes of these preliminary calculations, a hydraulic gradient of 0.05 was assumed
based on land surface topography. The implicit assumptions are that the aquifer is unconfined and the water
table mimics topography. This is the approximate slope of the relatively flat area directly north of the spring,
and also the slope from a surface drainage north of the spring to the spring collection area. This gradient appears
to be relatively high and significant uncertainty exists in applying this estimate throughout the potential protec-
tion zone over the required time periods.
Minor seeps and discharge into the collection tunnel indicate that much of the ground water at the dis-
charge point is moving through a sandstone layer. Effective porosity estimated from a core sample of this unit
was 13%. Using these parameters and a hydraulic conductivity of 0.3 m/d (1 ft/d), the estimated average
interstitial velocity is 0.12 m/d (0.4 ft/d). This velocity yields a 250-day travel distance of 30 m (100 ft) and a
15-year travel distance of 660 m (2200 ft) topographically upgradient of the spring (Figure 32). These potential
protection zones cover only a part of the potential zone of contribution and would be a more manageable and,
perhaps, more reasonable protection zone for the water supplier. As previously noted, significantly more
hydrogeologic data from subsurface investigations would be required to estimate more reliable protection zones
using such criteria.
113°37.5' W
37°7.5' N
0 2000 feet
B . ' Contour Interval = 200 feet
500 meters
Datum is sea level
Figure 32. Example of potential protection zones at Sheep Spring based on preliminary time-of-travel
calculations. Small zone around spring is 30 m (100 ft) travel distance and large area is 660 m
(2200 ft) travel distance.
52
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Chapter 4
CONCLUSIONS AND RECOMMENDATIONS
EVALUATION OF HYDROGEOLOGIC MAPPING METHODS
Results from these case studies support the conclusion that hydrogeologic mapping methods may be used
at many sites to provide basic site characterization information. Case studies performed in significantly differ-
ent settings would be required to evaluate the relative utility of different mapping methods in each setting. This
investigation took place in the arid Southwest where vegetation is sparse and population density is relatively
low. One of the more powerful hydrogeologic characterization techniques (potentiometric surface mapping)
could not be used in this study. This method of ground-water flow field analysis may be very effective in other
areas where population density is higher and more ground-water elevation data are available. Other methods
used in this study, such as detailed geologic mapping, may not be effective in many areas due to dense vegetative
cover.
Potential ground-water flow controls and boundaries may be located using these techniques. These con-
trols may be geologic (e.g., areal limits of aquifer formations) or hydrologic (e.g., ground-water divides). At
some sites, hydrogeologic mapping may result in delineation of ground-water flow boundaries suitable for use
in directly defining protection zones around springs. Although these methods will not be capable of fully
defining the zone of contribution to springs at all sites, results of such investigations will provide information
needed for more informed decisions regarding potential zones of contribution and appropriate protection zones
at most sites. As in this study, such mapping may also serve as a basis from which additional investigations to
better define the zone of contribution, hydraulic parameters, and subsurface heterogeneity may be assessed.
In these case studies, hydrogeologic mapping provided information on aquifer characteristics including
aquifer lithologies, relative ground-water residence times, qualitative evaluation of hydraulic parameters, poten-
tial recharge areas, and potential physical/hydrologic controls on ground-water flow. This information was
valuable in building the conceptual models for site hydrogeology, estimating potential zones of contribution,
and constraining potential protection zones around each spring. However, ground-water flow boundaries suit-
able for delineating practicable protection zones with a high degree of confidence could not be identified at
either site using the techniques applied in this study. Additional subsurface studies may be warranted at each
site to support delineation of more reliable and more practicable protection zones.
Application of these techniques in the carbonate aquifer (Olsens Spring study site) provided sufficient
information to indicate that the portion of the aquifer supplying water to Olsens Spring may behave predomi-
nantly as a diffuse-flow system as termed by Shuster and White (1971). In other words, flow may be predomi-
nantly laminar along joints, fractures, bedding planes, and other features interconnected at a scale much smaller
than the site scale and, potentially, on the scale of centimeters to meters. However, additional studies would be
required to better define potential conduit-flow components.
Geologic mapping indicated the area surrounding the spring may be relatively vulnerable to ground-water
contamination from surface sources. Ground-water velocity near the springs, estimated from a tracer study and
hydraulic parameters, was relatively high. Results also indicated that the zone of contribution to Olseris Spring
and West Hallings Spring may be larger than the topographic basin surrounding the springs. This implies that
53
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reliable protection zones may also extend beyond the basin. Available information indicated that potential
protection zones should, at a minimum, encompass the entire surface drainage basin.
Additional studies designed to obtain piezometric data within and surrounding the drainage basin may
have provided valuable information for evaluation of the shallow flow system, but were not feasible in this
investigation due to a lack of available wells. Evaluation of piezometric information may be more feasible in
other areas where private wells are installed into the aquifer under investigation. Information regarding the role
of faults in ground-water movement, potential presence of shallow ground-water divides, and improved fracture
characterization may be obtained from such studies. Additional tracer and more extensive aquifer tests could be
conducted within the potential zone of contribution to better estimate ground-water velocity and anisotropy.
Use of naturally occurring isotopes may also be valuable in tracing ground water from recharge to discharge
following installation of additional piezometers. Additional monitoring of mineral saturation indices and fre-
quent monitoring of discharge rates of individual springs could be instituted to confirm temporal trends ob-
served in this study. Other evaluation tools include analysis of chlorofiuorocarbon compounds and tritium/
helium-3 dating techniques to better estimate ground-water residence time. Detailed monitoring of spring parame-
ters such as discharge rate, water temperature, turbidity, or specific conductance before, during, and after re-
charge events may also provide a better indication of whether a significant component of conduit flow exists.
In the fractured clastic-rock aquifer (Sheep Spring study site), the potential zone of contribution to springs
in this area may be bounded by a ground-water divide located about 24 km (15 mi) north of Sheep Spring.
However, flow boundaries closer to the spring could not be inferred using the methods applied in this study. Due
to the potentially long ground-water residence time and the relatively large extent of the aquifer that was indi-
cated, protection zones delineated using other techniques, such as estimation of ground-water time-of-travel,
may be more practicable.
Additional studies designed to provide more detailed information regarding hydraulic parameters and
potential subsurface heterogeneity would be required to more reliably estimate ground-water time-of-travel to
Sheep Spring. Such studies could include aquifer testing to define anisotropy, potentiometric surface mapping,
and use of tools such as the borehole flowmeter to define the relative flow contribution from various lithologic
units and from fractures within those units. A more quantitative determination of ground-water residence time,
potentially comparing information from radiocarbon, chlorofiuorocarbon, and tritium/helium-3 analyses with
estimates derived from hydraulic information, may also be useful in evaluating potential protection zones.
GUIDANCE FOR APPLICATION OF TECHNIQUES
The hydrogeologic mapping methods applied at any spring site will depend on such factors as hydrogeologic
setting, availability of existing hydrogeologic information, and management objectives to be applied within the
delineated protection zones. Integration of multiple mapping techniques will be required to develop the
hydrogeologic conceptual model and define potential ground-water flow controls and boundaries to support the
delineation. The success of all of these methods depends greatly on the hydrogeologic complexity of the spring
site.
A general methodology for aquifer characterization using hydrogeologic mapping begins with a review of
site-specific hydrogeologic literature to develop an initial conceptual model for site hydrogeology. The concep-
54
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tual model is the'synthesis of available information designed to .convey the assumptions and hypotheses regard-
ing ground-water flow to the spring. Development of a conceptual model is an essential element in the delinea-
tion process. The model serves as the focus for characterization efforts.
Important assumptions and potential ground-water flow controls should be identified from this model.
Hydrogeologic mapping and subsurface investigations may then be designed and implemented to test these
assumptions and delineate specific physical and hydrologic controls. Results from these studies may be used to
refine the conceptual model, estimate the zone of contribution, and, potentially, delineate protection zones based
on controls. Important information gaps may be identified and additional investigations proposed to fill these
gaps.
Review of previous investigations, analyses of topographic maps, fracture trace analyses from aerial pho-
tographs, estimation of catchment area, evaluation of geochemical parameters, and construction of potentiomet-
ric maps should provide valuable information for model development at most sites. Detailed geologic mapping,
supported by surface and borehole geophysical surveys, may be useful at many sites in defining the boundaries
of formations acting as aquifers and aquitards and locating features that may serve as pathways or boundaries for
ground-water flow (e.g., fracture systems and faults). Isotope studies may be used to constrain ground-water
age, residence time, and location of recharge. Specialized tools, such as tracer studies, may provide direct
information regarding ground-water velocity and zone of contribution, particularly in karst settings.
Hydrogeologic mapping techniques are relatively low cost characterization tools. Many of these tech-
niques are also simple to apply. The applicability of particular tools at spring sites depends on site conditions
and the nature of potential ground-water flow controls. Many of these methods may be most applicable at sites
where little information concerning the hydrogeologic setting is available. It is recommended that the potential
utility of such techniques be evaluated early in all spring protection zone delineation projects.
55
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60 &U.S. GOVERNMENT PRINTING OFFICE:
1997- 549-001/«OISZ
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