Milli"

                                                                            

 

 

 

 

 

 

 

321 ‘2‘93\ 1‘0681 736

 
 
 

Llamm
fiiehigan State
University

 

This is to certify that the

thesis entitled

THE VARlATlON I‘N HYDRAULIC CONDUCTIVI‘TY
OF HETEROGENEOUS DRIFT DEPOSlTS

presented by

CHERYL A. KEHRES

has been accepted towards fulfillment
of the requirements for

fl' 5:. degree in%,

QM:%

Dr. Grahame J. Larson

 

professor

Date game a; may

0-7639 MS U is an Affirmative Action/Equal Opportunity Institution

 

 

 

MSU ‘ RETURNING MATERIALS:
Place in book drop to
LIBRARIES remove this checkout from
Ailing-III. your record. FINES will

be charged if book is
returned after the date
stamped below.

 

 

 

 

  

{Us 1 :45- 1.3.3

 

 

THE VARIATION IN HYDRAULIC CONDUCTIVITY
OP HETEROGENEOUS DRIFT DEPOSITS

by
Cheryl A. Kehres

A THESIS

Submitted to
Michigan State University
in partial fulfillment of the requirements
for the degree of

MASTER OF SCIENCE
Department of Geology
1984

{‘3'“

ABSTRACT

THE VARIATION IN HYDRAULIC CONDUCTIVITY
OF HETEROGENEOUS DRIFT DEPOSITS

BY
Cheryl A. Kehres

Aquifer test analysis and laboratory permeability studies
are the most common methods which have been used to
determine the hydraulic conductivities of a leaky artesian
system in a glaciated area. These techniques, however, can
not adequately determine the variability in hydraulic
conductivity of drift deposits that is expected over a broad
regional area. In this study a finite-element model is
applied to such an area in an attempt to estimate this
anticipated range in hydraulic conductivity.

The area chosen for this investigation lies in south-
central Michigan and encompasses 16 townships. It is
underlain by Pleistocene drift that rests directly on the
major bedrock aquifer for the region. A data base of
approximately 2000 well logs was used to define drift and
aquifer lithology, and areas of probable high recharge. The
hydraulic conductivity values of drift determined from this
research are 1.8*10'43ft/sec (S.5""10'7 cm/sec) for till
deposits and 1.8*10'7 ft/sec (E'>.5*10"6 cm/sec) for sand and
gravel deposits. It can be concluded that application of a
finite-element model can assess the variability in hydraulic
conductivity of drift deposits over a regional area and
should provide a more accurate description of the

hydrogeology of regions characterized by such deposits.

To Matt

ii

ACKNOWLEDGEMENTS

There have been many individuals who deserve
acknowledgement for their assistance during the course of
my research. I would like to extend my appreciation to my
thesis committee members, Dr. John Wilband and Dr. James
Trow for their review of this manuscript. I also wish to
extend my deepest appreciation to Dr. David Wiggert for his
patience and guidance in teaching me the application of
finite-element analysis. In addition, I wish to extend my
sincere appreciation to Dr. Grahame Larson, my committee
chairman, who not only taught me how to approach a research
problem and provided expertise in the field of
hydrogeology, but also gave support to me as a friend.

I would also like to thank Mr. Ron Hohenstein of the
Lansing Board of Water and Light, Mr. David Jansen of the
East Lansing - Meridian Township Water and Sewer Authority,
Mr. Bob Roller of the Tri-County Regional Planning
Commission, and individuals of Delta Township, Snell
Environmental Group, and Keck Consulting. These
individuals were most helpful in extending the technical
data which made my research possible.

Also, I wish to thank Dr. Genevieve Segol and Mr. Tim
Kelly who gave me the opportunity to gain professional
experience while I was working on my Master‘s degree, as
well as Dr. Satya Sehgal and Kim Kesler of GMC Associates
for giving me the flexibility to finish my thesis.
Finally, I wish to thank Bill Monaghan for all his help as

well as Sue, Sandy, Bruce, Steve, and my parents.

iii

TABLE OF CONTENTS
LIST OF FIGURES............................................v
LIST OF TABLES............................................Vi
INTRODUCTION...............................................1
STUDY AREA.................................................2
HYDRAULIC CONDITIONS.......................................8
FINITE-ELEMENT MODEL......................................15

Finite-Element GridOOOOOOOOO00....00.0.0000000000000016
AssumptionSOOOO0.0..0.0.0...I...00.0.000000000000000016

CALIBRATIONOOOOOO...O...OOOOOOOOOOOOOOO0.00.00.00.0000000021
DISCUSSIONOOOOOOOOOOOOOOOOOOOOOOO00......0.0.0.0000000000042

Electric-Analog MOdeIOOOOOOOOCOOIO0.0.00.00000000000042
DewateringOIOOOOCCOO0.0.0.0000...0.0.0.0000000000000043

CONCLUSIONSOOOOOOO0.0...OOOOOOOOOOOOOOOOOOOOOO0.000.000.0045
BIBLIOGRAPHYOOOOO0.00.00.00.00...O...00.00.000.0000000000047

APPENDIXOOOOOOOOOOCOOOOO0.0.0....O...OOOOOOOOOOOOOOOOOOOOO49

iv

Figure

Figure

Figure

Figure

Figure
Figure
Figure
Figure
Figure
Figure
Figure
Figure
Figure

Figure

LIST OF FIGURES
Study AreaOOOOOOOOOOOOOOOOOOOOOOOOOOOOOOOOO0.04

Schematic of Drift Overlying the Saginaw
FormationOOOO0.0.0..OOOOOOOOOOOOOOOOOOOOOOOO0.7

BedIOCkLith0109y MaPOOOO0.0.0.0000000000000010

Areas of Direct Hydraulic Connection through
Drift to BedIOCk....00...0.00.00.00.00000000013

Finite-Element Grid..........................18
Finite-Element Grid with Pumping Centers.....50
Simulated 1945 Potentiometric Surface........24
Actual 1945 Potentiometric Surface...........26
Simulated 1964 Potentiometric Surface........28
Actual 1964 Potentiometric Surface...........30
Simulated 1979 Potentiometric Surface........33
Actual 1979 Potentiometric Surface...........35
Final Simulated 1945 Potentiometric Surface..38

Final Simulated 1964 Potentiometric Surface..40

LIST OF TABLES
Table 1: Pumping Schedules..............................Sl

Table 2: Hydraulic Conductivity of Drift................46

vi

INTRODUCTION

Often in glaciated areas the hydrologic regime is that
of a leaky artesian system with drift acting as a semi-
pervious layer. This condition has prompted several
investigations in recent years that have focused primarily
on the hydraulic properties of drift. They have included
laboratory permeability studies (Desaulniers, et a1. 1981;
Norris, 1962), aquifer testing (Walton, 1962), and digital
modeling (Grisak and Cherry, 1975). Collectively, these
studies have shown that drift can have a range in hydraulic
conductivity on the order of 10'13 to 10-7 ft/sec (10'11 to
10'5 cm/sec).

The above investigations have been mainly site
specific rather than regional, and have generally treated
the drift as a homogeneous unit. However, in a broad
region characterized by heterogeneous drift deposits a
range in hydraulic conductivity values as opposed to a
single value can be expected. Therefore, in this
investigation an analysis utilizing a finite-element model
is applied to a broad region characterized by varied drift
deposits in an attempt to estimate this anticipated range

in hydraulic conductivity.

STUDY AREA

The area chosen for this investigation lies in south-
central Michigan (Figure 1) and encompasses 16 townships
that make up the Lansing - East Lansing metropolitan
region. It is underlain by Pleistocene drift that varies
from 50 to 150 feet (15 to 46 m) in thickness (Stuart,
1945; Vanlier, et al. 1973). This drift rests directly on
the Saginaw Formation which is the major bedrock aquifer
for the region. The Saginaw Formation is confined
underneath by the Bayport Limestone. The entire area
depends almost exclusively upon groundwater for its water
supply (39.6 mgd) and contains approximately 1800 private
wells and 167 municipal and industrial wells.

The Saginaw. Formation is composed chiefly of
interbedded sandstone and shale (Kelly, 1936; Vanlier, et
a1. 1973). It dips slightly towards the northeast and
ranges in thickness from approximately 350 feet (107 m) in
the southern part of the study area to 400 feet (122 m) in
the northern part. The transmissibility of the aquifer has
been determined from previous aquifer studies (Stuart,
1945; Firouzian, 1963) and ranges between 6.2“10-3 ft2/sec
(5.76 cm2/sec) and 0.12 ft2/sec (111.5 cmzlsec). Assuming
an average aquifer thickness of 375 ft (114 m) these values
result in a range of hydraulic conductivity from 1.2“?10'S
ft/sec (3.6*1o'4 cm/sec) to 3.3*1o'4 ft/sec (1.o*10"2

cm/sec).

'V
I I

m

'C

Figure 1. Study Area. Darkened circles indicate

municipal wells.

 

 

 

 

 

 

 

 

 

 

 

 

°F

1 #1
6d!“

 

 

 

5

The drift in the region consists mostly of clay till.
However, in some areas there are localized deposits of sand
and gravel associated with eskers and rivers. It has been
proposed that recharge through the drift to the underlying
aquifer occurs mainly through the permeable materials
associated with the esker and river deposits; whereas, less
recharge occurs through the till or where shales are in
direct contact with the drift (Stuart, 1945; Vanlier, et
al. 1973; Long, et al. 1981; Monaghan and Ritter, 1982),
(Figure 2). It is this geologic setting which creates a
leaky artesian groundwater system in the study area. A
value of 1.8*10'7 ft/sec (5.5""10-6 cm/sec) has been
suggested for the hydraulic conductivity of the sands and
gravel (Lovato and Larson, 1979). However, conductivity

values for the till have not been determined.

Figure 2. Schematic of Drift Overlying the Saginaw

Formation.

 

 

DRIFT

SAGINAW

 

Sandstone

 

 

HYDRAULIC CONDITIONS

A bedrock lithology map was prepared from a
computerized data base of 893 private and municipal wells
in the study area. To determine the lateral and vertical
change in lithology of the aquifer, the Saginaw Formation
was divided into three major lithologic units: sandstone,
shale, and sandstone and shale (i.e. sandy shale, shaley
sandstone, and thin beds of alternating shale and
sandstone). If a well log recorded greater than 33%
sandstone, the bedrock in the vicinity of the well was
defined as sandstone. Similarily, if the well log recorded
greater than 33% shale or greater than 33% shale and
sandstone, the bedrock in the vicinity of the well was
defined as shale or shale and sandstone, respectively.
This averaging technique was applied at depths of 10 feet,
100 feet, 150 feet, and greater than 300 feet into the
aquifer. It was found that the aquifer lithology was
fairly constant with depth and consisted of a higher
percentage of sandstone in the center of the study area
than to the east or west (Figure 3). In conjunction with a
review of results of numerous pump tests which have been
performed in the study area (Keck Consulting; Stuart, 1945;
East Lansing - Meridian Water and Sewer Authority),
defining the bedrock lithology enabled assignment of
hydraulic conductivity values to the aquifer. As a result,
the center of study area was initially assigned a

8

.‘fl

Figure 3. Bedrock Lithology Map. Shaded area
represents mostly sandstone; unshaded area represents
mostly shale. Darkened circles indicate municipal

wells; open circles indicate private wells.

‘II.

 

 

 

 

 

 

 

 

 

I‘ll

11
conductivity value of 7.9*10"5 ft/sec (2.4*10"3 cm/sec),
whereas, the area consisting of a higher percent of shale
was assigned a conductivity value of 2.3*"10"S ft/sec
(7.0“'10"'4 cm/sec).

Once the aquifer conductivity was defined, three
criteria were applied to define areas of direct hydraulic
connection through the drift to the bedrock. First, the
well file was searched for those wells which recorded
continuous units of permeable sand and gravel from the
ground surface to the bedrock aquifer. Next, it was
determined if those wells which recorded continuous sand
and gravel were in direct contact with sandstone. Finally,

it was determined if the areas defined by the first two

conditions were near any rivers or other recharge sources._

From this assessment, the area where the Red Cedar and
Grand Rivers converge was identified as a location where
substantial recharge to the aquifer seemed probable (Figure
4). Also identified were a few deposits of smaller areal
extent that were in direct contact with sandstone. With
these areas of suspected high recharge differentiated, a
conductivity value of 3.3“'10"9 ft/sec (10'7 cm/sec) was
assigned to the till; a value of 3.3*10-8 ft/sec (10"6
cm/sec) was assigned to the permeable sands and gravel of
the suspected high recharge areas.

In conjunction with the hydraulic values of the drift
and bedrock aquifer, pumping data were assembled for the

period 1945-79. Because of the configuration of the

-41

12

Figure 4. Areas of Direct Hydraulic Connection through
Drift to Bedrock. Circles represent wells that
recorded continuous sand and gravel to bedrock. Shaded

area represents identified high recharge area.

 

 

 

 

 

 

 

 

 

 

 

 

5 miles

14
municipal well field, the wells were then clustered into
pumping centers. With this information a finite element

grid was designed.

FINITE-ELEMENT MODEL

The finite-element model (GEOFLOW) used in this study
is a two dimensional computer program which utilizes the
Galerkin-based finite-element method to transform the
mathematical equations governing groundwater flow to a
system of first-order partial differential equations as a
means of evaluating the potentiometric surface over the
study area (Haji Djafari, 1982). Features of the
hydrodynamic model are:

1) Time-variable pumping rates can be modeled.

2) The system can be divided into regions, and
each region can have different hydraulic
conductivity parameters such as the
transmissibility, recharge, and storage
coefficients.

3) Initial hydraulic heads can be specified or
can be computed by the program.

4) Along any designated element, a line source
can be incorporated.

5) At any designated node, the value of the
hydraulic head can be specified (Dirichlet
boundary condition).

6) The thickness of aquifer can be updated if
required.

7) Flow vectors are computed at the center of
each element.

15

16

Finite-Element Grid

In the design of the finite-element grid (Figure 5)
three major factors were considered. First was the aquifer
hydraulic conductivity, second was the configuration of the
municipal well field, and third was the hydrologic
character of the drift. In regards to the aquifer
'hydraulic conductivity the elements of the grid were
designed to conform to the facies relationship between the
two conductivity zones as previously defined. To establish
more detail near the municipal well field, the elements
were also made smaller nearer the well field and generally
larger farther away. The finite-element grid delineating
pumping centers is included in the Appendix (Figure 6).
The heterogeneous nature of the drift was also delineated
by assigning a higher vertical conductivity value over the
elements representing the suspected high recharge area
(Figure 5). In total, the finite-element grid consisted of

209 elements and 118 nodes.

Assumptions

 

The main assumptions that were used to implement the
program were as follows:

1) The aquifer was treated as being under semi-

confined conditions. Although the drift is

acting as the overlying confining bed, there

is significant amount of leakage through the

17

Figure 5. Finite-Element Grid. Shaded area represents

identified high recharge area.

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

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19
heterogeneous deposits characterizing the
drift to the Saginaw aquifer.
2) Each node along the boundary of the study
area was assumed to have a constant hydraulic
head as prescribed by the Dirichlet condition
(Prescribed Potential Boundary Condition)
since historical data has shown that the
potentiometric surface around the perimeter
of the study area has not changed appreciably
since 1900.
The thickness of the aquifer was assumed
constant at 375 feet (114 m). An isopach map
of the thickness of the Saginaw Formation was
constructed from oil well logs in the area.
The thickness was found to vary from 300 feet
in the south to greater than 400 feet in the
northeast. From the map it was determined
that assuming a constant thickness of 375
feet for the Saginaw aquifer over the study
area was reasonable.
Leakage from the rivers was only assumed to
be important in those areas where the well
logs recorded continuous units of sand and
gravel to the bedrock.
Initial hydraulic heads were determined from
historical data in 1900 because at that time

there were no significant stresses imposed on

6)

20

the groundwater system (Wheeler, 1967).

Resistivity values for drift were computed
assuming an average thickness of 87.5 feet
(27 m), where resistivity is defined as the
thickness of a semi-confining layer divided
by its' hydraulic conductivity. An isopach
map of the thickness of the drift was
computer generated. From this map an average

drift thickness of 87.5 feet was chosen.

CALIBRATION

Calibration of the finite—element model was achieved
with application of the indirect approach (Neuman, 1973)
whereby input parameters are adjusted with repetitive
application of the model until computed heads matched the
observed field values (Freeze and Cherry, 1979). Over 60
simulation runs were made adjusting the resistivity of the
drift, the hydraulic conductivity of the aquifer, the
storage coefficient, and the boundary conditions to obtain
a solution which best approximated the actual hydrologic
conditions. Included in the Appendix are the various
pumping schedules used in calibrating the finite element
model (Table 1) and a sample computer run.

The model was initially calibrated to simulate 1945
steady-state conditions. Conductivity parameters were
assigned uniform values so as not to force artificial
complexities on the system. Therefore, the hydraulic
conductivites of the drift and the aquifer were assigned
the estimated values of 4.1*10-9 ft/sec (1.251'i10"7 cm/sec)
and 9.8"*10"6 ft/sec (3.0*10'4 cm/sec), respectively. The
pumpage in 1945 totaled 14 mgd over the study area.

Initial hydraulic heads were input as ¢ (head) zones
based on data for the year 1900 when the aquifer was
virtually undeveloped. The ¢ zones and hydraulic
conductivity of the aquifer were then adjusted singularly
keeping the hydraulic conductivity of the drift uniform

21

22

until the simulated potentiometric surface (Figure 7)
matched the historical potentiometric surface (Figure 8)
fairly well. Five initial ¢ zones proved to be a good
representation of initial conditions in 1900. In addition,
differentiating the hydraulic conductivity of the aquifer
into two zones (7.9“10'”5 ft/sec for the sandstone facies
and 2.3*10'5 ft/sec for the shale facies) as originally
anticipated, produced a good fit for the 1945 conditions.

After synthesis of the 1945 potentiometric surface a
transient analysis was performed for the period 1945-64.
In this analysis the initial ¢ zones were altered slightly
in addition to checking the model sensitivity to various
time steps. The values of drift conductivity and aquifer
conductivity as determined from the 1945 calibration were
used in this calibration. In addition, a storage
coefficient of 10-4 was assigned to the bedrock aquifer.
The model was not sensitive to the time step as values of
l, 2.5, and 5 years produced very little difference in the
resulting hydraulic heads for 1964. Therefore, a 1 year
time step was considered adequate. In 1964 27.5 mgd of
water was withdrawn from the aquifer. The fit obtained
from this calibration was reasonable considering that the
drift heterogeneity had not yet been addressed. Figure 9
shows the 1964 simulated potentiometric surface and Figure
10 shows the actual 1964 potentiometric surface.

With the 1964 match completed, a transient analysis

was performed for the period 1956-79. In this analysis the

23

Figure 7. Simulated 1945 Potentiometric Surface.

25

Figure 8. Actual 1945 Potentiometric Surface.

27

Figure 9. Simulated 1964 Potentiometric Surface.

29

Figure 10. Actual 1964 Potentiometric Surface.

 

 

 

 

/A

Q
T“\

 

 

 

 

 

 

 

31
drift hydraulic conductivity was addressed. Using the
values of the aquifer conductivity, storage coefficient,
time step, and 6 zones as determined from the 1945 and 1964
calibrations, the drift hydraulic conductivity was adjusted

until the simulated potentiometric surface matched the

actual 1979 conditions. Values of 1.5*10-8 to 3.2"10-8

7 to 9.8""10"7 cm/sec) were assigned to the

conductivity of the till; whereas, values of 5.3"*10-8 to

5 6 4

ft/sec (4.6*10-

2.3*10- ft/sec (1.6*10' to 7.0*10- cm/sec) were assigned
to the conductivity of the suspected high recharge areas.
From this analysis it was found that assigning a
differential conductivity value to the permeable sand and
gravel deposits of smaller areal extent produced a mounding
effect. Therefore, these deposits may be important for
recharge locally but on a regional level they were
determined to be insignificant. The area where the Red
Cedar and Grand Rivers converge, however, was determined to
be an important avenue for recharge. Conductivity values

7 6

of 1.8*10- ft/sec (5.5*10- cm/sec) for the river deposit

8 7

and 1.8*10- ft/sec (5.5*10- cm/sec) for the till produced
a good fit when compared with water well level data for
1979 from USGS observation wells (Huffman, 1980), Lansing
Board of Water and Light municpal wells, and State of
Michigan water well records. Figure 11 shows the 1979
simulated potentiometric surface and Figure 12 shows the
1979 actual potentiometric surface. The slight

irregularity in the contours of the simulated

32

Figure 11. Simulated 1979 Potentiometric Surface.

 

 

 

 

 

 

 

 

 

 

 

 

Figure 12.

34

Actual 1979 Potentiometric Surface.

_-A¢4—.—___.-_..-_ _—-.. _.___._.__._.

 

~_ -I-_‘.__A .—

36
potentiometric surface near the center of the cone of
depression can be attributed to artificial influence caused
by clustering pumping wells into a limited number of
pumping nodes. Pumpage for 1979 totaled 39.6 mgd in the
Lansing area.

With an acceptable match obtained, sensitivity
analyses were performed by varying the aquifer conductivity
and storage coefficient. Values ranging from 6.2*10"5 to
9.3*1o'5 ft/sec (1.9*1o‘3 to 2.8*1o'3 cm/sec) and 1.5*1o’5
to 3.1”!10"5 ft/sec (4.6*10'4 to 9.4*10'4 cm/sec) were
assigned to the sandstone facies and shale facies
respectively, to determine the best estimate of the aquifer
conductivity. The model was only slightly sensitive to
these changes and produced the best fit using conductivity
values of 7.9*10"5 ft/sec (2.41"10'3 cm/sec) for the
sandstone and 2.3*10'5 ft/sec (7.1"*10"4 cm/sec) for the
shale. The model was run with storage coefficients ranging
from 10'2 to 10-6, however, the simulations showed no
sensitivity to the storage coefficient. Therefore, a
storage coefficient of 10'4 was assumed adequate for
calibration. With the model calibrated to the 1979 data
the 1945—64 transient analysis and the 1945 steady state
analysis was again performed to check the calibration.
Figure 13 shows the final simulated 1945 potentiometric
surface and Figure 14 shows the final simulated 1964

potentiometric surface. In both cases the computed

37

Figure 13. Final Simulated 1945 Potentiometric

Surface.

 

 

 

 

 

 

 

 

w m fi/N

x we I
W \\\.
fly... x .\ _

 

 

 

39

Figure 14. Final Simulated 1964 Potentiometric

Surface.

potentiometric

measured surface.

41

surface matched' well with the

actual

DISCUSSION

Electric-Analog Model

 

Approximately 15 years ago, an electric-analog model
(Wheeler, 1967; Vanlier and Wheeler, 1968) was utilized in
the Lansing area to predict the change in the
potentiometric surface with time due to increased pumpage.
In the application of the analog model leakage was
contributed along the total length of the rivers. However,
even though the amount of pumpage input to the analog model
in 1975 (43 mgd) virtually matched that actually pumped in
1979 (39.5 mgd), drawdown in the downtown Lansing area in
1975 was overpredicted by 50 feet when compared to actual
1979 conditions (Huffman, 1980; Lansing Board of Water and
Light municipal well data). Furthermore, the drawdown in
the East Lansing area predicted by the 1975 analog
simulation was 83 feet greater than the actual 1976
conditions (Van Til, 1977).

To assess the hydraulic character of the drift,
leakage values were used in the analog analysis, whereas,
in the finite-element analysis hydraulic conductivity
values were used. However, the recharge values input to
the analog model appear to have been underestimated since
all other parameters used in the analog model are fairly

close to those used in the finite-element model.

42

43

Dewatering

 

In 1979 the potentiometric surface in the northern
Lansing area was reported to be 75 feet (23 m) below the
surface of the Saginaw aquifer, while in the southern
Lansing area, near the identified high recharge area, it
was reported to be as much as 25 feet (8 m) below the
surface of the aquifer (Keck Consulting; Huffman, 1980).
These data which were collected from wells that penetrate
at least 155 feet (47 m) into the aquifer would suggest
that the Saginaw aquifer is locally being dewatered.

In order to evaluate the significance of dewatering in
the Lansing area, a dewatering option (DEWA) was applied to
the finite element model. The potentiometric heads
calculated when using this option were found to be nearly
identical (3 5 ft) to those simulated when no dewatering
option was applied. .This lack of significant head
difference would suggest that perhaps the finite element
grid is not refined enough in the vicinity exhibiting
dewatered conditions to adequately simulate such
conditions. An alternative explanation is that the
observation well data may reflect mainly vertical pressure
differences due to shale lenses occuring within the
aquifer, rather than actual dewatering conditions.

To determine which of the two explanations is correct,
shallow observation wells would have to be drilled to test

if dewatering is occuring. In addition, a more refined

 

 

 

 

 

44
grid could be applied to the area exhibiting dewatered
conditions.

If dewatering is occuring in the Lansing area, water
may be drawn into the Saginaw aquifer from the underlying
Bayport Limestone. This could result in contamination of
the Saginaw aquifer since the Bayport Limestone produces
saline water in areas where it is overlain by the Saginaw

Formation (Vanlier, et al., 1973).

CONCLUSIONS

The methods which have been most commonly used to
determine the hydraulic conductivities of a leaky aretesian
system in a glaciated area are aquifer test analyses and
laboratory permeability studies: These techniques, though,
can not adequately determine the variabilty in hydraulic
conductivity of drift deposits that is expected over a
broad regional area. Application of a finite-element
model, however, can assess this variability and should
provide a more accurate description of the hydrologic
regime of a region characterized by heterogeneous deposits.

The hydraulic conductivity values for drift obtained
in this investigation can be compared to values obtained by
other researchers in the United States and Canada (Table
2). The conductivity values of drift as determined from
previous investigations range from 3.3“10-13 to 3.3*10-7
ft/sec (10"11 to 10'"5 cm/sec). The final values determined
from this research are 1.8*10-8 ft/sec (5.5*10-7 cm/sec)
for till deposits and 1.8*10'7 ft/sec (5.5*10'6 cm/sec) for
sand and gravel deposits. These hydraulic conductivity

values are clearly within the range defined by others.

45

 

 

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BIBLIOGRAPHY

Desaulniers, D. 8., Cherry, J. A., and Fritz, P., 1981.
Origin, Age and Movement of Pore Water in Argillaceous
Quaternary Deposits at Four Sites in Southwestern
Ontario. Journal of Hydrology 50, p.231-257.

Firouzian, A., 1963. Hydrological Studies of the Saginaw
Formation in the Lansing, Michigan Area - 1962.
Michigan State University, Unpub. M.S. Thesis.

Freeze, R. A. and Cherry, J. A., 1979. Groundwater.
Prentice-Hall, Bnglewood Cliffs, New Jersy.

Grisak, G. E. and Cherry, J. A., 1975. Hydrologic
Characteristics and Response of Fractured Till and Clay
Confining a Shallow Aquifer. Can. Geotech. Jrnl. 12,
p.23-43.

Haji-Djafari. S., 1982. Users' Manual, GEOFLOW Ground Water
Flow and Mass Transport Computer Program. D'Appolonia
Waste Management Services, Inc., Pittsburgh, Pa.

Huffman, G. C., 1980. Ground-Water Data for Michigan, 1979.

Kelly, W. A., 1936. The Pennsylvanian System in Michigan.
Michigan Dept. of Conservation, Geol. Surv. Div. Pub.
40, Geol. Ser. 34.

Long, D. T., Larson, G. J., and Slayton, D. S., 1981. Field
Evidence for Shale Membrane Filtration. Geol. Soc.
Amer. Abstr. with Programs 13, p.499.

Lovato, J. and Larson, G. J., 1979. The Geohydrologic
Characteristics of the Glacial Drift within the Water
Quality Management Project Site, Michigan State
University. Inst. of Water Research, Tech. Report No.
47.

Monaghan, G. W. and Ritter, M. H., 1982. Identification and
Verification of Natural Recharge Areas in Glaciated

Terrain. Geol. Soc. Amer. Abstr. with Programs 15, no.
7.

Neuman, S. P., 1973. Calibration of Distributed Parameter
Groundwater Flow Models Viewed as Multiple-Objective
Decision Process under Uncertainty. Water Resource
Res. 9, p.1006-1021.

Norris, S. E., 1962. Permeability of Glacial Till. Geol.

Sur. Prof. Paper 450-E, Short Papers in Geology,
Hydrology, Topography Art.224, p.E150-E151.

47

48
Stuart, W. T., 1945. Progress Report on the Ground-water
Resources of the Lansing Area, Michigan. Mich. Geol.
Sur. Prog. Rept. 13.

Vanlier, K. E. and Wheeler, M. L., 1968. Analog Simulation
of Ground-Water Development of the Saginaw Formation,
Lansing, Michigan area. Prepared by the U.S.G.S. in
cooperation with Tri-County Planning Commission and
Mich. Geol. Sur.

Vanlier, K. E., Wood, W. W., and Brunett, J. 0., 1973.
Water-Supply Development and Management Alternatives
for Clinton, Eaton, and Ingham Counties, Michigan.
Geol. Sur. Water-Supply Paper 1969.

Van Til, R. L., 1977. An Analysis of Hydrologic Simulation
and Groundwater Withdrawal Patterns in Eastern Lansing,
Michigan. Michigan State Univ., Unpub. M.S. Thesis.

Walton, W. C., 1960. Leaky Artesian Aquifer Conditions in
Illinois. Illinois State Water Survey Rept. of Invest.
39.

Wheeler, M. L., 1967. Electric Analog Model Study of the
Hydrology of the Saginaw Formation in the Lansing,
Michigan Area. Michigan State Univ., Unpub. M.S.
Thesis.

APPENDI X

49

Figure 6. Finite Element Grid with Pumping Centers.

Darkened circles represent pumping nodes.

 

 

 

 

3| 5'

 

 

59

 

 

 

 

 

 

 

 

51
TABLE 1
PUMPING SCHEDULES

1945 STEADY STATE
PUMPING SCHEDULE*

 

 

AMOUNT——
noon punpso
52 6.55*1o§
63 5.48*10
77 1.87*105
78 2.94*1o5
76 2.67*1o4
99 6.68*104

101 9.36*1o4

1945 - 1964 TRANSIENT ANALYSIS
PUMPING SCHEDULE*

 

 

NODE 1945-1949 1950-1954 1955-1959 1960-1964
52 6.95*1o5 6.82*105 6.15*105 6.1S*105
63 6.95*105 6.28*105 4.52*105 4.41*105
78 2.27*1o5 2.91*105 3.88*105 4.81*105
77 2.27*1o5 2.27*1o5 2.94*1o5 2.94*105
51 o 2.46*105 4.01*105 6.42*105
76 2.57*1o4 5.35*1o4 6.68*104 5.3s*1o4
99 6.68*104 9.36*1o4 1.60*105 2.14*1o5
101 1.47*105 2.54*105 6.68*104 6.68*104
102 o o 2.41*105 3.21*105
53 4.01*1o4 1.2o*105 1.47*105 1.74*105

52
TABLE 1 (Continued)

1964 - 1979 TRANSIENT ANALYSIS
PUMPING SCEEDULE*

 

 

noon 1956-1963 1964-1971 1972-1979
52 5.82*105 6.02*105 5.07*1o5
63 4.74*105 5.4o*1o5 5.32*1oS
78 4.57*105 5.75*1o5 5.52*1o5
77 3.18*105 2.81*105 3.26*105
51 4.99*105 6.68*105 5.59*1o5
76 5.35*1o4 6.6s*1o4 o

99 1.86*105 2.54*105 3.68*105
101 6.6s*1o4 3.02*1o4 1.66*105
102 2.86*105 4.o1*105 4.49*105
53 1.75*105 2.oo*105 o

98 o o 2.19*105
79 o o 3.44*105
so 0 o 5.37*1o5
55 o o 1.02*105
59 o o 3.22*105
31 o o 2.46*105
4o 0 o 6.42*1o4

* Values in ft3/day

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