66th Annual Tri-State
Geological Field Conference
September 24-25, 2005
Iowa State University
Ames, Iowa 50011
ROCKIN' IN THE HEARTLAND:
The Paleozoic/Quaternary Geology and
Hydrogeology of Central Iowa
Co-sponsored by
Department of Geological and Atmospheric Sciences
Iowa State University
and
National Association of
Geoscience Teachers
Iowa Groundwater
Association
ROCKIN' IN THE HEARTLAND:
The Paleozoic/Quaternary Geology and
Hydrogeology Of Central Iowa
66th Annual Tri-State Geological Field Conference
Iowa State University
Ames, Iowa
September 24-25, 2005
Co-sponsored by
Department of Geological and Atmospheric Sciences
Iowa State University
and
National Association of
Geoscience Teachers
Iowa Groundwater
Association
Trip Leaders:
Jane Pedrick Dawson
Matthew Graesch
Neal Iverson
Bill Simpkins
Carl F. Vondra
Field Trip Coordinators:
Jiasong Fang
Chris Harding
Iowa State University
Ames, Iowa 50011
TABLE OF CONTENTS
WELCOME
Carl E. Jacobson ……………………………………..……………………………… 1
STOP 1: WHATOFF'S BORROW PIT: TILLS, MORAINES, AND
DYNAMICS OF THE DES MOINES LOBE
Neal R. Iverson ……………..…………………………………………..…………… 3
STOP 1A: WHATOFF’S BORROW PIT: FRACTURES IN TILL OF THE DES
MOINES LOBE
William W. Simpkins and Martin F. Helmke ……………..……………………….. 12
STOP 2: ADA HAYDEN HERITAGE PARK: BURIED CHANNELS, QUARRY
LAKES, BEAVER DAMS, AND THE WATER SUPPLY OF AMES,
IOWA
William W. Simpkins and Evan G. Christianson ……………………………..…… 20
STOP 3: DOLLIVER PARK: DES MOINES CHEROKEE GROUP
Carl F. Vondra ……………………………………………………………………... 31
STOP 4: BJORKBODA MARSH: MORAINES, KAMES, AND DRAINS
Neal R. Iverson ……………………..……………………..…………………..…… 44
STOP 5: DES MOINES RIVER VALLEY: LATE-WISCONSINAN HISTORY
OF THE UPPER DES MOINES RIVER
Neal R. Iverson ………………………………………………..………..………….. 50
STOP 6: MISSISSIPPIAN AND PLEISTOCENE GEOLOGY AT MONTOUR
QUARRY
Jane Pedrick Dawson and Matt Graesch ……………………………………....…… 56
Appendix A: Maps
Map of Ames ……………...……………..………………………………….….….. 70
Location map of field trip Stops 1 to 5 ……………...……...……..…..…..….….. 71
Location map of field trip Stop 6 ……………...……………..…..………..….….. 72
Appendix B: 2005 Tri-State Road Log ……………………………….………….……….…. 73
WELCOME TO THE
66 ANNUAL TRI-STATE GEOLOGICAL FIELD CONFERENCE
th
by
Carl E. Jacobson
Department of Geological and Atmospheric Sciences
Iowa State University
Ames, Iowa
As is traditional for Tri-State, the
Sunday trip is only a half-day and this year
includes just a single stop, Montour Quarry
located east of Marshalltown. The quarry is
within Lower Mississippian limestones and
dolomites, and we will consider various
aspects
of
the
stratigraphy
and
sedimentology of these units. Also to be
examined are exceptional exposures of
Pleistocene till, loess, and paleosols exposed
by the quarrying operations. Following a
snack break, the trip ends here.
The Department of Geological and
Atmospheric Sciences at Iowa State
University warmly welcomes you to this
year’s Tri-State Geological Field Conference.
Tri-Sate was last hosted by our department in
1993, a year of record rainfall and
exceptional flooding within Iowa. With the
recent devastation induced by hurricane
Katrina, we are again reminded of the
enormous impact of atmospheric and Earth
processes on society. We hope for the
duration of this trip that the forces of nature
will be kinder to us.
We thank you for coming to Ames
and hope you enjoy the trip.
Our trip this year will cover a
diversity of geologic features in central Iowa.
The first stop on Saturday will be at
Whatoff’s Borrow Pit in southeast Ames,
where we will view the basal till of the Des
Moines Lobe and consider various processes
related to the geometry and movement of the
lobe. Next comes Ada Hayden Heritage Park
north of Ames, where we will delve into a
number of considerations pertaining to the
water supply for the City of Ames. From
there, we travel northwest to Dolliver State
Park to view Pennsylvanian sandstones and
have lunch. The middle to late afternoon
includes two stops. The first is at Bjorkboda
Marsh within the Altamont Moraine. Here
we will discuss processes of end-moraine
formation and the transformation of the Des
Moines Lobe landscape caused by
agricultural tile drainage. The final stop for
Saturday deals with the late Wisconsinan
evolution of the Des Moines River valley.
ACKNOWLEDGEMENTS
We are indebted to the Iowa
Groundwater Association and National
Association of Geoscience Teachers for their
generous financial support of the trip. Thanks
go to Tom Neumann, Director of the Water
and Pollution Control Department, for
pumping data on Ames’ well field; Harris
Seidel, retired Director of the Water and
Pollution Control Department, for the water
history of Ames; and Nancy Carroll, Director
of the Ames Parks and Recreation
Department, for access to Ada Hayden
Heritage Park. We are grateful to Marc
Whitman of Wendling, Inc. for providing
access to Montour Quarry. Numerous
discussions with Brian Gossman and Robert
Dawson of the Iowa DOT improved our
understanding
of
the
Mississippian
1
Mathison and Jason White contributed
substantially to the preparation of the
guidebook. Dave Flory was instrumental in
setting up the on-line registration system.
nomenclature problems in Iowa. DeAnn
Frisk is thanked for assistance with
innumerable aspects of logistics, registration,
and preparation of the guidebook. Mark
2
STOP 1 – WHATOFF’S BORROW PIT
TILLS, MORAINES, AND DYNAMICS OF THE DES MOINES LOBE
Neal R. Iverson
Iowa State University
lobe and the origin of one of its most
perplexing landforms: its minor moraines.
INTRODUCTION
The Des Moines Lobe was the largest
of several lobes of the Laurentide Ice Sheet
that extended into the mid-continent region
of North America near the end of the last
glaciation. At its maximum extent, ~13,800
radiocarbon years ago, the lobe was up to
250 km wide and covered greater than 105
km2 in southern Minnesota and north-central
Iowa (Fig. 1). The lobe deposited diverse
sediments, but only its basal till, seen here in
SEDIMENTS AND MORAINES
Several end moraines delineate
marginal positions of the Des Moines Lobe
in Iowa from ~12,000 to 14,000 radiocarbon
years before present (RCYBP). The most
prominent of these are the Bemis (~13,800
RCYBP), Altamont (~13,500 RCYBP), and
Algona moraines (~12,300 RCYBP),
characterized by broad,
concentric belts of
ridges and hummocky
topography, roughly 520 km wide (Fig. 1).
These moraines reflect
ice-marginal positions
that were sufficiently
steady for sustained
sedimentation in one
area.
Radiocarbon
dates
on
wood
summarized by Clayton
Ames
and Moran (1982)
suggest that the Algona
moraine was built by a
later advance of the
lobe, rather than during
Figure 1. Footprint of the Des Moines Lobe in Iowa with end moraines (Prior, a hiatus in the lobe’s
1991).
recession.
Whatoff’s borrow pit in Ames (Fig. 2), was
deposited over the full extent of the lobe. The
properties of this till and associated
sediments, together with reconstructions of
the lobe’s surface morphology, provide a
basis for inferring both the dynamics of the
Upland sediments of the Des Moines
Lobe in Iowa are collectively called the
Dows Formation (Kemmis et al., 1981; Bettis
et al., 1996). Two till members that constitute
the bulk of this formation have been
identified: the Alden Member, which is
3
and sorted sediment, and the matrix texture
of the diamicton layers is far more variable
than that of the massive Alden Member till
(Fig. 3). The degree of alignment of elongate
gravel and larger-sized clasts—the so-called
clast fabric—is also stronger in the Alden
Member than in the overlying till. The
stronger
fabric
and
homogeneous texture of the
Ice movement
Alden
Member
are
consistent with deposition
at the base of the lobe;
shear deformation there of
sediment-laden basal ice or
of the sediment bed would
Whatoff's pit
align clasts and mix
primary
heterogeneities.
Although the bulk density
of both till units is highly
variable, the mean density
of the Alden Member is
distinctly larger than that of
the Morgan Member (Fig.
4), consistent with its
interpretation as a basal till
compacted
under
the
Figure 2. Location of Whatoff's pit in southeast Ames. The Skunk River
valley is to the west and minor moraines are to the east.
weight of the glacier.
interpreted to be a basal till, and the
overlying Morgan Member, interpreted to be
a supraglacial till. Together they are
commonly 15-30 m thick, although the
Morgan Member is usually
present only
within end moraines. Both tills are yellowish
brown where they have been oxidized,
commonly to a depth of 3-5 m below the
ground surface; at greater depths where the
till is unweathered it is dark gray. Only the
Alden Member is present at Whatoff’s pit,
which lies about 7 km south of the Altamont
Moraine.
Thin (< 0.3 m) isolated layers of
sorted sand and gravel, visible in Whatoff’s
pit are present in the Alden Member till and
may reflect water movement and associated
sediment transport in zones where the base of
the lobe was separated from the bed. Such
zones are common beneath modern
temperate and polythermal glaciers in the
form of either discrete channels or
interconnected cavities (Paterson, 1994).
The Alden Member is interpreted to
be a basal till based primarily on three
properties that distinguish it from the Morgan
Member: its textural homogeneity, clast
fabric, and density (Lutenegger et al., 1983;
Kemmis, 1991). The Alden Member is a
loam with a very narrow range of matrix
texture (Fig. 3). At any one location this
range is even narrower. Isolated layers of
sand and gravel are present locally but are
volumetrically minor. In contrast the Morgan
Member consists of interbedded diamicton
Enigmatic landforms associated with
the Alden Member till are minor moraines
(also called corrugated or washboard
moraines), which are ubiquitous over much
of the upland regions of central and northcentral Iowa and are quite prominent near
Ames (e.g., Kemmis et al., 1981; Stewart et
4
of the Alden Member but can
also contain lenses of crossbedded sands.
DES MOINES LOBE
RECONSTRUCTIONS
Reconstructions of the
surface morphology of the DML
at its maximum extent provide
some basis for evaluating the
lobe’s dynamics and the origins
of
minor
moraines.
Reconstructions indicate that the
lobe was probably thin and gently
sloping (Mathews, 1974; Clark,
1992; Brevik, 2000; Hooyer and
Iverson, 2002).
Basal shear
stresses (down-slope component
of the glacier’s weight per unit
bed area) calculated from the
widely cited reconstruction of
Clark (1992) range from 0.7 to
4.3 kPa—one to two orders of
magnitude smaller than is typical
Figure 3. Textures of the Morgan Member and Alden Member tills for modern glaciers (Paterson,
(from Lutenegger et al., 1983).
al., 1988). Whatoff’s pit is excavated in
minor moraines at the edge of the Skunk
River valley; these moraines are welldeveloped east of the pit (Fig. 2). Although
they are barely discernable from the ground
they can be conspicuous on air photographs,
owing to differences in soil-moisture content
at the tops and bottoms if the ridges. The
moraines are subtle, roughly concentric
ridges with heights of 1-2 m and spacings of
30-180 m (average spacing of 105 m). Their
original relief may have been reduced by a
factor of 3 by slope processes and farming
(Burras and Scholtes, 1987). Their crests are
generally parallel to the Bemis Moraine (Fig.
1) and are perpendicular to the ice movement
direction. The moraines consist largely of till
Figure 4. Densities of the Morgan Member
(top) and Alden Member (bottom) tills (from
Lutenegger et al., 1983).
5
1994). To put these numbers in perspective,
0.7 kPa is the shear stress beneath a thick
dictionary resting on a 20° slope! Clark’s
(1992) reconstruction indicates that ice was
only ~ 80 m thick in the vicinity of Ames,
despite its location about 40 km upstream as
measured along a flow line from the lobe’s
margin at the Bemis Moraine. Clark (1992)
used the elevation of the Bemis Moraine,
together with flow-direction indicators (e.g.,
minor moraines) and the maxim that glaciersurface contours must lie perpendicular to
flow, to reconstruct the lobe.
A more recent reconstruction (Fig. 5)
considers the possibility that the Bemis
Moraine was ice-cored at the glacier
maximum, such that the current elevation of
the moraine underestimates the elevation of
the lobe at its edge (Hooyer and Iverson,
2000). Reconstructions based on two
different volumetric fractions (Cr in Fig. 5) of
debris in the Bemis Moraine, assuming it was
ice-cored, yielded a glacier thickness as
much as three times larger than that
estimated by Clark (1992) (~ 250 m thick at
Ames), with driving stresses as high as 15
kPa. Although these values are less extreme
than Clark’s, they do not alter the general
conclusion that the Des Moines Lobe was
unusually thin and gently sloping.
DYNAMICS OF THE DES MOINES
LOBE
Figure 5. (a) Reconstructed morphology and flow
lines of the Des Moines Lobe, based on the modern
elevation of the Bemis Moraines. Flowline A-A' is
the trace of the profiles shown in Figure 5b and 5c.
(b) Longitudinal ice surface profiles compared with
those of Clark (1992). (c) Basal shear stresses
calculated for the three reconstructions shown in (b)
(from Hooyer and Iverson, 2002).
6
The low shear stresses at the bed of
the Des Moines Lobe indicate that little of its
motion was by internal ice deformation
(Hooyer and Iverson, 2002). The velocity
due to internal shear deformation of ice,
averaged over the ice thickness, is Ui =
2AEτbnH/(n + 2), where τb is the basal shear
stress, H is the ice thickness, A is an icecreep parameter that is inversely proportional
to the effective ice viscosity, E is an
enhancement factor for soft Wisconsin-age
ice, and n is the stress exponent in the flow
law for ice (Paterson, 1994). Using
maximum values of τb and H (ice-cored
The rapid movement of the lobe into
Iowa, despite warm conditions, suggests that
the lobe was likely out of balance with the
climate, such that the lobe’s rapid advance
into Iowa was a pulse (or pulses), induced by
rapid slip, that could not be sustained. This
would have ultimately resulted in thinning
and eventual stagnation of the glacier, much
like that which occurs after a glacier surges
(e.g. Kamb et al., 1985). As a result many
have referred to “surges” of the Des Moines
Lobe and have called the lobe a “surging
glacier.” This usage should be avoided.
Surge-type glaciers undergo quasi-periodic
rapid motion or, more specifically, periods of
rapid motion with longer and relatively
uniform intervening periods of quiescence.
There is no evidence that the Des Moines
Lobe displayed such periodicity.
Bemis Moraine, Cd = 0.05) to maximize Ui
and reasonable values of A (7 x 10-15 kPa-3 s1
), n (3), and E (2.5) (Paterson, 1994), Ui is <
1.0 m yr-1. This upper velocity limit is
approximately three orders of magnitude less
than advance rates of the lobe, as inferred
from its radiocarbon chronology (Clayton
and Moran, 1982) after correction for
variable atmospheric production of C14
(Stuvier et al., 1998). Possible variability of
A and E (Paterson, 1994) falls well short of
accounting for this difference in speed.
If the glacier did not shear much
internally, it had to move primarily by
slipping over or shearing its bed, which in
turn implies that the bed was thawed (at the
pressure-melting temperature of the basal
ice). Considerable indirect evidence supports
these inferences. The degree of compaction
of the Alden Member till, as indicated by
preconsolidation stresses determined in
consolidation tests on intact till specimens (<
300 kPa for 13 samples tested, Hooyer and
Iverson, 2002), indicates that only a small
fraction of the weight of the glacier (< 15 %)
was supported by the grains of the till. The
rest of the weight was supported by
pressurized water in the till pores, indicating
that the basal water pressure was near the
ice-overburden pressure. Such glaciers that
are nearly “floating” are prone to rapid basal
movement (e.g., Englehardt and Kamb,
1997). Additional evidence for a thawed bed
includes fossil insects (Schwert and Torpen,
1996), herbaceous plants (Baker, 1996), and
trees (Bettis et al, 1996) found near the base
of the Alden Member. These fossils and the
general lack of evidence of permafrost
features in the area indicate that the lobe
advanced into a relatively warm (non-arctic)
climate. Moreover, tunnel valleys mapped
along the former margin of the lobe in
southern Minnesota indicate that there was
significant meltwater at the bed (Patterson,
1996).
A better modern analog may be some
of the Siple Coast ice streams in West
Antarctica (Clark, 1992). Like the Des
Moines Lobe these ice streams move fast,
despite low driving stresses, due to high
basal water pressure that “lubricates” the
glacier sole. They differ, however, in that
they are bounded on their sides by much
slower moving ice, which supports most of
the down-slope component of their weight.
The Des Moines Lobe, at least in Iowa, had
no such lateral support (Fig. 1) and hence
may been intrinsically more susceptible to
rapid motion than the Siple Coast ice
streams.
DID THE DES MOINES LOBE MOVE
BY DEFORMING ITS BED?
If the Des Moines Lobe slipped
rapidly at its base, where exactly, relative to
the glacier sole, did that motion occur?
Some of have argued that lobes along the
southern margin of the Laruentide ice sheet,
including the Des Moines Lobe, moved
primarily by shearing their water-saturated
7
till beds (Alley, 1991; Clark, 1994). An
implication of this hypothesis is that the till
of the Alden Member may have been
transported into Iowa largely beneath the
glacier, rather than within it.
fabric formed by the alignment of gravelsized elongate clasts becomes steady and
strong at low strains (< 25) (Hooyer and
Iverson, 2000). Steady-state S1 eigenvalues,
a measure of the degree of alignment of the
long axes of clasts that can vary from 0.33
(uniform distribution) to 1.0 (perfectly
aligned), were 0.78-0.87. If the beddeformation hypothesis is correct, the till of
the Alden Member should display similarly
strong fabrics.
Figure 6. Clast-fabric stereograms for the Alden
Member till at various locations. (b) S1 and S3
eigenvalues for the Alden Member till, for till
deformed in ring-shear tests (Hooyer and Iverson,
2000) and for till of selected drumlins.
Measurements of clast fabrics along
the centerline of the Des Moines Lobe (Fig.
6a), including measurements in Whatoff’s pit
(“Ames” in Fig. 6), yielded eignenvalues of
0.44-0.66, far smaller than the steady-state
values from ring-shear experiments (Fig. 6b).
Hooyer and Iverson (2002) argued, on that
basis, that although the till likely underwent
some shear during deposition from ice, it
probably did not shear pervasively over its
thickness to the high strain required of the
bed-deformation hypothesis.
However,
earlier measurements of clast fabric in
Whatoff’s pit yielded a mean eigenvalue of
0.72 (Stewart et al., 1987), a value less than
those expected at high strains but
significantly larger than the values of Hooyer
and Iverson (2002). Moreover, unpublished
data gathered by Kemmis (1991) at three
sites within end moraines indicate S1 = 0.680.91.
If during the full duration of the Des
Moines Lobe in central Iowa the glacier
essentially rode “piggy-back” on till shearing
beneath the ice, then the bed would have
been sheared to very high strains (> 1000,
such that glacier displacement was at least
1000 times greater than the bed thickness).
The alignment of clasts in the till should,
therefore, reflect this high strain. Laboratory
ring-shear experiments, in which tills were
sheared to strains up to 475 indicate that the
One reasonable interpretation of this
apparent variability in fabric strengths among
different studies is that deformation of the
bed was highly heterogeneous. A second,
more unsettling interpretation is that clastfabric measurements require so much human
subjectivity and hence uncertainty, that such
measurements are not very meaningful. A
controlled study that isolates the uncertainty
of clast-fabric measurements, although not
very glamorous (or fundable), is probably
needed more at present than additional field
8
melting ice (Paterson, 1994). Vigorous water
flow also is consistent with the cross-bedded
sands observed in the moraines. Some of
these sands have been described in Whatoff’s
pit (Stewart et al., 1987) but are no longer
visible.
measurements that continue to neglect this
uncertainty.
GENESIS OF MINOR MORAINES
The strong likelihood that the Des
Moines Lobe moved rapidly into Iowa under
very low basal shear stresses and high basal
water pressures has important implications
for models of minor moraine formation.
Such glaciers that advance out of balance
with the climate due to basal lubrication
commonly undergo transient extending flow,
even in their ablation areas where a glacier
that is in balance with the climate would
normally undergo compressive flow. The
extending flow of a rapidly sliding glacier is
somewhat analogous to flow of ice shelves
(parts of glaciers that float on the ocean),
which can extend rapidly under their own
weight owing to a lack of basal slip
resistance.
ACKNOWLEDGEMENTS
I thank Tom Hooyer whose
dissertation research at ISU provided the
basis for much of this section.
REFERENCES
Alley, R.B. 1991. Deforming-bed origin for
southern Laurentide till sheets? J.
Glaciol., 37(125), 67-76.
Baker, R.G. 1996. Pollen and plant
macrofossils. In Bettis, E.A., D.J. Quade
and T.J. Kemmis, eds., Hogs, Bogs, and
Logs:
Quaternary
deposits
and
environmental geology of the Des Moines
Lobe. Iowa Department of Natural
Resources, Guidebook Series, 18, 105109.
Building on an earlier hypothesis by
Kemmis et al. (1981), Stewart et al. (1987)
appealed to extending flow to explain the
minor moraines near Ames (Fig. 2). These
authors attributed them to preferential
deposition of basal till and fluvial sediment
in crevasses that opened in the basal ice. If
ice was extending parallel to the direction of
flow, these crevasses would have opened
perpendicular to the flow direction,
consistent with the orientations of minor
moraines. A possible problem with this
hypothesis is that it is difficult for crevasses
to open in basal ice because tensile stresses
that promote crevasse opening must be very
high to overcome the basal confining
pressure that tends to squeeze crevasses shut.
This is true even for glaciers as thin as 100
m. A potential resolution to this problem is
that crevasses may have opened under the
combined effect of high basal water pressure
and vigorous water flow. The latter dissipates
heat, which can hold basal cavities open by
Bettis, E.A., D.J. Quade and T.J. Kemmis,
1996. Overview, In Bettis, E.A., D.J.
Quade and T.J. Kemmis, eds., Hogs,
Bogs, and Logs: Quaternary deposits and
environmental geology of the Des Moines
Lobe. Iowa Department of Natural
Resources, Guidebook Series, 18, 1-79.
Brevik, E. 2000. Limits to ice thickness in
Iowa during the Late Wisconsian. J. of
the Iowa Acad. Sci., 107(2), 46-50.
Burras, C.L. and W.H. Scholtes, 1987. Basin
properties and post-glacial erosion rates
of minor moraines in Iowa. Soil Sci.
Amer. J., 51(6), 1541-1547.
9
Kemmis, T.J., G.R. Hallberg, and A.J.
Lutenegger,
1981.
Depositional
Environments of Glacial Sediments and
Landforms on the Des Moines Lobe,
Iowa.
Iowa
Geological
Survey
Guidebook Series Number 6.
Clark, P.U. 1992. Surface form of the
southern Laurentide Ice Sheet and its
implications to ice-sheet dynamics. Geol.
Soc. Am. Bull., 104(5), 595-605.
Clark, P.U. 1994. Unstable behavior of the
Laurentide Ice Sheet over deforming
sediment and its implications for climate
change. Quat. Res., 41(1), 19-25.
Lutenegger, A.J., T.J. Kemis, and G.R.
Hallberg, 1983. Origin and properties of
glacial till and diamictons. Special
Publication on Geological Environment
and Soil Properties, American Society of
Civil
Engineers,
Geotechnical
Engineering Division, 310-331.
Clayton, L. and S.R. Moran, 1982.
Chronology
of
late
Wisconsinan
Glaciation in middle North America.
Quat. Sci. Rev., 1, 55-82.
Mathews, W.H. 1974. Surface profiles of the
Laurentide Ice Sheet in its marginal
areas. J. Glaciol., 13(7), 37-43.
Engelhardt, H. and B. Kamb, 1997. Basal
hydraulic system of a West Antarctic ice
stream: constraints from borehole
observations. J. Glaciol., 43(144), 207230.
Paterson, W.S.B. 1994. The Physics of
Glaciers. New York, Pergamon Press.
Hooyer, T.S. and N.R. Iverson, 2000. Clastfabric development in a shearing granular
material: Implications for subglacial till
and fault gouge. Geol. Soc. Am. Bull.,
112(5), 683-692.
Patterson, C.J. 1996. The glacial geology of
southwestern Minnesota with emphasis
on the deposits and dynamics of the Des
Moines Lobe. (Ph.D. thesis, University of
Minnesota.)
Hooyer, T.S. and N.R. Iverson (2002), Flow
mechanism of the Des Moines lobe of the
Laurentide ice sheet, J. Glaciol., 48(163),
575-586.
Prior, J.C. 1991. Landforms of Iowa, 153 pp.,
University of Iowa Press, Iowa City.
Schwert, D.P. and H.J. Torpen, 1996. Insect
remains: a faceted eye’s perspective on
the advance of the Des Moines Lobe into
north-central Iowa. In Bettis, E.A., D.J.
Quade and T.J. Kemmis, eds., Hogs,
Bogs, and Logs: Quaternary deposits and
environmental geology of the Des Moines
Lobe. Iowa Department of Natural
Resources, Guidebook Series, 18, 99104.
Kamb, B. and 7 others, 1985. Glacier surge
mechanism: 1982-1983 surge of
Variegated Glacier, Alaska. Science,
227(4686), 469-479.
Kemmis, T.J. 1991. Glacial landforms,
sedimentology,
and
depositional
environments of the Des Moines Lobe,
northern Iowa. (Ph.D. thesis, University
of Iowa.)
10
Stuvier, M. and 9 others, 1998.
INTERCAL98
radiocarbon
age
calibration,
24,000-0
cal
BP.
Radiocarbon,
40(3),
1041-1083
Stewart, R.A., D. Bryant, and M.J. Sweat,
1988. Nature and origin of corrugated
ground moraine of the Des Moines Lobe,
Story County, Iowa. Geomorphology, 1,
111-130.
11
STOP 1A – WHATOFF’S PIT
FRACTURES IN TILL OF THE DES MOINES LOBE
William W. Simpkins and Martin F. Helmke
Iowa State University and West Chester University
freeze/thaw, and lateral unloading, may also
play a role (Boulton and Paul, 1976;
Mitchell, 1976; Connell, 1984). Helmke
(2003) suggested that the fractures in till here
and at his trench site south of Ames are shear
fractures because their primary orientations
are at 45 to 90° angles to the ice-flow
direction. Fractures in Pre-Illinoian till show
three main orientations indicative of the
predominance of 6-sided (desiccation?)
polygons.
INTRODUCTION
Tills are considered by many to be
impermeable and thus should prevent vertical
and horizontal transport of point- and
nonpoint-source contaminants.
However,
studies in Iowa have shown that aquifers
underlying till are contaminated and that
streams contain high concentrations of
nutrients and pesticides (e.g., Kolpin et al.,
1995; Burkart et al., 2004). We observed
fractures in till of the Des Moines Lobe at the
Whatoff Pit (Fig. 1) in the early 1990s, at the
suggestion of George Hallberg of the Iowa
Geological Survey. Martin Helmke later
obtained some of the test samples for his
Ph.D. research on fractures from this pit. His
dissertation (Helmke, 2003) showed that
groundwater flow and contaminant transport
in the till are controlled by these fractures.
OCCURRENCE AND FORMATION OF
FRACTURES IN TILL
Fractures or zones of preferential
flow in till have been reported previously in
Iowa near Iowa City (Kemmis et al., 1992)
and elsewhere in the U.S. (Connell, 1984;
Simpkins and Bradbury, 1992; Brockman
and Szabo, 2000), Canada (Keller et al.,
1988; McKay et al., 1993a,b), and Denmark
(Klint and Gravensen, 1999). Consolidation,
unloading during glaciation, subglacial
shearing, and stresses generated by glacer
flow have been proposed as mechanisms for
fracture formation (Boulton, 1970; McGown
et al., 1974; Johnson, 1983; Connell, 1984;
Feeser, 1988). Secondary processes, such as
chemical alteration, desiccation, syneresis,
Figure 1. Exposure of Alden Member till at the
Whatoff Pit (c. 1990) showing extensive Festained fracture surfaces (Lee, 1991).
12
686 rural wells and revealed that 35 percent
of the state’s shallow groundwater was
contaminated by NO3-N concentrations
above the US EPA MCL of 10 mg/L NO3-N,
and that 18 percent contained detectable
concentrations of herbicides.
Aquifers
confined by thinner (and presumably
fractured) till units (<15 m) showed
significantly more nitrate (35.1 vs. 12.8
percent) and pesticide (17.9 vs. 11.9 percent)
detections than those confined by thicker
(>15 m) and less fractured till units (Kross et
al., 1990).
WHY ARE FRACTURES IN TILL
IMPORTANT?
Fractures
create
preferential
flowpaths that promote greater velocities in
till than would otherwise be expected under
porous media assumptions (Freeze and
Cherry, 1979; Grisak and Pickens, 1980). A
velocity increase occurs by increasing bulk
hydraulic conductivity (Kb) and reducing
effective porosity (ne). The Kb of a fractured
till is typically one to 3 orders of magnitude
greater than Kb for an unfractured till (Keller
et al., 1989), whereas fracture porosity (nf)
may be one to 4 orders of magnitude less
than the total porosity (nT) of till (McKay et
al., 1993a; Jørgensen et al., 1998).
Advective velocity of solutes in fractured
systems may be estimated by the average
linear velocity equation:
V =
Kbi
nf
INVESTIGATIONS OF FRACTURED
TILL IN IOWA
Helmke (2003) used large till
columns to investigate contaminant transport
in till from the Des Moines Lobe, the Iowan
Erosion Surface, and the Southern Iowa Drift
Plain (landform regions) in Iowa. The study
site in the Des Moines Lobe was located
within the Walnut Creek watershed, about 7
km south and slightly west of Whatoff’s Pit.
The surficial deposit at the site is the Alden
Member till of the Dows Formation and is
identical in texture and bulk density to the till
at the pit. Previous investigations in the
Walnut Creek watershed (visited during the
1993 Tri-State Field Conference) revealed
that the till is extensively fractured (Eidem et
al. 1999).
To provide large-diameter
columns for tracer tests, a 4-m-deep trench
was excavated using a backhoe to provide
access to the till. The trench was carved
using a bench and tier method to provide
multiple faces for fracture mapping and to
ease column collection. Fractures were
identified as planes with iron-oxide staining
or as leached zones in the till and were
mapped using sheets of clear acetate on both
vertical and horizontal faces in the trench.
Fracture strike and dip were measured using
a Brunton compass.
[1]
where V is velocity and i is the hydraulic
gradient. Fluid velocities up to 200 m/day
have been calculated for fractured till using
Eq. [1](Jørgensen et al., 1998). Fortunately,
the processes of matrix diffusion, sorption,
and
degradation
typically
retard
contaminants as they pass through fractured
till, allowing only a small percentage of a
solute to travel at velocities calculated by Eq.
[1] (Freeze and Cherry, 1979).
There is now good evidence that
fractures allow contaminants to move
through till more than previously thought.
Vertical transport of contaminants has been
documented in till in Canada (McKay and
Fredericia, 1995) and Denmark (Jørgensen
and Fredericia, 1992; Jørgensen and Spliid,
1992) as well as lateral transport to streams
(D’Astous et al., 1989; Herzog et al., 1989;
McKay et al., 1998). At a larger scale, the
Iowa State Rural Water Survey (SWRL) of
the late 1980s sampled groundwater from
13
The excavation revealed that the till
contains numerous sub-horizontal and subvertical fractures from ground surface to the
base of the pit. Fracture spacing ranged from
< 2 cm near the surface to approximately 4.6
cm at a depth of 4 m. The most prominent
fractures were observed below 3 m depth
where the till is partially weathered. At this
depth, the fracture surfaces were stained
reddish brown (Munsell color: 10YR 5/8) in
contrast to the olive-brown (Munsell color:
2.5Y 5/4) till matrix. The fractures were
primarily sub-vertical and oriented northeast
to southwest and northwest to southeast. The
average fracture spacing at a depth of 3.3 m
was 4.3 cm and the fracture density was 260
fractures/m2 (Fig. 2).
Figure 3. Photograph of the till column
(43-cm diameter and 45-cm length) prior
to encasement. Sub-vertical, Fe-stained
fracture surfaces are prominent. Putty
knife for scale (Helmke et al., in press).
The tracer solution was introduced to
each column under a constant hydraulic
gradient using a Mariotte bottle (Helmke et
al., 2005).
Diffusion coefficients were
determined directly (Helmke et al., 2004). In
a separate study, four fracture transport
models (Fig. 4) – the Mobile-Immobile
Model (MIM), Parallel-plate Discrete
Fracture Model (PDFM), and Stochastic and
Deterministic Discrete Fracture Models
(DFMs) – were used to simulate transport of
conservative solutes through the till (Helmke
et al., in press). Tranport of nitrate and
atrazine was also modeled using the MIM
(Helmke et al., 2005).
Intact columns of till, 43 cm diameter
by 45 cm length, were carved from steps in
the excavation trench using a shovel and
putty knife (Fig. 3). Additional details on the
core preparation process are given in Helmke
(2003). Five tracers were used in the till
column experiments: KBr [potassium
bromide], PFBA [pentafluorobenzoic acid],
PIPES [1, 4-piperazinediethanesulfonic acid
disodium salt], KNO3 [potassium nitrate],
and
atrazine
[6-chloro-N-ethyl-N’-(1methylethyl)-1,3,5-triazine-2,4-diamine].
Figure 2. Plan-view map of fractures at 3.3
m depth in a pit near Ames. Fractures are
predominantly sub-vertical in orientation and
trend NW-SE and NE-SW. Ice-flow direction
is from the NW (Helmke, 2003).
Figure 4. (a) Parallel-plate Discrete Fracture
Model, (b) Stochastic Discrete Fracture Model ,
and (c) Deterministic Fracture Model
representations of the till column (Helmke et al.,
in press).
14
Differences of BTC morphology
among the conservative tracers (Br, PFBA,
Breakthrough curves (BTCs) showed
that solute transport in the till is controlled by
macropores or fractures (Fig. 5). In the
absence of such features, breakthrough
should have occurred after one pore volume
(PV) had passed through each column.
Instead, breakthrough occurred prior to one
PV.
Measurable concentrations of the
conservative tracers (Br, PFBA, and PIPES)
appeared in the column effluent (C/C0 >
0.02; the instrument detection limit) at least
10 times faster than 1 PV. Helmke et al.
(2005) defined a velocity of Br, or VBR,
corresponding to this time of first arrival. As
a result, VBR velocities of 658, 106, and 9.7
m/d were calculated for DML-1, -2, and -3,
1.0
V (m/d)
10 -4 10
0
Relative Concentration (C/C0)
1.0
0.6
0.5
1.0
1.5
2.0
2.5
Depth (m)
10-1
DML-3
10
0
10DML-1
10 1
DML-2
H2
10 2
H1
AO
15
AT
20
25
30
ALB
VPM
0.2
0.0
10-2
5
DML-1 Column
PV = 1.95 d
0.6
-3
3.0
DML-2 Column
PV = 0.69 d
VBr
Figure 6. Plot of velocity (log scale) versus
depth for eight till columns from Iowa. Time of
first arrival of the Br tracer (VBR, filled arrows)
and for plug flow through the bulk porosity of
the sample (VPM, open circles) are shown
(Helmke et al., 2005).
and PIPES) provide additional evidence
of macropore- or fracture-controlled
0.2
solute transport. Matrix diffusion, the
0.0
0.5
1.0
1.5
2.0
process whereby solutes are exchanged
1.0
between the matrix (immobile region)
DML-3
Column
and macropore/fracture (mobile region)
0.6
PV = 20.3 d
due to a concentration gradient,
0.2
effectively retards solutes as they pass
through the column. If matrix diffusion
0
10
20
30
40
50
60
70
is occurring, the rate at which solutes
Days
increase in concentration during the
rising limbs of BTCs should be inversely
PFBA PIPES NO3 Atrazine
MIM Fit
Br
proportional to the respective D0 values
Figure 5. (a) Observed and modeled breakthrough curve
(i.e., PIPES will increase in concentration
from the DML-1, -2, and -3 columns. Dashed vertical line
first, followed by PFBA and then Br).
indicates time for 1 pore volume (PV) to pass through the
There should be a similar separation of
column (Helmke et al., 2005).
the solute concentrations during the
falling limbs, or tails, of the BTCs
respectively. These are among the highest
(Moline
et al., 1997; Gwo et al., 1998). This
velocities recorded in the overall study (Fig.
phenomenon
occurs (Fig. 5), although it is
6).
more pronounced in the longer experiments
(DML-2 and DML-3 columns). Additional
15
retardation or complete degradation in
columns from partially weathered and
unweathered till. This may be due to the
increase in organic carbon in this material
and the onset of denitrification in the matrix
(Parkin and Simpkins, 1995). Fracturecontrolled transport of atrazine occurs in
weathered till, but due to sorption its velocity
is retarded with respect to nitrate. Velocities
in unweathered fractured till may be low
enough for complete degradation of atrazine
to occur prior to entering an aquifer or
stream.
evidence of matrix diffusion is shown by the
response to rinsing the columns, where low
concentrations of solutes were detected (socalled “elongated tails”) even when rinsed
for twice the time of injection. Mass balance
calculations indicate that 15 to 35 percent of
the conservative solutes remained in the
shallow columns. Hence, nonpoint-source
contaminants could be stored in the matrix
for later release into the environment
(Helmke et al., 2005).
Nitrate behaved as a conservative
tracer during short-term experiments (fewer
than 3 days) in the shallow columns (DML-1
and DML-2) and in a non-conservative
manner during longer-term experiments for
deeper columns (DML-3). The nitrate BTCs
from the DML-1 and DML-2 column
experiments were nearly identical to the BrBTCs (Fig. 5). This was not the case for the
DML-3 column, where the relative
concentration of nitrate remained below 0.05
for the duration of the experiment (Fig. 5),
suggesting that it degraded quickly
(presumably by denitrification). Atrazine
behaved non-conservatively in all columns.
Tailing phenomenon suggests that sorption,
rather than degradation, is the main process
acting to retard atrazine, particularly in some
of the longer-term experiments (Helmke et
al., 2005).
Model simulations indicate that the
close fracture spacing of the till allowed
diffusive equilibrium to occur between the
fractures and matrix over a relatively short
time period (several weeks). This effect
caused the system to behave similar to an
Equivalent Porous Medium (EPM), even
though a dye trace study (Fig. 7) showed that
flow occurred primarily through fractures.
Halos surrounding the fractures provide
evidence of rapid matrix diffusion (Helmke
et al., 2005). Thus, an EPM approach could
ENVIRONMENTAL IMPLICATIONS
OF FRACTURES IN TILL
Experiments with three conservative
solutes (KBr, PFBA, and PIPES) and two
non-conservative solutes (nitrate and
atrazine) showed that transport through till of
the Des Moines Lobe is controlled by
fractures. The potential for fractures to
transport nitrate and atrazine and impact
water quality varies with depth. Nitrate was
unaltered in shallowest columns from
weathered till, but showed evidence of
N
20 cm
Figure 7. Plan view of a slice taken from the
center of the DML-3 column showing
fractures (dark lines) and extent of FDC
Brilliant Blue Dye no. 1 (gray zones) at a
depth of 3.65 m (Helmke et al., 2005).
16
geotechnical properties of glacial tills.
Quaterly J. of Eng. Geol., 9(3), 159-194.
be used in these deposits for large spatial or
temporal scales.
However, for short
timescales or situations where fracture
spacing is large with respect to the scale of
investigation, EPM assumptions would
clearly be inappropriate. Additionally, for
cases where boundary conditions are
transient (e.g., recharge or remediation),
fractured systems would remain in a state of
constant disequilibrium and would require
models that simulate diffusion explicitly.
Brockman, C. S. and J. P. Szabo. 2000.
Fractures and their distribution in the tills
of Ohio. Ohio J. of Sci., 100(3/4), 39-55.
Burkart, M.R., W.W. Simpkins, W.J.
Morrow, and J.M. Gannon. 2004.
Occurrence of total dissolved phosphorus
in unconsolidated aquifers and aquitards
in Iowa. JAWRA, 40(3), 827-834.
In summary, thin, weathered till units
will not protect underlying or adjacent
aquifers (e.g., alluvial aquifers) and surface
waters from contamination. In addition,
matrix diffusion may store nonpoint-source
contaminants in the till matrix for later
release, providing a legacy of past
contamination activities well into the future
(Rodvang and Simpkins, 2001). Thus, we
believe that the concept “till impermeability”
for protection of aquifers and surface waters
in Iowa should be re-evaluated.
Connell, D.E.
1984.
Distribution,
characteristics, and genesis of joints in
fine-grained till and lacustrine sediments,
eastern and northwestern Wisconsin.
M.S thesis, University of WisconsinMadison.
D’Astous, A.Y., W.W. Ruland, J.R.G. Bruce,
J.A. Cherry, and R.W. Gillham. 1989.
Fracture effects in the shallow
groundwater zone in weathered Sarniaarea clay. Canadian Geotech. Jour., 26,
43-56.
ACKNOWLEDGEMENTS
We thank former ISU hydrogeology
graduate students Beth Johnson and Jim
Eidem for doing the fieldwork that paved the
way for this research. We thank ISU
Veterinary Farm for allowing us to dig up
part of a pasture to retrieve the till core. We
also acknowledge the helpful insights of our
collaborator, Dr. Robert Horton, in the
Agronomy Department at ISU.
Eidem, J. M., W.W. Simpkins, and M.R.
Burkart. 1999. Geology, groundwater
flow, and water quality in the Walnut
Creek watershed. J. Environ. Qual., 28,
60-69.
REFERENCES
Freeze, R.A. and J.A. Cherry. 1979.
Groundwater. Prentice Hall, New York.
Feeser, V. 1988. On the mechanics of
glaciotectonic contortion of clays.
Glaciotectonics: Forms and Processes,
63-76.
Boulton, G.S. 1970. The deposition of
subglacial and meltout tills at the margin
of certain Svalbard glaciers. J. Glaciol.,
9(56), 231-245.
Grisak, G.E. and J.F. Pickens. 1980. Solute
transport through fractured media: 1. The
effect of matrix diffusion. Water Resour.
Res., 16, 719-730.
Boulton, G.S. and M.A. Paul. 1976. The
influence of genetic processes on some
17
Jørgensen, P. R. and N. H. Spliid. 1992.
Mechanisms and rates of pesticide
leaching in shallow clayey till In
European Conference on Integrated
Research for Soil and Sediment
Protection and Remediation. MECC,
Maastricht, the Netherlands.
Gwo, J.P., R.O’Brien, and P.M. Jardine.
1998. Mass transfer in structured porous
media: embedding mesoscale structure
and microscale hydrodynamics in a tworegion model. J. of Hydrol., 208, 204222.
Helmke, M.F. 2003. Studies of solute
transport through fractured till in Iowa.
Ph.D. dissertation. Iowa State University.
Ames, Iowa.
Jørgensen, P.R., L.D. McKay, and N.ZH.
Spliid. 1998. Evaluation of chloride and
pesticide transport in a fractured clayey
till using large undisturbed columns and
numerical modeling. Water Resour. Res.,
34, 539-553.
Helmke, M.F., W.W. Simpkins, and R.
Horton.
2004.
Experimental
determination of effective diffusion
parameters in fractured till. Vadose Zone
Journal, 3, 1050-1056
Keller, C.K., G. van der Kamp, and J.A.
Cherry. 1988. Hydrogeology of two
Saskatchewan tills, I. Fractures, bulk
permeability, and special variability of
downward flow. J. of Hydrol., 101, 97121.
Helmke, M.F., W.W. Simpkins, and R.
Horton. 2005. Fracture-controlled
transport of nitrate and atrazine in four
Iowa till units. J. of Environ. Qual., 34,
227-236.
Keller, C.K., G. van der Kamp, and J.A.
Cherry. 1989. A multiscale study of the
permeability of a thick clayey till. Water
Resour. Res,. 25(11), 2299-2317.
Helmke, M.F., W.W. Simpkins, and R.
Horton. in press. Simulating conservative
tracers in fractured till under realistic
timescales. Ground Water.
Kemmis, T.J., E.A. Bettis III, and G.R.
Hallberg. 1992. Quaternary geology of
Conklin Quarry. Iowa Department of
Natural Resources Guidebook Series no.
13.
Herzog, B.L., R.A. Griffin, C.J. Stohr, L. R.
Follmer, W.J. Morse, and W.J. Su. 1989.
Investigation of failure mechanisms and
migration of organic chemicals at
Wilsonville, Illinois. Ground Water
Monitoring and Review, 9, 82-89.
Klint, K. E. S. and P. Gravensen. 1999.
Fractures and biopores in Weichselian
clayey till Aquitards at Flakkebjerg,
Denmark. Nordic Hydrol., 30(4/5), 267284.
Johnson, M.D. 1983. The origin and
microfabric of Lake Superior red clay. J.
Sed. Petrol., 53(3), 859-873.
Kolpin D.W., S.J. Kalkhoff, D.A. Goolsby,
D.A. Sneck-Fahrer, and E.M. Thurman.
1995. Occurrence of selected herbicides
and herbicide degradation products in
Iowa’s ground water, 1995. Ground
Water, 35(4), 679-688.
Jørgensen, P.R. and J. Fredericia. 1992.
Migration of nutrients, pesticides and
heavy metals in fractured clayey till.
Geotechnique, 42, 67-77.
18
McKay, L.D., D.J. Balfour, and J.A. Cherry.
1998. Lateral chloride migration from a
landfill in a fractured clay-rich glacial
deposit. Ground Water, 36, 988-999.
Kross, B.C., G.R. Hallberg, D.R. Bruner,
R.D. Libra, K.D. Rex, L.M.B. Weih,
M.E. Vermace, L.F. Burmeister, N.H.
Hall, K.L. Cherryhomes, J.K. Johnson,
M.I. Selim, B.K. Nations, L.S. Seigley,
D.J. Quaide, A.G. Dudler, K.D. Sesker,
M.A. Culp, C.F. Lynch, H.F. Nicholson,
and J. Hughes. 1990. The Iowa StateWide Rural Well-Water Survey, Water
Quality Data: Initial Analysis. Iowa
Department of Natural Resources
Technical Information Series 19.
Lee, S-H. 1991. Genesis and distribution of
fractures in late-Wisconsin till of the Des
Moines Lobe in central Iowa. Iowa State
University, unpubl. M.S. thesis, 85 p.
Mitchell, J.K. 1976. Fundamentals of Soil
Behavior. New York. John Wiley and
Sons. 422 p.
McGown, A., A. Alvidar-Sali, A.M.
Radwan. 1974. Fissure patterns and slope
failures in till at Hurlford, Ayrshire.
Quarterly J. of Eng. Geol., 7(1), 1-26.
Moline, G.R., C.R. Knight, and R.Ketcham.
1997.
Laboratory measurement of
transport processes in a fractured
limestone/shale saprolite using solute and
colloid tracers. Geological Society of
America Absts. with Progs., 29(6), 370.
McKay, L.D. and J. Fredericia.
1995.
Distribution, origin, and hydraulic
influence of fractures in a clay-rich
glacial deposit. Canadian Geotech. J.,
32, 957-975.
Parkin, T.B. and Simpkins, W.W. 1995.
Contemporary groundwater methane
production from Pleistocene carbon. J.
Environ. Qual., 24(2), 367-372.
McKay, L.D., J.A. Cherry, and R.W.
Gillham. 1993a. Field experiments in a
fractured clay till: 1. Hydraulic
conductivity and fracture aperture. Water
Resour. Res., 29, 1149-1162.
Rodvang, S.J. and W.W. Simpkins. 2001.
Agricultural contaminants in Quaternary
aquitards: a review of occurrence and fate
in North America. Hydrogeology J., 9(1),
44-59.
McKay, L.D., J.A. Cherry, and R.W.
Gillham. 1993b. Field experiments in a
fractured clay till: 2. Solute and colloid
transport. Water Resour. Res., 29, 38793890.
Simpkins, W.W. and K.R. Bradbury. 1992.
Groundwater flow, velocity, and age in a
thick, fine-grained till unit in southeastern
Wisconsin. J. of Hydrol., 132, 283-319.
19
STOP 2 – ADA HAYDEN HERITAGE PARK
BURIED CHANNELS, QUARRY LAKES, BEAVER DAMS,
AND THE WATER SUPPLY OF AMES, IOWA
William W. Simpkins and Evan G. Christianson
Iowa State University
Figure 1. View looking north towards Ames, Iowa, showing alluvial valleys
(and unconfined aquifers) of Squaw Creek (left) and the South Skunk River
(right). Three-dimensional, shaded representation courtesy of David James of
the National Soil Tilth Laboratory.
of Ames during the past 40 years (e.g., Kent,
1969; Dougal et al., 1971; Austin et al.,
1984; Wille, 1984). However, increased
water demand, indications of contamination,
and the arrival of more sophisticated
modeling techniques and visualization tools
all suggest that a new study of the aquifer is
needed.
INTRODUCTION
The City of Ames, Iowa, located at
the confluence of Squaw Creek and the
South Skunk River (Fig. 1), receives its
drinking water from an alluvial aquifer. The
Ames aquifer, as it is collectively known, has
provided the city with excellent quality,
award-winning, water since the first
municipal well was drilled in 1891.
Why a new study? First, demand for
water has increased. Since the last major
study of the aquifer by Maroney (1994),
Ames and surrounding communities have
The Ames aquifer has been studied
extensively through cooperative efforts
between Iowa State University and the City
20
grown. For example, net-to-service pumpage
in Ames on July 11, 2005, reached an all
time high of nearly 9.5 million gallons.
Industries, including ethanol plants, have
come into the region partly because of a
plentiful supply of good quality water. A 50
Mgal/yr ethanol plant, Lincoln Way Energy,
LLC, requiring up to 500 gpm for regular
operation, is being installed nearby in
Nevada. It will draw water from same the
alluvial aquifer that supplies Ames. Future
supplies from the alluvial aquifer are also
uncertain, because climate change will affect
how much water will be available.
Second, there are indications that the
Ames aquifer is being contaminated.
Concentrations of Cl in two wells sampled in
1998 in the Downtown and Southeast well
fields were (Fig. 2) 26 and 33 mg/L,
respectively – typical values for groundwater
contaminated by road salt or fertilizer. In
addition, the Cl/Br ratios were 745.4 and
729.9, respectively, well above the normal
value of 50 expected from an atmospheric
source of Cl for the region (Davis et al.,
2004).
Contaminants may also enter
groundwater via induced infiltration from
streams. Recently, part of the South Skunk
River in direct hydraulic connection with the
Southeast and Downtown well fields was
identified as “impaired” on the Draft 2004
Clean Water Act Section 303(d) list
compiled by the Iowa Department of Natural
Resources. The listing was the result of
indicator bacteria above 400 cfu/100 mL in >
10% of the samples.
Finally, the Hallett Materials quarry –
the emergency water supply for the City of
Ames – was converted into Ada Hayden
Heritage Park in 2004. Although the City of
Ames exerts more control on the lakes
(including the stage), there is now increased
usage of the lake for fishing and other waterborne activities. The relationship among the
21
lake, the South Skunk River, and the Ames
aquifer is also not well understood.
Understanding the controls on water volume
and quality in the lakes at the park is critical
for their use as an emergency water supply.
This article highlights some of the
extensive previous work on the Ames aquifer
and presents some preliminary results of a
new three-year study of the Ames aquifer and
funded by the City of Ames. The primary
objective of the new study is to understand
the aquifer and the emergency water supply
within a holistic framework of water supply
and water quality. We will take advantage of
existing wealth of borehole, geophysical,
stratigraphic data and, will also apply new
tools, such as Geographic Information
Systems
(GIS),
and
3-dimensional
groundwater flow modeling, optimization,
and visualization techniques that are now
available. The results will provide the city
with the tools to manage the aquifer well into
the future.
DRINKING WATER IN AMES
The site of Ames, Iowa, was chosen
by the promoters of the Chicago and
Northwestern Railway in 1864 on high
ground between two streams, the South
Skunk River and Squaw Creek, which
migrated through wetlands (Seidel, 1991).
This proved to be fortuitous, because the
streams occupied buried bedrock channels
filled with outwash sand and gravel. After
drilling the first municipal well, additional
wells were drilled in 1906 at the site of the
present water treatment plant. Iowa State
University (ISU) also drilled its first well in
1897 and maintained its own water supply.
The city system expanded westward in 1924
to serve businesses and residences in that
area. Although the drought of 1977-78
moved the city and ISU to share water, ISU
continued to pump some of its own wells for
Figure 2. Location map of the Ames area showing major streams,
well fields (ISU, DT, SE, and SC), unconfined alluvial aquifer (gray),
and the low-head dam at River Valley Park. Dashed lines show
trends of major buried channel aquifers. Flags indicate golf courses.
Mining symbol is Whatoff's Pit.
drinking water until the mid-1990s, after
which it switched totally to city water.
Presently, ISU uses four wells in the vicinity
of the power plant for processed water use.
The city pumps water from two wells south
of the Lied Recreation Facility at the east end
of College Creek.
Ames currently has four well fields –
Downtown, ISU, Southeast, and the Sports
Complex (Fig. 2). The Downtown well field,
which is the oldest, is composed of 10 wells
in the confined portion of the aquifer, known
locally as the buried Skunk channel. The
earliest of these was drilled in the 1940s and
the last one was drilled in 1974. Wells range
from 33.5 to 40 m (110 to 130 ft) deep and
are often finished in the top of the
Mississippian limestone. Most of the water is
22
drawn from sand and gravel units
in the lower 12 to 21 m (40 to 70
ft) of the wells. These are large
diameter wells (gravel packs up to
5 ft or 1.5 m) with screen lengths
up to 9 m (30 ft) long. Well yields
range from 1908 to 4633 m3/d (350
to 850 gpm).
Pumping tests
performed as part of the Field
Methods in Hydrogeology course
(Geology 410/510) during the past
15 years show the aquifer to have
transmissivities (T) of 1200 to
hydraulic
1400
m2/day,
conductivity (K) values of 1 x 10-3
m/s, and storativities (S) near 2 x
10-4. Because of Iowa Department
of Natural Resources (IDNR)
regulations and the well field’s
location in downtown Ames (i.e.,
potential hazards of sewer lines and
underground storage tanks), it is
unlikely that the Downtown field
will be allowed to expand.
Till of the Dows Formation
comprises the main aquitard
(confining unit) for the aquifer
(Fig. 3); hence, the outwash sand and gravel
in the underlying buried channel could be
associated with earlier Wisconsinan or PreIllinoian ice. Peoria Formation loess has
been encountered below the till, suggesting
that the outwash might be Pre-Illinoian. The
buried channel is thought to connect directly
with the modern alluvium in the South Skunk
River close to River Valley Park, providing a
convenient point for vertical infiltration and
induced infiltration (via pumping) into the
confined aquifer.
The ISU well field is located on the
east side of the ISU campus (Figures 2 and 3)
and is physically separated from the
Downtown well field by Squaw Creek. The
well field lies in the buried Squaw Creek
1000
south in the late 1980s, as a
result of investigations by
Squaw Creek
Skunk River
Brookside
Cemetery
Austin et al. (1984).
The
Lincoln
River
Brookside Soccer well
well
Sand Way
Sand
Valley
Woods
Southeast well field (Fig. 2) is
and East
Park
well
and
wells
gravel well
gravel
Wisconsinan
situated just north of the
900
till
confluence of the South Skunk
River (flowing from the north)
Silt (loess)
and Squaw Creek (flowing
Ames Aquifer
(buried channel)
from the west). Because the
800
land was in private hands, there
was much controversy involved
Lower till unit (Pre-Illinoian)
in siting these wells.
The
Mississippian limestone
landowners,
who
hired
attorneys
all
the
way
from
New
700
York City to defend their
Figure 3. Schematic cross section of parts of the Ames aquifer
claim, refused to relinquish the
in the downtown and ISU areas based on Akhavi (1970), Nicklin
requested small parcel of land
(1974), and others.
for the wells; hence, the City of
Ames was forced to “take” the
channel, which is thought to be the older and
property under Eminent Domain. The first
deeper of the two bedrock channels.
well in this well field was drilled in 1986.
Reconnaissance borings in the in the 1970s
Downtown Well Field
Elevation (ft asl)
ISU Well Field
and 80s suggest that the aquifer is confined
in some areas and unconfined in others. The
Peoria Formation loess is present below the
alluvial sand and gravel associated with
Squaw Creek and, where present, is likely
the primary aquitard (confining unit) for the
sand and gravel in the buried channel (Fig.
3). Two city wells were drilled in 1978 and
1982 to depths of 43 m (140 ft). Although
production has been as high as 4088 to 4906
m3/d (750 to 900 gpm) the yield has been
quite variable. Values of T, K, and S are
about the same as those from the Downtown
well field based on previous studies. Just to
the north, four wells drilled between 1949
and 1969 are pumped for the ISU power
plant and have depths from 42 to 51 m (139
to 168 ft) – the deeper well may tap the
Mississipian limestone aquifer. Initial well
yields ranged from 5451 to 7359 m3/day
(1000 to 1350 gpm) and about 254 million
gallons of water are pumped annually (20002004 average).
The Ames aquifer expanded to the
23
Although the Southeast well field sits
within the flood plain of the modern South
Skunk River, the size of the valley and the
depth to bedrock (29 to 33 m; 95 to 100 ft)
suggest that it was carved by a much larger
stream graded to a lower based level. The
aquifer is unconfined and overlain only by
about 3.3 m (10 ft) of alluvial silt, which
grades into coarse sand and then very coarse
gravels with depth – a classic fining-upward
sequence. The aquifer could have formed
from outwash shed from the Altamont ice
margin just to the north. Five large-diameter
wells with gravel-pack diameters of 1.5 m (5
ft) are installed at the bottom of the aquifer.
Although the screens are on 12 to 15 m (40
to 50 ft) long, gravel packs are generally 26.5
to 30 m (87 to 99 ft) long. Yields range from
4633 to 8176 m3/day (850 to 1500 gpm),
although yields of 11,992 m3/day (2200 gpm)
are possible.
Pumping tests conducted
during the Field Methods in Hydrogeology
courses during the past 15 years suggest T, K
and Sy values of about 15,000 to 25,000
m2/day, 8 x 10-3 m/s and 0.2, respectively, for
this unconfined aquifer.
In 1999, two production wells were
installed in the Hunziker Youth Sports
Complex, just west of the South Skunk River
and south of U.S. Highway 30 (Fig. 2). The
hydrogeology at this site is similar to that in
the Southeast well field and the aquifer is
unconfined.
The main water-producing
zones occur in coarse gravel and boulders at
depths of about 15 to 30 m (50 to 97 ft). The
wells are 1.5 m (5 ft) in diameter with
stainless steel screens that cover the entire
production zone. An on-site well is used for
irrigating the soccer and baseball fields.
Seidel (1991) reported that the
percentage of water coming from each well
field was about 60 percent from the
Southeast well field and 30 percent from the
Downtown well field. Based on a total
pumpage of 2230 Mgal in 2004, the
percentages are similar – 35, 9, 37, and 19
percent for the Downtown, ISU, Southeast,
and Sports Complex, respectively.
THE DROUGHT OF 1976-77 AND ITS
HYDROLOGIC SOLUTION
Previous
investigators
have
documented the hydraulic connection
between streams and the aquifer – both in the
confined and unconfined portions (Dougal et
al., 1971). Under pre-pumping conditions, it
is likely that both the South Skunk River and
Squaw Creek received recharge from upland
areas (i.e., gaining streams) and down-valley
groundwater flow parallel to the stream. The
expansion of the Downtown well field,
however, induced groundwater flow towards
the pumping wells. Induced infiltration has
supplemented the depletion of aquifer storage
during wet periods when the streams contain
water; however, the relationship is strained
during drought periods, particularly in the
24
Downtown well field. The South Skunk
River now contains losing reaches and may
be a disconnected stream (see Winter et al.,
1999) at certain times of the year.
The hydraulic connection between the
modern South Skunk River alluvium and the
buried channel aquifer of the Downtown well
field was used to the City’s advantage during
a major drought. During July 1976 to June
1977, total precipitation was only 33 cm (13
in) – nearly 20 inches less than normal
(Seidel, 1991). Stream flow declined in the
fall of 1976 and both the Skunk River and
Squaw Creek remained dry into the summer
of 1977. The hot, dry weather prompted
residents to use more water, which lowered
the potentiometric surface in the Downtown
well field (the only one) to dangerously low
levels
despite
stringent
conservation
measures. Parts of the aquifer probably
converted to unconfined conditions at this
time. Dr. Merwin Dougal and his colleagues
at the Iowa State Water Resources Research
Institute developed an ingenious plan to raise
the water levels in the well field. They
proposed building a temporary dam across
the South Skunk River at River Valley Park
on 13th Street and pumping water from the
Hallett Materials quarry (Ada Hayden
Heritage Park) across Highway 69 and into
the river. The dam was constructed by
bulldozing sand and gravel from the river bed
into a structure about 2.4 to 3.3 m (8 to 10 ft)
high, covering the upstream face with plastic
liner material, and then adding another 0.3 m
(1 ft) of sand on top to minimize leakage
through the dam. The dam was located at
this spot because it was believed that water
would infiltrate through the bed of the
channel and migrate southwestward toward
the Downtown well field along the axis of
the buried channel (Figs. 2 and 3).
Under an agreement with Hallett
Materials and a user fee of $1, pumping
Figure 4. Photograph of a surface pump at the Hallett
Materials quarry on July 12, 1977. Courtesy of the City of
Ames Water and Pollution Control Department.
commenced on July 12, 1977 from the shores
of the quarry (Fig. 4). Four surface pumps,
with a total estimated yield (4-day average)
of about 43,608 m3/day (8000 gpm),
discharged water to the nearby South Skunk
River at a point nearly 4 km (2.5 mi)
upstream of the low-head dam in River
Valley Park. Despite the best efforts of
beavers to dam the river upstream and hold
the water above the sand dam, water pumped
into the channel stretched about 1.6 km (1
mi) upstream from the dam after 4 days.
Only about 25 percent of the pumped water
was unaccounted for and presumably had
leaked into alluvium prior to the dam (Seidel,
1991). As a result, normal water levels in the
Downtown well field were restored within a
few days (Seidel, 1991). Because of the
distance of about 1.6 km (1 mi) between the
dam and the closest well in the Downtown
well field, it is likely that the rapid rise of
water levels to extra water in the South
Skunk River reflects a pressure response
induced at the boundary of a confined aquifer
(the “backwater effect” discussed by Todd,
1980), rather than physical migration of
water from the stream bed to the Downtown
well field.
Assuming the river was a
disconnected stream, it would only have been
necessary to wet the bottom of the channel
and convert it to a losing stream. At that
25
point, the pressure response would have been
felt almost instantaneously in the Downtown
well field. The sand dam washed away later
that summer and was rebuilt again in 19811982 during a drought period. A permanent
low-head concrete dam was built at the site
during the winter of 1983-84. Temporary
risers allow water to pool up to 1.5 m (5 ft)
behind the dam (Seidel, 1991). Pumpage of
water into the South Skunk River occurred
during drought periods of 1981-82 and 1988.
Peterson Pits, about 4 km (2.4 mi) upstream,
were used in 2000 when the Hallett Materials
quarry was not available.
ADA HAYDEN HERITAGE PARK:
DEVELOPMENT, RECREATION, AND
AN EMERGENCY WATER SUPPLY
The lakes at Ada Hayden Heritage
Park (Fig. 2) were formed by gravel
excavation (i.e., a dredging operation) that
began with the E.I. Sargent Company in the
mid-1930s and continued under the auspices
of the Hallett Materials Co. until 1996.
Coarse sand and gravel, which appeared to
be in plentiful supply until about the mid1990s, was the main product of the
operation.
Because of the quarry’s ties to the
water supply history of Ames, there was
cause for concern when mining ceased in
1996. It was clear that land might be sold to
parties not interested in an emergency water
supply for the city. The following timeline
of events was provided by the City of Ames:
•
Summer 1998 – Des Moines developer
approaches Hallett Materials Co. for
purchase option and agrees to buy land.
Ames’ access to the quarry for water
supply is denied;
•
February 1999 – Des Moines developer
approaches City of Ames about land
annexation and development of a 1400unit “Grand Lakes” subdivision;
•
March to December 1999 - City examines
proposal and hires lake consultant to
report by April 2000;
•
May 2000 – Consultant reports that
water quality in the lakes will likely
deteriorate over time, particularly with
1400 homes near the shoreline;
•
June 2000 - Dry conditions cause water
levels to drop in the Downtown well field,
but access to lake water is not possible;
•
Late Summer 2000 – City rejects
developer proposal and decides to
investigate purchasing the land for a
park;
•
•
September-October 2000 – About 6 x 105
m3 of water is pumped into the South
Skunk River from the lake at Peterson
Pits; the lake was pumped dry after 491.5
hours (~ 20 days);
November 2001 – Ames residents
approve (86%) a bond issue to purchase
the lake area.
The lakes were purchased by the city
and converted into Ada Hayden Heritage
Park, which was officially dedicated on
August 28, 2004.
Its namesake, Ada
Hayden, grew up just south of the park, was
an Ames High School graduate, and was the
first woman to receive a Ph.D. from Iowa
State College (now ISU). She taught Botany,
was Curator of the ISU Herbarium, worked
in the Experiment Station at ISU, and was
inducted into the Iowa Conservation Hall of
Fame in 1967 among other honors. Funds
for the purchase and $6.47 million
renovation of the area were obtained from a
$1.5 million Vision Iowa CAT grant and a
26
$4.97 million bond referendum. An
additional $1 million “in-kind” service grant
from the Story County Conservation Board
and SWC District will be used to plant
prairie grasses at the park.
Purchase of the land for a park
guaranteed that the city would maintain a
backup water supply of good quality water.
The water supply potential of the north and
south lakes is huge (Table 1). They are up to
18.6 m (61 ft) deep and the total surface area
of the two lakes is 50 hectares (123 acres).
They are fed by a watershed of 1150 hectares
(2840 acres) containing agricultural and
urban uses and moderately sloping to steeply
sloping topography.
Table 1. Characteristics of the North and
South Lakes at Ada Hayden Heritage Park
(from Downing et al., 2005).
North Lake:
Lake surface area: 16.0 ha (40 ac)
Maximum depth: 14.6 m (48 ft)
Mean depth: 7.3 m (24 ft)
Volume: 1.9 million m3 (0.34 billion gal)
Length of shoreline: 1.9 km (1.2 mi)
Watershed area: 96 ha (236 acres)
Watershed area/Lake area ratio: 6
South Lake:
Lake surface area: 33.6 ha (83 ac)
Maximum depth: 18.6 meters (61 ft)
Mean depth: 10.4 meters (34 ft)
Volume: 3.3 million m3 (0.86 billion gal)
Length of shoreline: 3.2 km (2.0 mi)
Watershed area: 1054 ha (2604 acres)
Watershed area/Lake area ratio: 31
Research by Downing et al. (2005)
suggests that both north and south lakes are
dimictic lakes; i.e., they mix from top to
bottom only in the spring and fall.
Phosphorus is the principal limiting nutrient,
owing to the high ratio of nitrogen (N) to
phosphorus (P) in both lakes. The average
total phosphorus (total P) concentrations at
urban-rural lands (Downing et al. 2005).
Monitoring indicates that all three tributaries
exhibit high P concentrations, with total P
concentrations > 100 µg/L. Nitrate-N
concentrations > 10 mg/L occur mainly in
tributaries A and D; tributary B shows much
lower concentrations. Samples of inflow and
outflow water from the cells since their
construction in 2003 suggest that they
remove some, but not all, P and N prior to
entering the lakes.
Hopefully, the
effectiveness of these treatment wetlands will
increase with time, although agricultural Best
Management Practices and a reduction in Pbased fertilizer by the urban residents could
help to reduce nutrient and sediment inputs
to the lake.
Figure 5. Aerial photo (c. 2004) showing the Ada
Hayden Heritage Park area (north and south Lakes),
Skunk River, and treatment wetlands. The watershed
contains row crop agriculture to the north and a large
subdivision to the south.
shallow depths are 20 and 25 ppb for the
north and south lakes, respectively. Owing to
the depth of the lakes and the moderate
degree of development in the watershed,
these P concentrations are lower than most of
the 132 lakes surveyed routinely in the Iowa
Lakes Survey (Downing et al., 2005).
Nitrogen concentrations are about five times
greater in the south lake than in the north
lake, which may suggest a greater
groundwater flow component in the south
lake.
A consulting report and subsequent
Master Plan compiled by the city advocated
the creation of treatment wetlands to
intercept surface water from the three main
tributaries to the lake. The purpose of the
wetlands is to intercept nutrients (N and P)
and suspended sediment and process them
prior to entering the lake. The exit points of
these tributaries into the lake are shown as
sites A, B, and D in Figures 5 and 6.
Tributary A drains agricultural, tributary B
drains urban, and tributary D drains mixed
27
Although studies of the limnology
and surface hydrology have occurred at the
park since 2001, no groundwater work has
been done.
However, preliminary mass
balance calculations suggest that the
concentrations of nutrients in the lakes
cannot be accounted for by surface water
inflow alone and that groundwater must be
involved in transporting nutrients (John
Downing, verbal communication, 2004).
Thus, we hypothesized initially that
groundwater may discharge into the lake on
the north and west sides, leave the lake on
Figure 6. View to the southeast overlooking two large
treatment wetlands related to sites D and B in Figure 5.
The parking lot and trail used during Stop 2 are also
shown.
the east side, and flow down the Skunk River
floodplain – a classic flow-through lake (see
Simpkins, 2006). As part of the new study of
the Ames aquifer, we will install nested
piezometers (wells) at 7 sites around the
lakes in 2006 to test this hypothesis.
In order to guide the piezometer
location and provide an educational
opportunity for the Hydrogeology and
Watershed Hydrology classes at ISU, we
sampled groundwater at three locations on
the lake margin on August 25, 2005, using a
Geoprobe, Iowa DNR staff, and a Screen
Point groundwater sampler. We pushed the
rod until refusal (usually about 18 m – close
to or below the bottom of the lakes), exposed
the well screen, measured hydraulic head,
and sampled groundwater about every 1.2 to
2.4 m as we pulled the probe rod to the
surface. Samples for Cl, SO4, and NO3-N
were analyzed by ion chromatography at the
National Soil Tilth Laboratory. Ortho-P of
the supernatant fraction of unfiltered samples
was analyzed using the Murphy and
Riley
(1962)
method
(after
centrifugation at 4500 rpm for 5
minutes in the Soil Geochemistry
Laboratory, Agronomy Department,
ISU).
Absolute elevations were
determined using a GPS unit on
September 6, 2005.
Figure 7. Profiles depicting hydraulic head, temperature, and
water quality with depth on August 25, 2005 at sites AH-1 and
AH-3 (see Figure 5 for location). The elevation of the lake
surface was 273.4 m asl.
28
The results supported the
conceptual model, but contained a
few surprises. Sediments consisted of
unoxidized, gray, fine to mediumgrained sand with some silt. The
direction of the vertical hydraulic
gradient at AH-1 along with decrease
in
temperature
and
anion
concentrations with depth suggest
groundwater is being recharged here –
right next to the lake (Fig. 5). The
water table elevation of 273.5 m asl is
slightly above the lake level at 273.4;
hence flow is toward the lake at the
water table. Recharge might occur
here because of the wetland
upgradient from this spot. At depth,
the hydraulic heads are slightly lower
than those in the lake, suggesting that
outflow may be occurring through the
lake bottom. The smell of H2S in all
the samples and lack of sulfate and
nitrate-N in the profile at depth
suggest that reducing conditions
prevail and that sulfate reduction is
discussed earlier.
Figure 8
summarizes the new conceptual
model for the site, subject to rigorous
testing by drilling and well
installation in 2006.
SUMMARY
The City of Ames relies
solely on groundwater derived from
confined (buried channel) aquifers
and unconfined alluvial aquifers for
its drinking water. Its backup water
supply consists of a former gravel quarry
transformed into a recreational lake.
Groundwater inflow probably replenishes the
nearly 5.2 million m3 (1 billion gallons) of
water in the lake and lake outflow feeds
groundwater that flows along the axis of the
South Skunk River. Droughts in 1976, 19811982, and 1988 (and almost in 2000) caused
the City of Ames to pump water from the
lake, across Highway 69, and into the river to
pond behind a dam. This action raised water
levels in the Downtown well field within
days – probably a result of a pressure
response or backwater effect.
Figure 8. Revised conceptual model of groundwater flow in the
vicinity of the North Lake at Ada Hayden Heritage Park depicting
the lake as a flow-through system.
occurring.
Although we are unable to
explain the pattern of ortho-P concentration
with depth, the concentrations are similar to
those found in surface Tributary A and
generally higher than in the lake. The
concentration of 1922 µg/L at 11.6 m may
overlie a change in lithology. Data from AH2 (data not shown) are similar to AH-1 in all
respects except that the hydraulic head
increases slightly with depth, suggesting
groundwater discharge at that location.
Coincidentally, wetlands are not present
upgradient of AH-2 (Fig. 5).
Site AH-3 on the eastern side of the
lake (Fig. 5) provides evidence that this is a
flow-through lake system. Hydraulic heads
are below lake level, indicating that lake
water can flow into groundwater. Hydraulic
heads are nearly equal with depth (suggesting
horizontal flow) and are also above the
elevation of the Skunk River at 272.2 m asl
(see Stage location on Fig. 5). In addition,
concentrations of Cl, SO4, and ortho-P are
also fairly constant with depth and are
similar to those found in the lake. The lack
of NO3-N is not surprising given that
concentrations in the north lake are below the
detection limit (Downing et al., 2005).
Presumably, water from the lake provides
base flow for the South Skunk River, but the
bulk of the water may flow down valley
parallel to the stream – a model that is
consistent with the disconnected stream idea
29
Research as part of the new study of
the Ames aquifer will help provide a
plentiful supply of good quality water for
Ames into the 21st century. It will include a
groundwater flow investigation at Ada
Hayden Heritage Park, a re-analysis of
stratigraphic,
hydrogeologic,
and
geochemical data of the aquifer, and
production of a state-of-the art, 3-D model of
groundwater flow, including delineation of
well capture zones, optimization of pumping
schemes and a siting scheme for new wells in
the aquifer.
ACKNOWLEDGEMENTS
For this article, we are grateful for the
help provided by the staff of the City of
Ames Department of Water and Pollution
Control, particularly Tom Neumann and Dr.
Harris Seidel, the present and former
directors, respectively, of that department, as
well as John Dunn, Lyle Hammes, Phil
Propes, Karla Tebben, and Barbara Schendel.
We also thank Bob Drustrup and Hylton
Jackson of IDNR, Kevin Cole and Colin
Greenan of NSTL and Dr. Michael
Thompson of the Agronomy Department for
their assistance in the laboratory and the
field.
REFERENCES
Akhavi, M.S. 1970. Occurrence, movement,
and evaluation of shallow groundwater in
the Ames, Iowa area. Unpublished M.S.
thesis. Iowa State University, Ames,
Iowa.
Austin, T.A., R. Drustrup, L. Antosch,
L.Wille, and W.W. Parsons. 1984.
Supplemental water supply studies, City
of Ames completion report. Unpubl.
Report to the Iowa State Water Resources
Research Institute, Iowa State University,
45 p. with appendices.
Davis, S.N., J.T. Fabryka-Martin, and L.E.
Wolfsberg, 2004. Variations of bromide
in potable ground water in the United
States. Ground Water, 42(6), 902-909.
Dougal, M.D., L.V.A. Sendlein, R.L.
Johnson, and M.S. Akhavi. 1971.
Groundwater
and
surface
water
relationships for the Skunk River at
Ames, Iowa. Special Report, Engineering
Research
Institute
ISU-ERI-AMES
99984, Project 893-S, 157 p. with
appendices.
Downing, J. A., G. Antoniou, D. Kendall, D.
Stipp-Bethune, and J. Li. 2005. Ada
30
Hayden Heritage Park Lakes Monitoring
– Interim Report. January 2005. 36 p.
Kent,
D.B.
1969.
A
preliminary
hydrogeologic investigation of the upper
Skunk River basin. Unpubl. Ph.D.
dissertation, Iowa State University,
Ames. 375 p.
Maroney, C.L.R. 1994. Evaluation of the
future water supply alternatives for the
city of Ames, Iowa. Unpubl. M.S. thesis,
Department of Civil and Construction
Engineering, Iowa State University, 190
p.
Murphy, J, and J.P. Riley. 1962. A modified
single solution method for the
determination of phosphate in natural
waters. Anal. Chem. Acta., 27, 31-36.
Nicklin, M.E., 1974. The hydrogeology of
the regolith aquifer supplying the Iowa
State University well field. Unpubl. M.S.
thesis, Iowa State University, Ames, IA.
131 p.
Seidel, H. 1991. Groundwater supply of
Ames,
Iowa.
Iowa
Groundwater
Quarterly, 2(9), 20-21.
Simpkins, W.W. 2006. A multi-scale
investigation of groundwater flow at
Clear Lake, Iowa. Ground Water 44(1).
Todd, D.K. 1980. Ground water hydrology.
New York, Wiley and Sons.
Wille, L.E. 1984. The hydrogeologic
investigation of the southeast well field
and McCallsburg Arm, Ames, Iowa.
Unpublished M.S. thesis. Iowa State
University, Ames, Iowa.
Winter, T.C., J.W. Harvey, O.L. Franke, and
W.M. Alley. 2002. Ground water and
surface water: a single resource. U.S.
Geological Survey Circular 1139. 77 p.
Yazicigil, H. 1977. Mathematical modeling
and management of groundwater
31
contaminated by aromatic hydrocarbons
in the aquifer supplying Ames, IA
Unpublished M.S. thesis, Iowa State
University, Ames, IA.
STOP 3 – DOLLIVER PARK, DES MOINES CHEROKEE GROUP
Carl F. Vondra
Iowa State University
(The following is reprinted from Lemish, J., Chamberlain, R.E., and Mason, E.W., 1981,
Part 1: Introduction and Regional geology, Iowa Geological Survey Guidebook, Series
No. 5, pp. 2-22)
GEOLOGIC SETTING
The area of this study is located on
the northwest flank of the Forest City Basin.
Major structural elements of regional extent
(Fig. 2) include the Wisconsin Dome to the
north, the Mississippi Arch to the east, the
Ozark Dome and Bourbon Arch to the south,
syncline plunging southward toward the
depositional center of the Basin in northern
Missouri.
The Paleozoic and younger rocks in
the Basin reach a thickness of over 5,200 feet
in southwestern Iowa (Fig. 3). The Paleozoic
section includes Cambrian sands overlain by
Ordovician, Silurian, Devonian, and
Mississippian sediments, predominantly
carbonates. Their combined thickness is
about 3,400 feet. The Pennsylvanian
System reaches 1,700 feet in thickness
in the southwestern part of the state and
covers 20,000 square miles of Iowa.
These rocks are unconformably overlain
by Cretaceous sandstone, shale, and
limestone which have a total thickness
of 500 feet but locally seldom exceed
100 feet. A cover of Pleistocene drift
from 0 to 500 feet thick mantles much
of the state.
Six major unconformities exist
within the Paleozoic section: at the
bases of the St. Peter Sandstone,
Maquoketa Formation and Maple Mill
Shale, and at the tops of the Gilmore
City Limestone (Mississippian), the
Mississippian
System,
and
the
Desmoinesian
Series.
Minor
unconformities were developed at
Figure 1. Index map showing field trip area in Webster
various intervals throughout the
County. The geologic map showing distribution of
Paleozoic section. The unconformity of
Pennsylvanian rocks is presented.
greatest
importance
influencing
subsequent
deposition
of
the
and the Nemaha Ridge to the west. In Iowa,
Pennsylvanian System is the widespread
the Forest City Basin forms a shallow
erosion surface with over 200 feet of relief
32
series of northwest trending structures
along the southern flank, called the
Thurman-Wilson
Structural
Zone,
includes faults and anticlines (i.e. -- the
Redfield and Ames anticlines). Over 400
feet of structural relief in Paleozoic rocks
have been related to this positive element
in the basement (Figs. 2 and 4). The
study area of the field trip is situated on
the northern margin of this feature.
DESMOINESIAN SERIES
STRATIGRAPHY
Figure 2. Major structural features of the mid-continent
region.
which
developed
on
the
exposed
Mississippian rocks during the ChesterMorrowan interval.
The major subsurface feature in the
basin is the Midcontinent Geophysical
Anomaly (Fig. 2) trending northeast across
Iowa from Nebraska to the Duluth area of
Minnesota (Coons, et al., 1967; Van Eck, et
al., 1979). This feature consists of a series of
gravity and magnetic highs and has been
interpreted as a rift containing basalts and
associated sediments of Keweenawan age. It
presently is considered to be a fault-bounded
horst about 40 or more miles wide flanked by
a great thickness of clastic fill. This feature
has been tectonically active exerting
basement control on structural development
within the basin since the Precambrian. A
33
The Desmoinesian Series consists
of two groups, the predominantly clastic
Cherokee Group and the overlying
Marmaton (Fig. 5). The Marmaton and
Cherokee Groups are of major economic
importance because of their coal deposits.
With the exception of the Mystic seam in
the Marmaton and a minor amount of
production from the Nodaway seam in
the Wabaunsee Group (Virgilian Series),
the bulk of Iowa's coal resources base
occurs within the Cherokee.
Stratigraphic study by Wanless
(1975) and McKee and Crosby (1975)
indicates that Atokan Series sediments occur
in the subsurface beneath the Desmoinesian
Series. Up to 400 feet of Atokan sediments
may be present as sandstone and shale
similar to Lower Cherokee clastic rocks. The
Atokan is recognized in Illinois and
Missouri. It is difficult to separate these units
in the subsurface records in Iowa, and as a
result, Atokan sediments are tentatively
included as part of the Cherokee Group.
The overlying Missourian and
Virgilian Series (Fig. 5) are deposited on top
of the Desmoinesian as part of an overall
marine transgression and represents an onlap
sequence with numerous transgressive and
regressive phases (related to cyclothem
deposits) (Heckel, 1977). Marine deposits of
the Cherokee Group occur throughout eastern
Iowa and are evidence that Pennsylvanian
seas traversed the entire state. By the time of
deposition of the Upper Cherokee, the
Illinois and Forest City Basins were
connected. Thus, the present boundaries of
the Cherokee Group represent the erosional
remnants rather than actual aerial extent of
Cherokee Seas in the midcontinent (Dapples
and Hopkins, 1969).
The Cherokee Group is composed
primarily of deltaic sediments deposited on
an irregular Mississippian erosion surface of
considerable relief. As a result, lithologies
are extremely variable laterally and vertically
and consist of mudrocks with subordinate
sandstone and limestone and localized coal
seams. Factors contributing to basal
lithologic variability include inherited
Mississippian
paleotopography,
prePennsylvanian structural trends, contemporaneous structural movement, and
differential subsidence and
related sediment compaction.
These elements exerted a
major
control
on
the
distribution of the Cherokee
sand bodies.
The
overlying
Marmaton Group, in general,
exhibits the characteristics of
marine-dominated sedimentation. It has been divided into
four cyclothems: the St.
David, Bereton, Sparland, and
Gimlet (Wanless and Wright,
1978), and these were
subdivided
into
seven
formations which can be
generally recognized over
wide areas (Fig. 6). The
Marmaton Group is composed
of two basic litho-logic
associations:
a
marine
transgressive phase represented by their persistent
limestones and minor black
shales and a regressive phase
represented
by
thin
sandstones, shales, underclays,
and coals. The coals and
carbonaceous zones occur at
several stratigraphic horizons
representing distal deltaic
platforms built up during
Figure 3. Generalized stratigraphic column for Iowa.
34
regressive phases. The coals tend to be thin,
and only two have been mined, the Lonsdale
and the Mystic (Keys, 1894). The Mystic
averages 2 1/2 to 3 feet in thickness across
south-central Iowa and northern Missouri.
Correlative deltaic coals of the Illinois Basin
are much thicker; the Herrin (No. 6) coal
represents one of the most widely mined
coals in Illinois and is the equivalent of the
Mystic coal seam (Lexington in Missouri).
Characteristics of the Cherokee Group
In Iowa the Cherokee Group
(Haworth and Kirk, 1894) is defined as those
rocks between the overlying Ft. Scott
Formation of the Marmaton Group and the
erosional
surface
formed
on
the
Mississippian aged rocks. Cherokee Group
nomenclature varies according to location in
Iowa. The correlations are tenuous because
of lateral and vertical lithologic variability
prevalent in the lower Cherokee. Many of the
units of the upper portion of the Cherokee in
southeastern and south-central Iowa where
they were named may extend to north-central
Iowa. Nevertheless, the Cherokee has been
subdivided into an Upper and Lower unit
(Fig. 7) with an indistinct boundary; it is
generally agreed that the base of the Upper
Cherokee coincides with the Hanging
Rock Sandstone (Landis and Van Eck,
1965).
In subsurface studies of the
geology of the deep coal in the Iowa
portion of the Forest City Basin,
correlations with the Cherokee units,
shown in Figure 7, are difficult to
achieve. The one prominent unit
which serves as a marker bed is the
Ardmore
limestone,
readily
recognizable throughout the subsurface from Iowa to Oklahoma. Thus,
it is used to establish the position of
sandstone bodies in the subsurface.
Cherokee Lithologies
Figure 4. Structure section on NW-SE direction across the
Thurman-Wilson structural zone showing the structure and
corresponding gravity anomaly. Displacement in the lower
Paleozoic is greater than displacement in the upper Paleozoic
indicating continued tectonic activity along the zone.
35
Derynck (1980) analyzed the
"undifferentiated" Cherokee Group
(Cherokee Group and rocks of the
Atokan
Series
combined)
for
lithologic composition based on well
log information in southwestern Iowa.
His results indicate that the group is
composed of 80% shales, 17%
sandstones
and
siltstones,
3%
limestones, dolomites, coals, and
underclays; 14% of the shales are
black. Approximately 3.2 feet of coal
Limestones
are
light
gray,
fossiliferous, and generally less than one foot
thick. Frequently, limestones are represented
by isolated nodules embedded in a shale or
mudstone. Underclays are light gray, silty,
and/or sandy and generally about two to four
feet in thickness (Derynck, 1980).
Subsurface studies by Reese (1977)
and Mason (1980) in the Madrid area of
central Iowa provide evidence of the
distribution of Cherokee lithologies. Their
studies indicate that the Cherokee section
averages 350 feet in thickness with limestone
units occurring only in the uppermost 80-100
feet of the section. Thus, the Cherokee strata
appear to become more dominated by continental sedimentation to the northeast.
Figure 5. Detailed stratigaphy of the Cherokee Group and its
relationship to other Pennsylvanian rocks of Iowa.
Experience to date indicates that the
Lower Cherokee, especially in Central Iowa,
is characteristically more clastic than the
occur per 400 feet of section, less than 1% in
overall composition.
Shales are commonly gray, silty,
and/or sandy with interlaminations of gray
siltstone or fine-grained sandstone common
throughout. Some gray shales are calcareous
with fossils or calcareous ironstone
concretions. The black shales are laminated,
phosphatic in part, and commonly pyretic.
The sandstones are generally very
fine-grained to fine-grained and argillaceous.
Thin sand units typically grade into thicker
units, which are commonly massive to crossbedded with erosional basal contacts. Thick
units consist of vertically stacked sand bodies
typified by the sandstone exposed at Redrock
Reservoir, Marion County, Iowa, and likely
indicate possible structural or topographic
control on sedimentation (Brown, 1975).
Offsetting of sand units within the Cherokee
may be evidenceofcompactional influence as
well (Derynck, 1980; Mason, 1980).
36
Figure 6.Detailed stratigraphy of the Marmaton
Group.
Upper Cherokee. It is characterized by a
scarcity of key marker horizons and consists
of units of variable thickness and lithology,
including lenticular and local sandstones,
marine
limestones
(characteristically
represented by isolated concretionary masses
within shale units), and localized lenticular
coal beds. The one limestone which is
recognizable and persistent is the Laddsdale
or Seville Limestone. It is characteristically
dark gray, hard to earthy, fossiliferous, and
lenticular, grading locally to a fossiliferous
sandstone. It is known to exist in southeast
Iowa (Davis Co.), but its presence has not
been recognized in the Fort Dodge area.
The Upper Cherokee is far more
consistent in lithologic character and demonstrates cyclothemic characteristics. It consists
of several persistent limestone, sandstone,
and coal units (Fig. 7). In southeastern Iowa
the basal unit is the Hanging Rock Sandstone
(Landis and Van Eck, 1965) which directly
underlies the Munterville Coal (Fig. 7).
Recognizable marine and continental units
alternate throughout the section with the
Whitebreast Coal and Ardmore Limestone
representing the most widespread deposition
during Cherokee times. The Whitebreast and
its correlative units comprise one of the most
widespread coals in North America; during
its deposition, a vast vegetated plain
extended from eastern Kansas to western
Indiana (Wanless and Wright, 1978). The
White-breast Coal in Iowa is the equivalent
of the Colchester (No. 2) Coal of Illinois and
the Croweburg Coal of Missouri.
In summary, the Cherokee Group
consists of variable lithologies representing
numerous cyclothems. These grade upward
into more continuous units as the
topographic
irregularities
of
the
Mississippian unconformity became less
dominant. Typical lithologies include a
basal sandstone (with channel and sheet
phases) overlain by an underclay-coal
sequence. These lithologies are typically
restricted geographically to a one to three
county area (Landis and Van Eck, 1965).
Succeeding this sandstone-underclay-coal
sequence are black shales and limestones
representing minor transgressive sequences
over coal swamps and deltaic platforms.
Correlations over large areas are tenuous,
and detailed descriptions are valid only on
a local paleontologic and palynologic
scale. Studies initiated as part of the Iowa
Coal Program (conducted by the Iowa
Geological Survey) under the direction of
Dr. Matthew Avcin are essential to the
solution of this problem.
Figure 7. Detailed stratigraphy of the Cherokee Group and
component cyclothems.
37
into the coals
lithologies.
Coal
According to the Iowa Geological
Survey (Avcin, pers. com., 1978), up to 19
individual coal beds occur in Iowa.
Characteristically the lowermost coal seams
are thickest and, therefore, most often mined.
The upper coals tend to be thinner but more
widespread. Detailed studies at Madrid
(Reese, 1977) confirm this. The recognized
coal seams are presented in figure 7; the nomenclature shown is that used for those coals
in south-central and southeastern Iowa.
Field experience and other studies
confirm the economic importance of the
lower coals. Work by Chamberlain (1980),
Derynck (1980), and Robertson (1976)
indicates that the Cherokee represents an
overall onlap sequence and that basal coals
get younger to the east as the basin was
filled.
In the Madrid area (Reese, 1977;
Mason, 1980), a total of 10 coals were recognized; the two mineable seams occur in the
lower part of the section 80 to 125 feet above
the Mississippian erosion surface. Again, the
upper coals tend to be thinner and more
widespread than the lower coals.
Two coal beds in the Lower Cherokee
were mined in the Fort Dodge area of Webster County. According to Landis and Van
Eck (1965), the local names were the Pretty
(also known as the Big or Upper
Bitumunous) coal and a bed known as the
Big Seam (also Calhoun-Camel Bed) about
35 feet below the Pretty seam. In all, up to 4
seams appear to be present. Stratigraphically
these coals are believed to occur in the
Lower Cherokee because of the proximity
(100' or less) of the Mississippian erosion
surface. The sand bodies occur lateral to the
coals and occasionally can be shown to cut
38
and
other
associated
Sandstone
During studies of the deep coal in the
Forest City Basin, well log records indicated
the presence of sand bodies having
thicknesses of greater than 100 feet. The
positions of the sand bodies appear to be
controlled by a well-defined drainage system
developed in the uppermost Mississippian
deposits. The presence of several thick
sandstone bodies exposed along the Des
Moines River led to a field study of their
occurrence in an attempt to better understand
their relationship to paleodrainage and the
incidence of coal.
Exposures
Sandstone
bodies
outcrop
discontinuously along the Des Moines River
and its tributaries throughout much of central
Iowa. In particular many significant
exposures occur between Fort Dodge, in
Webster County, and Red Rock Reservoir in
Marion County. Study of these and other
sandstone outcrops in central Iowa forms the
basis for the interpretations discussed in this
guidebook.
In general, primary sedimentary
structures identified and measured on the
outcrop indicate a strong southwesterly
paleocurrent direction. Sand bodies are
marked by erosional basal contacts; coarsegrained basal intervals characteristically rich
in fragmented clasts derived from underlying
and/or lateral rock bodies; cosets of largescale planar cross-strata; abundant carbonaceous debris, and a general finingupward grain size. On a regional scale, it is
believed that the sand bodies studied in
central Iowa indicate the occurrence during
deposition of the Cherokee strata of a series
of predominantly southwestardly flowing
rivers. Farther south, in Marion County,
studies by Hansen (1978) suggest a transition
from a convergent to a divergent channel
pattern which is interpreted to indicate a
deltaic distributary system developed
basinward from an area of dominant fluvial
deposition. A general paleogeographic map
representing conditions prevalent during the
deposition of Lower Cherokee sediments
includes a pronounced southwestward
drainage network throughout central Iowa
(Fig. 8). This reconstruction is in marked
occurrence in the deeper part of the basin. As
a result, studies were undertaken by Derynck
(1980), Chamberlain (1980), and Mason
(1980) in a series of Iowa State University
master's theses. Derynck (1980) studied the
distribution of the sands and their position in
the Cherokee; Chamberlain related the
distribution and stratigraphic interval to the
Mississippian erosion surface; and Mason
(1980) made an in-depth study of the
sandstone, shale, and sandy shale occurring
between coal beds in a 23-square-mile area
near Madrid.
A critical aspect of the
subsurface sandstone distribution is its
relationship to the Mississippian
erosion surface. Chamberlain (1980)
studied
the
unconformity
and
constructed a paleotopographic map
(Fig. 9A) of this critical surface to see
what effect it had on the distribution of
sandstone bodies and coal. The map is
based partly on the work of others in
the Forest City and Illinois Basins
(Bretz, 1950; Lee, 1943; Lee and
Payne, 1944; Bransen, 1962; Siever,
1951; Wanless, 1975; Wanless and
Figure 8. Inferred Cherokee pleodrainage based on outcrop
Wright, 1978; Potter, 1979). The map
and subsurface evidence of sand bodies. The sand bodies are
at different stratigraphic levels within the Cherokee and are
shows the paleotopographic patterns
considered to be of different ages. (1) Nemaha ridge source,
developed in the Mississippian erosion
(2) source from the north, (3) Fort Dodge-Holiday Creek area,
surface during initial Pennsylvanian
(4) Dolliver Park area, (5) Ledges area, (6) Altoona (subsurface)
sedimentation. These patterns exerted a
area, (7) Red Rock area, (8) Southeast Iowa-northeast Missouri
progressively decreasing control on
area.
sedimentation until deposition of the
Ardmore Limestone during Upper
contrast to the ideas of earlier workers
Cherokee
time (Wanless and Wright, 1978).
(McKee and Crosby, 1975) which suggested
At this point, paleotopographic irregularities
a single dominantly southwardly trending
were covered, and the Lower Cherokee
river during this interval of geologic time.
deposits formed a broad platform on which
the Whitebreast and later coals were
Subsurface Sand Bodies
deposited.
The surface distribution and the
predominant southwesterly drainage directions indicated by surface sand bodies
focused attention on their subsurface
39
Paleotopography
township to none in Cass County.
The Mississippian plaeotopographic
map (Fig. 9A) was constructed using the
unconformity contour method (Anderson,
1962) which requires that the base of the
valley fill must be recognized in subsurface
records. Following this method, the altitude
relating to the top of Mississippian strata was
recorded on a map, and a structure contour
map of the Mississippian surface was made
with present day sea level as the reference
plane.
Later structural movements can
distort the valley fill relations.
Other
methods to determine paleo-drainage were
not used because of the extreme variation in
density of bore hole data used for control,
which ranges from several points per
Contours were drawn assuming that
drainage initially developed on a relatively
level surface developed on a uniform
lithology, predominantly of Mississippian
limestones. This results in a radial or
modified dendritic drainage pattern trending
to the south-southwest from the eastern
margin of the basin toward the center and to
the east and south under the influence of the
Nemaha Ridge along the west flank of the
basin. Laury (1968) used a similar approach
in his study of the Upper Cherokee
Pleasantview Sandstone in central and southcentral Iowa.
Figure 9. A. Paleotopography of the Mississippian erosion surface.
B.-F. Sand body occurrence in subsurface shown in successive 50
foot stratigraphic intervals.
40
Karst topography is believed to have
been relatively unimportant for the
presentstatus of the study because
only 22 boreholes penetrated
possible
karstfill
deposits.
Extensive cave and sink formation
probably occurred in the lower
Paleozoic
limestones
and
dolomites of northeast Iowa,
similar to those documented by
Bretz (1950) in the Ozark Region.
Northeast Iowa, where carbonate
rocks predominate, was probably
emergent long enough for ground
water flow to initiate substantial
karst formation which may
account for some of the karst
present there today.
Pennsylvanian deposits are
not observed in the northeast Iowa
karst at present suggesting either
post-Pennsylvanian erosion or
non-deposition. Further evidence
supporting
extensive
karst
development to the east and
northeast is the evidence of
Pennsylvanian-aged
materials
filling solution features in eastern
Illinois. Siever (1951) has suggested that
karst formation was not an important process
deeper in the Illinois Basin because exposed
thin limestones and shales were present in
the basin interior. By analogy in the Forest
City Basin of Iowa, the pre-Pennsylvanian
surface is underlain by predominantly sandy
limestones, shales, sandstones, and thinbedded limestones comprising the St.
Genevieve and St. Louis Formations. These
lithologies suggest limited karst development
in the Basin.
Review of additional literature on the
Mississippian erosion surface in Illinois and
western Kentucky by Howard (1979), Shaw
and Gildersleeve (1969), and Garner (1974)
provides another interpretation of the
configuration of the erosion surface. This
interpretation (Howard, 1979) suggests
multiple erosion and sedimentation episodes
under alternating arid and humid paleoclimatic conditions eventually developing a
multiple anastomosing drainage system in
contrast to the dendritic system favored by
Siever (1959). Shaw and Gildersleeve (1969)
describe an anastomosing paleodrainage
system on the erosion surface in western
Kentucky. Hansen (1978) and Hooper
(1978), in their study of sandstone bodies in
Marion County, had sufficient
subsurface data to indicate that
the erosion surface had aspects of
a rectangular drainage pattern
with valleys up to 200 feet deep
incised into a predominantly
Mississippian limestone surface.
Comparion of their paleocontour
map with that of Shaw and
Gildersleeve
(1969)
shows
remarkable similarity. This may
be an indication that an
anastomosing
paleodrainage
system is present in parts of Iowa.
In view of the limited subsurface
data for vast parts of the erosion
surface within the Forest City
Basin, the radial dendritic pattern
is tentatively accepted until
further data become available.
Figure 10. Index map showing well control and direction of crosssections. 1.-8. Structure sections showing distribution and occurrence
of sandstone bodies in the subsurface with reference to the Ardmore
Limestone.
41
A study of the occurrence
of sandstone bodies to the
paleotopographic
surface
indicates a variable relationship.
Near the Cherokee outcrop area
where data are abundant, a strong
correlation of sandstone bodies to
paleodrainage exists (Hansen,
1978). The correlation fails in a
down
dip
(southwesterly)
sands has occurred over valleys and
sometimes over hills in the
paleotopography. Unit B sandstones
show
little
relationship
to
paleotopography. Tracing of bodies
between cross sections on the basis
of thickness, vertical position, and
paleovalley trends supports the idea
of a dominant southwest drainage.
The sand bodies occur at all
stratigraphic levels in the Cherokee.
This can be readily seen in the maps
(Figs. 9B-9F) by Chamberlain
(1980) in which the percent sand in
each 50-foot interval below the
Ardmore Limestone was plotted; the
figures show the areas with greater
than 25% sand in the logs
superimposed on the Mississippian
paleotopo-graphic map. This study
supports the following observations:
1. The basal Cherokee Group
sediments contain more widely
distributed sand and fewer channel
sandstones.
Figure 10,continued.
direction going deeper into the basin.
Paleovalley control exists, as does sand body
stacking; the weaker correlation is believed
to be the result of fewer data points rather
than lack of paleovalley control.
Derynck (1980) approached the
subsurface sand study by drawing a series of
cross sections trending NW and SW and by
subdividing the Cherokee into two units
above and below the Ardmore Limestone,
which is recognizable throughout the
subsurface. Unit A includes the section
below the Ardmore, unit B above the
Ardmore. The sections (Fig. 10) show that
unit A contains more sandstone than unit B
and that drainage systems are confined, in
many cases, to paleovalleys. Stacking of
42
2. A broad south-trending
belt of basal sand occurs along the
western edge of Iowa and probably
represents sediment derived from the
Nemaha Ridge to the west.
3. The 2nd and 3rd intervals between
150 and 200 feet below the Ardmore contain
the most channel sands.
4. Cherokee Group sand bodies
correlate closely with topographic lows on
the erosion surface in the Mississippian
rocks, and consequently stacked sand bodies
occur within narrow geographic ranges.
It was also found that isochore
contours of the Cherokee Group correlate
well with paleotopographic trends and
percent sand contour trends.
In a subsurface study of the
continuation of the sand bodies outcropping
in the Fort Dodge area of north-central Iowa
with the subsurface Cherokee sand bodies in
a down dip direction 60 miles distant, a
strong correlation was found to exist. The
sand bodies are confined to southwest
trending valleys developed on the
Mississippian erosion surface (Fig. 11).
Continuing in a southwest direction into the
deeper part of the basin, considerable subsurface sandstone is present. The correlation
is not as strong primarily because the number
of data points diminishes rapidly toward the
deeper part of the basin. The strong
southwest current direction indicated in the
outcropping sandstone bodies and their close
relationship to the subsurface sand bodies
basinward strongly supports a southwest
paleodrainage direction during Cherokee
time.
PALEOENVIRONMENT
The encroachment of the Middle
Pennsylvanian sea upon the Mississippian
erosion surface led to deposition of the
predominantly clastic and coal bearing
Desmoinesian Series in the Forest City
Basin. Paleodrainage and
sandstone body occurrence
suggest the existence of
deltaic systems fringing the
Basin during deposition of the
Cherokee Group. Similar
conditions probably existed
during deposition of the
Virgilian and Missourian
Series although marine transgression is believed to have
advanced farther to the
northwest
leaving
cyclic
limestone and shale deposits
overlying
the
clastic
Desmoinesian Series in the
southwestern corner of Iowa.
The Virgilian and Missourian
clastic
and
coal-bearing
equivalents
of
the
Desmoinesian Series are not
preserved in central Iowa.
Figure 11. Maps showing the distribution of sand bodies in the subsurface in the
area immediately southwest of the sandstone outcrops along the Des Moines River.
The paleodrainage is inferred form well log data.
43
Recently
published
paleoenvironmental studies by
Wanless and Wright (1979) on
Cherokee,
Marmaton,
Pleasanton, Kansas City, and
Lansing Groups for the
northern midcontinent and
Illinois
Basin
provide
considerable
information on the nature of sedimentation
and geological setting of these groups. A
summary of these studies includes the
following:
1. Deltas composed dominantly of
fine-grained sediments formed in the Forest
City Basin;
2. Deltaic clastic wedges formed
during periods of regression;
3. Widespread coals formed on
deltaic platforms constructed by one or more
deltas;
4. Most sandstone deposits over 20
feet thick are deltaic in origin;
5. The sea was present for a longer
period of time in the northern midcontinent
than in the southern midcontinent and Illinois
Basin;
6. A marine connection developed
between the northern midcontinent and
Illinois Basin when the upper Cherokee was
deposited.
As a result of these observations, the
Iowa environment for peat deposition is
considered to be a deltaic platform with
interfingering marine environments. The
44
Desmoinesian of Iowa is characteristically
more marine in nature than the equivalent
units in Illinois (especially true of the
Marmaton Group). The deposition rates were
slow in contrast to eastern or Appalachian
sedimentation and probably account for the
higher sulfur content and thinner coal beds in
Iowa.
Because the paleoenvironment during
Cherokee deposition relates to a deltaic
complex, the coal-forming peat is considered
to have formed in fluvial, upper delta, lower
delta, and possibly, lagoonal environments.
Associated sand bodies result from coarsegrained sedimentation in these environments.
The various occurrences of sandstone bodies
can be related to the differing depositional
conditions found in deltas, alluvial channels,
estuaries prior to drowning, coastal and
alluvial plains.
Transition between environmental
facies is believed to occur over longer distances than those observed by Horne, et al.
(1978), for deltas on the Appalachians.
Insead of a 10-mile wide transition zone for
upper to lower delta environment in
Appalachia, a 20- to 60-mile zone existed
during deposition of Cherokee strata.
Wanless (1975) indicates that the coals may
retain their identity up to 80 miles down-dip.
STOP 4 – BJORKBODA MARSH
MORAINES, KAMES, AND DRAINS
Neal R. Iverson
Iowa State University
INTRODUCTION
The flat till plains that characterize
much of the Des Moines Lobe landscape in
Iowa are punctuated by curved, broad belts
of ridges and chaotic jumbles of tight hills
and depressions.
These belts are end
moraines where the location of the ice
margin was stationary for a sufficiently long
period to localize sediment deposition.
Bjorkboda Marsh lies in a depression within
the Altamont Moraine in extreme southern
Hamilton County (Fig. 1, see also Fig. 1 of
Stop 1). The marsh and surrounding hills
provide the setting for this discussion of the
Altamont Moraine and its formation, kame
Bjorkboda Marsh
Crest of Altamont Moraine
development, and the history of tile drainage
that drastically changed the face of this part
of Iowa.
ALTAMONT MORAINE
The Altamont Moraine extends across
the footprint of the Des Moines Lobe,
curving sharply north near the edges of the
lobe and merging with the Bemis moraine in
northern Iowa (Fig. 1 of Stop 1). Near
Bjorkboda Marsh the Altamont Moraine is
roughly 20 km wide. Its southern end is
defined by a ridge several kilometers wide
that rises ~15 m above the plain to the south
(Fig. 1). This ridge is locally called Mineral
Ridge, with its crest about 5
km south of the marsh (at
about the latitude of County
Road E18). The ridge is
locally transected by linked
depressions with intervening
hummocks (hills with short
steep slopes). Kemmis (1991)
developed
a
descriptive
classification of the landforms
of the Des Moines Lobe; he
called
such
ridges
“hummocky ridge systems.”
Farther north in the moraine,
the
hummocky
terrain
continues but at lower
elevation and locally is
interrupted by river valleys
(Bettis et al., 1996).
Figure 1. Bjorkboda Marsh and the approximate crest of the Altamont
Moraine.
45
The moraine consists
partly of the till members
described at Stop 1. The
supraglacial till of the Morgan Member is
present on both the flanks of hummocks and
within linked depressions and probably
usually accounts for most the moraine’s
relief. Much of this till is expected to have
been mobilized by mass-wasting processes
during its deposition. The basal till of the
Alden Member is also present and may reach
thicknesses of 30 m beneath end moraines
(Bettis et al., 1996). Also, any core taken
from the moraine would have a high
probability of containing fluvial sediments
associated with the movement of water
beneath, through, or over the glacier margin.
Some of the hills of the moraine consist
almost entirely of sand and gravel. These are
with the climate when it advanced to its
maximum position at present-day Des
Moines. Note that rapid retreat is also
consistent with the lobe being very thin at its
maximum extent; ice moves from high to
low elevations as a glacier thins during
extending flow, which leaves the glacier
especially vulnerable to rapid wastage once
the period of rapid flow and advance have
ended.
The Altamont Moraine was deposited
~ 13,500 radiocarbon years ago. Radiocarbon
dates indicate that the lobe’s margin had
apparently retreated to the location of the
Altamont Moraine, a distance of about 75 km
up-flow from Bemis Moraine (Fig. 1 of Stop
1), in ~300 years (Bettis et al., 1996). This
rapid retreat reinforces the hypothesis that
the Des Moines Lobe was out of balance
One popular myth regarding endmoraine formation, promulgated by the
Bjorkboda Marsh interpretive sign, should be
put to rest before continuing further: most
end moraines do not form as a result of
glaciers pushing sediment in front of them
like a bulldozer. Such moraines, called push
moraines, can form but are usually small
Although the topography of the
Altamont Moraine is distinct from flatter
regions to the north and south, the end
moraines of the Des Moines Lobe have
generally lower relief than those formed by
lobes of the Laurentide Ice Sheet that
were farther north (e.g., Attig et al;
1989). These northern lobes, such as the
Superior Lobe in Minnesota and the
Green Bay Lobe in Wisconsin, built far
more conspicuous end moraines, perhaps
due to the margins of these lobes being
frozen to the bed. Frozen glacier
margins, by enhancing compressive flow
near the margin, are thought by some to
increase rates of supraglacial till
accumulation (Hambrey et al., 1997).
The observation that the Lake Michigan
Lobe of Illinois also built low-relief end
moraines supports this hypothesis; this
Figure 2. Kame visible from the parking lot of Bjorkboda
lobe, like the Des Moines Lobe, was
Marsh.
thin, sloped gently, and advanced far
south of where there is evidence of a
kames; an example is the peaked hill a few
frozen margin.
hundred meters northeast of the Bjorkboda
Marsh parking lot (Fig. 2).
END-MORAINE FORMATION
46
cored with ice for many years after retreat of
the active margin (Fig. 3a). Differences in
the rate of surface melting cause depressions
on the ice surface. These depressions can fill
with sediment from higher adjacent regions
due to debris transport by mass-wasting or
melt water.
Wherever lots of debris
accumulates on the ice surface, either due to
locally high englacial debris concentrations
or sediment redistribution to topographic
lows, these areas eventually become hills
after ice has fully melted. The longer the
glacier margin remains in one area, the more
sediment is carried to that area and the more
prominent the resultant moraine.
ridges with widths that do not exceed a few
tens of meters. They are thus vastly different
in scale from the wide zones of hummocky
topography that characterize moraines of
major ice masses.
Most end moraines, regardless of
their scale, form as the simple result of
movement of ice to the glacier margin (Fig.
3). This ice contains debris that is deposited
on the glacier surface as it melts. The bed is a
major source for debris, so most debris is
contained in ice near the bed and thus melts
out on the ice surface near the margin. The
debris helps insulate the underlying ice,
which can result in moraine ridges that are
In the case of the Altamont Moraine,
owing to the thinness of the Des Moines
Lobe, debris melted out at the lobe’s surface
over a wide area, producing a broad belt of
hummocky topography (Fig. 3b). Figure 4
illustrates the margin of a modern glacier in
Alaska that is mantled in debris well back
from its margin. Perhaps the margin of the
Des Moines Lobe looked something like this
when it built the Altamont Moraine (sans
mountains).
a
Kemmis (1991) suggested that the
Des Moines Lobe fully stagnated over much
of its area after multiple advances, enabling a
glacial-karst system to develop after each
advance. In such a system, meltwater would
b
over ice
Figure 3. Formation of (a) single moraine crest at a
steep glacier margin and (b) broad belts of hummocky
moraine at the margin of a thin ice sheet. The latter is
a better model for the Des Moines Lobe (modified
from Bennet and Glasser, 1996).
47
Figure 4. Till-mantled terminus of an Alaskan
glacier.
would have left hills of bedded sand and
gravel, together with some till accumulated
through mass wasting (flow tills) (Fig. 5).
Removal of adjacent ice during melting
would have caused collapse of bedding near
kame margins. An alternative is that sand and
gravel accumulated at the bases of moulins,
near-vertical shafts in the ice that conduct
water to the bed. These shafts in modern
glaciers can be sustained, despite the pressure
percolate through the ice and melt it, causing
collapse of the debris mantle and resultant
systems of linked depressions. This
explanation for linked depressions seems
plausible for a debris-mantled marginal zone
of any glacier that has become isolated from
actively flowing ice up-glacier; such isolation
of marginal zones of stagnant ice in the form
of ice-cored moraine complexes is common
during glacier retreat. Thus, although the
conceptual model of Kemmis (1991) for
linked depressions may be largely correct, an
aspect of the model that seems superfluous is
stagnation of the Des Moines Lobe at a
regional scale.
KAME FORMATION
(3)
A kame is a landform composed
primarily of stratified drift deposited in
contact with glacier ice. The kames of the
Des Moines Lobe are mound-like features,
usually hundreds of meters in diameter and
6-13 m high (Bettis et al., 1996). They
consist predominantly of sand and gravel,
although uncommonly diamictons are present
as isolated pods. Kame margins commonly
display evidence of collapse in the form of
steep faults or bedding that dips away from
the kame center (Bettis et al., 1996).
(2)
Kames of the Des Moines Lobe,
which can be classified as hill kames (Benn
and Evans, 1998), likely formed either on the
glacier surface or at the base of the glacier
where water derived from the glacier surface
encountered the bed. During supraglacial till
deposition, depressions on the glacier surface
associated with non-uniform ablation were
likely sinks for sediment that was washed in
from higher elevations. Crevasses on the ice
surface that closed or narrowed sufficiently
at depth could have also been loci for
sediment accumulation. Filling of such
depressions, together with the eventual
melting of adjacent and underlying ice,
Depressions
on glacier
surface
(1)
Figure 5. Model for hill-kame formation involving
deposition of fluvial deposits and flow tills within
depressions on the ice surface (modified from Benn
and Evans, 1998).
48
were looking for ways to optimize
productivity. Although farmers had begun
draining the landscape of the lobe in the
1870s, by the early 1900s drainage efforts
had greatly expanded. State legislation was
enacted that encouraged farmers to band
together in “drainage enterprises,” which
stimulated drainage improvements and
resulted in the formation of 3000 drainage
districts in Iowa. During the decades up to
1940, over 126,000 km of underground
drainage tiles were installed, most of them
within the area spanned by the Des Moines
Lobe (Fig. 6b). Most tiles were installed 1.01.5 m below the surface, and trenches for
them were dug by hand. Networks of tiles
typically emptied into ditches, which were
excavated to feed creeks and rivers. These
drainage efforts were highly successful.
Today, for example, in Hamilton County,
on the ice that closes cavities at depth, due to
heat dissipation and consequent melting of
ice associated with the falling water. This
water can carry sediment from the glacier
surface and pile it in mounds at the bed. Such
kames are sometimes called moulin kames
and are usually difficult to distinguish from
kames that accumulate on the ice surface.
*DRAINING WETLANDS
“clay and cement tile is the most important
industry in Iowa and . . . Iowa puts out more
of these products than any other state, all
because of the great drainage movement now
under way.” 1913, D.A. Marston, an
engineer at Iowa State University.
“we should adopt a system of selective
drainage in place of the reckless system
which operates on the groundless assumption
that all drainage is beneficial.” 1920s, B.
Shimek, a geologist at the University of
Iowa.
a
Wetlands like Bjorkboda Marsh are
few and far between in central Iowa,
surrounded by tracks of row crops that are
vast in comparison. This was not true in the
late 1800s. Over the footprint of the Des
Moines Lobe, wetlands were the rule, rather
than the exception. A horseback ride through
the countryside at that time would have
followed curving roads that snaked along
high ground above adjacent ponds and
sloughs. Patches of corn were common but
isolated by large tracks of land that were too
wet for too long each year for crops. These
wetlands reflected the poorly developed
drainage of the Des Moines Lobe landscape;
the well-integrated river networks that drain
most of Iowa are far more subdued within the
margins of the Des Moines Lobe (Fig. 6a).
b
Figure 6. (a) Relief map of Iowa showing river
networks. Compare with (b) land in Iowa subject
to "drainage enterprises" in 1940.
By the late 1800s, however, the
settlement of Iowa was over, and farmers
49
which contains Bjorkboda Marsh, less than
0.5% of original wetlands remain.
Moines Register (July 23, 2000) written by historian
Lowell J. Soike.
Intensive
tile-drain
installation
transformed the distribution of agricultural
productivity and wealth in the state. By 1910
the formerly swampy land of the Des Moines
Lobe supported farms with greater cropproduction capacity and wealth than farms in
southern Iowa, where the land was naturally
better drained but more rolling and better
suited to pasture. To try to compete, farmers
in southern Iowa expanded row-crop
coverage, commonly increasing soil erosion
and worsening their already decreasing land
values. At the same time the formerly
diversified but modest farms in north-central
Iowa expanded and focused on a few
lucrative cash crops.
REFERENCES
Attig, J.W., D.M. Mickelson, and L. Clayton,
1989,
Late
Wisconsin
landform
distribution and glacier-bed conditions in
Wisconsin, Sed. Geol., 62, 399-405.
Benn, B.I., and D.J.A. Evans, 1998. Glaciers
and Glaciation, Oxford, London.
Bennet, M.R., and N.F. Glasser, 1996.
Glacial Geology, Ice Sheets and
Landforms, Wiley, New York.
Bettis, E.A., D.J. Quade and T.J. Kemmis,
1996. Overview, In Bettis, E.A., D.J.
Quade and T.J. Kemmis, eds., Hogs,
Bogs, and Logs: Quaternary deposits and
environmental geology of the Des
Moines Lobe. Iowa Department of
Natural Resources, Guidebook Series, 18,
1-79.
Today, of course, the benefits of
wetlands are better appreciated.
Their
positive effects on water quality, flood
control, wildlife, and recreation highlight the
need to preserve existing wetlands and have
motivated numerous restoration efforts. Of
course these efforts fall orders-of-magnitude
short of restoring the wetlands that once
existed over the Des Moines Lobe landscape
(Fig. 6b). Wetlands like Bjorkboda Marsh,
therefore, provide small glimpses of a part of
the Iowa landscape and ecology that, for the
most part, is probably gone forever.
Hambrey, M.J., D. Huddart, M.R. Bennett,
and N.F. Glasser, 1997, Genesis of
‘hummocky moraines’ by thrusting in
glacier ice: evidence from Svalbard and
Britain, J. Geol. Soc., 154, 623-32.
Kemmis, T.J. 1991. Glacial landforms,
sedimentology,
and
depositional
environments of the Des Moines Lobe,
northern Iowa (Ph.D. thesis, University
of Iowa).
*Data and quotations regarding the tile-drain history
of Iowa were obtained from an article in the Des
50
STOP 5 – DES MOINES RIVER VALLEY
LATE-WISCONSINAN HISTORY OF THE UPPER DES MOINES RIVER
Neal R. Iverson
Iowa State University
when and why the valley formed and the
Wisconsinan terraces within the valley. The
discussion leans heavily on the work of
Arthur Bettis and colleagues (Bettis and
Hoyer, 1986; Bettis et al., 1988; Bettis and
Hajic, 1995). The development of the valley
provides a good example of how river history
in the Midwest is commonly closely tied to
glacial history.
INTRODUCTION
Except for some glaciers at very high
altitudes or latitudes, glaciers seasonally
produce copious volumes of melt water,
resulting in outwash streams with high
discharges of both water and sediment. This
may have been especially true for the Des
Moines Lobe, which advanced into a climate
too warm to sustain the lobe. Melting and
retreat of the lobe in Iowa occurred rapidly
from ~14,000-12,000 radiocarbon years
before present (RCYBP). The outwash
stream that eroded the deepest and widest
valley within the footprint of the lobe was the
Des Moines River. This summary of the lateWisconsinan history of the river focuses on
CHRONOLOGY
Despite the prominence of the upper
Des Moines River valley (the part of the river
north of Des Moines), it is younger than the
smaller Skunk River valley to the east (see
Stop 1). While the margin of Des Moines
Lobe was at the Altamont moraine
~13,500 RCYBP (about 14 km north
of this stop), its major outwash
streams is this area were the Skunk
River and Beaver Creek, which is
now a small stream about 12 km to
the west. There is no evidence that
the Des Moines River existed at that
time. Instead, as argued by Bettis et
al. (1988), the river did not develop
until the lobe’s margin was building
the Algona Moraine, at ~12,300
RCYBP and about 115 kilometers to
the north of this stop (Fig. 1). The
evidence for this late development of
the river is convincing:
(1)
Stop 5
radiocarbon dates on wood from the
highest Des Moines River terrace near
Saylorville Dam just north of Des
Moines are essentially the same as
those for the Algona Moraine; (2) the
Figure 1. Course of the Des Moines River and end moraines
region north of the Algona Moraine
of the lobe (Bettis et al., 1988).
51
52
Figure 2. Wisconsinan-age terraces along a reach of the Des Moines River spanning the Altamont Moraine (from Bettis et al., 1988). Noah
Creek is nearest to Stop 5.
10 km
downcutting exceeded 65 m
in the vicinity of Boone.
TERRACE FORMATION
Formation of the Des
Moines
River valley included
High Late
Wisconsinan
the
formation
of many
Terrace
Rose Hill
Wisconsinan-age
terrace
Cemetary
segments. Periods of lateral
Low Late
migration of the river
Wisconsinan
Terrace
necessary to form the flat
treads of terraces were
punctuated by periods of
downcutting. Two of these
Holocene
Terraces
terrace surfaces are visible
from Rose Hill Cemetery
(Fig. 3). These are point-type
terraces (Fig. 4) consisting of
up to 6 m of alluvium resting
on pre-Illinoian till. Thus,
Figure 3. Location of Stop 5 at Rose Hill Cemetery. Terrace
because the terraces reflect
demarcation is based on Bettis and Hajic (1995).
planation of a deposit that predates the river, strictly they
forms the headwaters for both the East and
are strath rather than alluvial terraces. Bettis
West Forks of the Des Moines River, and
et al. (1988) called them “benched” terraces.
terraces of the West Fork just outside the
Gravel operations active in the 1980s
moraine contain kettles, indicating ice was at
provided several good exposures of the
the Algona Moraine when river terraces were
alluvial stratigraphy in these terraces. Bettis
being formed; (3) the Des Moines River is
et al. (1988) divided the alluvium into three
incised into older landforms of the Des
increments (Fig. 5). The lower increment
Moines Lobe, including the Altamont
consists of sand and gravel, one to several
Moraine, with terraces of the river well
below the elevation of the moraine (Fig. 2).
Holocene
Terraces
Erosion of the valley occurred over a
period of 1,000-1,600 years based on the
ages of the oldest Holocene sediments at the
base of the valley. By ~ 11,000 RCYBP, the
Des Moines Lobe margin had retreated far to
the north of Iowa, and the late-Wisconsinan
entrenchment of the river was over. The
river in the area around Boone eroded
through the deposits of the Des Moines Lobe,
then through far older pre-Illinoian tills, and
finally into Pennsylvanian bedrock. Total
Figure 4. Point-type terraces like those at Stop 5
(from Bettis et al., 1988).
53
Figure 5. Stratigraphy of a Wisconsinan-age terrace, showing three characteristic increments. This
section was about 500 m southeast of Stop 1 (from Bettis et al., 1988).
meters thick, with a wide variety of bedding
structures. Cross beds indicative of either
migrating bedforms (e.g., dunes, ripples) or
point-bar accretion are common, as are
channel fills 15-50 m wide. Sediments of the
so-called
middle
increment,
which
unconformably overly the lower increment,
are much coarser. Middle increment
sediments consist of cobble gravels that are
massive to crudely planar bedded and very
poorly sorted. Sediments of the upper
increment are usually less than 1 m thick and
are fine grained (sandy loam to loam). They
are either massive or fine upward and
abruptly contact the middle increment. This
succession of sediments was observed at
each of three terraces studied by Bettis et al.
(1988).
suspended load, approaching deposition
typical of a hyperconcentrated flow. They
interpret the fine-grained sediments of the
upper increment to be waning-flow/overbank
deposition associated with the floods that
deposited the middle increment. The most
likely cause of a jökulhlaup from the Des
Moines Lobe margin at the Algona Moraine
would be rapid draining of a proglacial lake.
Proglacial lake sediments are present near the
Algona Moraine; the lake, called Glacial
Lake Jones (Kemmis, 1981) could have
drained catastrophically (Bettis et al., 1988),
but there is no evidence that it did. Another
possible source for a catastrophic flood
would be water that is sometimes stored
subglacially during rapid glacier flow and
released when rapid flow stops (e.g., Kamb
Bettis et al. (1988) interpret the lower
increment to be the result of normal
deposition in a braided river, with fluctuating
water and sediment discharges. They
attribute the middle and upper increments,
however, to infrequent very high magnitude
floods called jökulhlaups, which occur when
a large reservoir of water is rapidly drained
from an ice-marginal or subglacial
environment. They interpret the poor sorting
of the middle increment to be a result of
simultaneous deposition of bedload and
Figure 6. Massive, poorly sorted gravel and cobble
outwash in Scotland unrelated to jokulhlaup
deposition (Maizels, 2002). Flow was from left to
right.
54
middle increment was deposited during the
rising stages of floods when sediment
discharge was high but before the waterdischarge peak; erosion to a new lower
terrace level was thought to have occurred
during the subsequent peak in water
discharge, after the river’s sediment load had
begun to decrease.
et al., 1985).
An alternative interpretation is that
both the lower and the middle increments
reflect deposition in a high-discharge braided
stream, with the upper increment a result of
normal overbank deposition. The poorly
sorted cobble gravel may well have been
deposited during high discharges, but the
need to invoke a jökulhlaup is unclear. For
example, Maizels (2002) describes facies
very similar to those of the middle increment
as being common in glacial outwash deposits
(Fig. 6) and not necessarily related to
jökulhlaups.
One difficulty with this hypothesis is
that, without invoking a long-term change in
either water discharge or sediment supply, it
appears to violate the concept that rivers
adjust their slopes toward a graded condition
(e.g. Mackin, 1948). A graded river is one
that is neither eroding nor aggrading its bed
because over a period of years it has adjusted
its slope to carry the sediment supplied to it
with the available water discharge. Although
no river strictly achieves this graded
condition, the slope of an alluvial river is
always adjusting to approach this equilibrium
in which the water discharge is just sufficient
to transport the sediment supply. In the
absence of tectonic uplift or a reduction in
base level, an increase in the ratio of water
discharge to sediment supply in a graded
river will cause it to downcut. Only an
increase in this ratio that is sustained can
cause sustained downcutting. This is why,
although a flood commonly scours a river
channel, that scour is usually transient; the
scoured river bed eventually aggrades to its
former level after the flood because the ratio
of water discharge to sediment supply is the
same before and after the flood. Thus, it is
unclear why each of the jökulhlaups
hypothesized by Bettis et al. (1988) would
have caused downcutting of the river that
was sustained between jökulhlaups. Instead,
the scour caused by one of these floods,
rather than initiating a new lower terrace
level, would have been erased by aggradation
of the river bed up to an elevation
comparable to that before the flood. Sporadic
floods alone, therefore, do not provide a
WHY DID THE RIVER DOWNCUT?
Perhaps the most important and
difficult question to answer regarding the
Des Moines River valley is why the river
downcut. Except for a tendency for outwash
streams to incise very near glacier margins,
most proglacial streams generally aggrade
their beds rather than erode them. This
reflects the high sediment fluxes brought to
glacier margins by a combination of ice and
water flow; the highest sediment fluxes in the
world are from catchments containing
glaciers, which are generally more efficient,
for a given catchment, than rivers in eroding
and transporting sediment (Hallet et al.,
1996).
Bettis et al. (1988) attributed the
cutting of the valley to sporadic jökulhlaups.
These authors interpreted the downcutting of
the river from one terrace level to another to
be the result of these severe but infrequent
floods. Between jökulhlaups, normal flows
were assumed to have resulted in lateral
migration and deposition of the lower
increment. This explanation for downcutting
is
consistent
with
these
authors’
interpretation of the middle increment as a
jökulhlaup deposit. They argued that the
55
(1988) was gone. Without a commensurately
large reduction in sediment supply,
downcutting stopped. During the Holocene
the river has underwent periods of both
downcutting and aggradation, but with little
net incision (Bettis et al., 1988).
completely satisfactory mechanism for the
erosion.
Hypotheses for erosion of the valley,
therefore, need to consider both the water
discharge of the river and its sediment
supply. Why was the water discharge
sufficiently large relative to the sediment
supply to cause sustained erosion, despite the
general tendency for outwash streams to
aggrade their beds? This question cannot be
answered with much certainty. One
possibility is that sediment produced by the
glacier was stored primarily near the glacier
margin, perhaps in an ice-marginal lake.
Thus, high seasonal water discharges from
the glacier would have been accompanied by
low sediment loads. The Minnesota River
valley to the north formed for this reason in
the wake of the retreating Des Moines Lobe
(Wright, 1973). A contributing factor could
have been vegetation and its effect on
reducing sediment supply. If the coniferous
forest characteristic of this period in Iowa
(e.g., Bettes and Hajic, 1995) occupied the
tributary drainages feeding the Des Moines
River, this would have enhanced storage of
sediment in proximal areas and reduced the
sediment supply to the river downstream.
Another possibility is that the hypothosis of
Bettis et al. (1988) is correct—that sporadic
catastrophic floods eroded the valley—but
that the time required to restore river
equilibrium was long relative to the
frequency of jökulhlaups, such that the slope
and elevation of the river were always
grossly out of equilibrium with the
predominant river discharge and sediment
supply.
REFEREENCES
Bettis, E.A. and E.R. Hajic, 1995. Landscape
development and the location of evidence
of evidence of Archaic cultures in the
Upper Midwest. Geol. Soc. Am. Spec.
Pap., 297, 87-113.
Bettis, E.A. and B.E. Hoyer, 1986. Late
Wisconsinan and Holocene lanscape
evolution and alluvial stratigraphy in the
Salorville Lake Area, central Des Moines
River valley, Iowa. Iowa Geol. Surv.
Open File Rep., 86-1.
Bettis, E.A. and 6 others, 1988. Natural
History of Ledges State Park and the Des
Moines River Valley in Boone County.
Geol. Soc. Iowa Guidebook, 48.
Hallet, B., Hunter, L., and Bogen, J., 1996,
Rates of erosion and sediment evacuation
by glaciers: A review of field data and
their implications. Glob. Planet. Change,
12, 213-235.
Kamb, B. and 7 others, 1985. Glacier surge
mechanism:
1982-1983
surge
of
Variegated Glacier, Alaska. Science,
227(4686), 469-479.
Kemmis, T.J., 1981. Glacial sedimentation
and the Algona Moraine in Iowa. Geol.
Soc. Iowa Guidebook, 35.
More certain is the reason why
downcutting stopped. By 11,000 RCYBP
glacier ice was gone from the Des Moines
River basin, and the source of water that kept
the seasonal discharge of the river high or
produced the jökulhlaups of Bettis et al.
Mackin, J.H., 1948. Concept of a graded
river. Geol. Soc. Am. Bull., 48, 813-893.
56
Wright, H.E., Jr., C.L. Matsch and E.J.
Cushing, 1973. Superior and Des Moines
lobes. In Black, R.F., R.P. Golthwait and
H.B. Willman, eds., The Wisconsinan
Stage. Geol. Soc. Am. Mem. 136, 153185.
Maizels. J., 2002. Sediments and landforms
of modern proglacial
terrestrial
environments. In Menzies, J., ed.,
Modern and Past Glacial Environments,
Butterworth Heinemann, Oxford, 279316.
57
STOP 6 – MISSISSIPPIAN AND PLEISTOCENE GEOLOGY AT MONTOUR QUARRY
Jane Pedrick Dawson and Matt Graesch
Iowa State University
deposits in east-central Iowa, we will follow
the ledges and ramps down to the quarry
floor, examining notable sedimentary
features along the way.
INTRODUCTION
Lower Mississippian (Kinderhookian)
limestones and dolomites are quarried here
by Wendling Inc. for use in highway and
road construction. Limestone dimension
stone was cut for one year in the 1990’s for
the purpose of restoring the old state
historical building in Des Moines.
OVERVIEW OF
STRATIGRAPHY
KINDERHOOKIAN
Witzke and Bunker (1996) concluded
from their study of Paleozoic cratonal
sediments in Iowa that most deposits were
representative of either inner shelf or middle
shelf epicontinental marine environments
(Fig. 1).
Inner shelf deposits are
characterized by shallow subtidal to peritidal
facies that preserve shallowing upward
cycles. Middle shelf facies are dominated by
subtidal deposits, which may have remained
submerged through successions of relative
sea-level changes.
The carbonates are underlain by a
siltstone that is the basal Mississippian unit
in this area. These strata are interpreted as
mainly shallow water, inner shelf deposits
that record the on-lap of Early Mississippian
epicontinental seas onto Upper Devonian
bedrock. The Montour Quarry location is
very close to the erosional edge of the
Mississippian
in
east-central
Iowa.
Mississippian bedrock extends only 10 km to
the east, where Devonian bedrock becomes
the youngest rock preserved (Glenister, 1987;
Witzke et al., 2003).
By
combining
stratigraphy,
biostratigraphy, and regional depositional
patterns, Witzke and Bunker (1996)
constructed a relative sea-level curve for the
Mississippian in SE Iowa (Fig. 2). These
data can also be related to east-central Iowa.
Summarizing their interpretations, the first
Mississippian transgressive-regressive (T-R)
cycle is recorded by deposition of
Kinderhookian thin transgressive, openmarine limestones (McCraney Formation)
onto Devonian strata in SE Iowa. In T-R
Cycle 2, more widespread shales and
siltstones of the Prospect Hill Formation
overlap the McCraney edge to lie on
Devonian bedrock across much of Iowa.
This cycle is characterized by progradational
silt deposits and regional expansion of
carbonate sedimentation.
T-R Cycle 3
deposits overlap the Prospect Hill formation
Quarrying operations at Montour
have removed the Pleistocene overburden,
exposing the top of the Mississippian
bedrock. Glacial striations are evident in
fresh exposures, but weather quickly. Cuts
in the overburden next to quarry walls reveal
three paleosols plus Wisconsinan aeolian
deposits.
We will drive past the quarry
entrance and leave our vehicles parked
alongside 290th Street and reach the quarry
by walking south through the bean field. We
will enter the quarry from the top and
examine the Pleistocene paleosols first.
After an explanation of stratigraphic
relationships of Lower Mississippian
58
Mississippian is shown in Fig. 4, but it is not
strictly followed.
to lie on Devonian strata in NW Iowa. In SE
Iowa, this cycle is characterized by the cherty
carbonate middle shelf deposits of the
Wassonville-Starrs Cave formations, whereas
in east-central and northern Iowa it is
The Kinderhookian carbonate rocks
between the Prospect Hill and Gilmore City
Figure 1. Large-scale marine epicontinental facies groupings of Witzke and Bunker (1996).
formations were originally named the
Hampton Formation, which included in
ascending order, the Chapin (limestone),
Maynes Creek (dolomite), Eagle City
(limestone), and Iowa Falls (dolomite)
members (Van Tuyl, 1925; Laudon, 1931).
The term Hampton has been discontinued
(Harris, 1947; Woodson and Bunker, 1989;
Witzke et al., 2001), but the member names
are still in use. They are now applied to
stratigraphic intervals that differ from their
original designations, as described below.
represented by thicker carbonate inner shelf
deposits of the Chapin, Maynes Creek, and
Eagle City units, the focus of our field trip.
A schematic cross-section of the entire
Mississippian section in Iowa (Fig. 3) by
Witzke and Bunker (2001) shows their
interpretation and correlation of inner-shelf
and middle-shelf facies from NW to SE
Iowa.
The
stratigraphic
nomenclature
applied to the Kinderhookian strata in central
and northern Iowa is somewhat confusing,
with various groups adopting different
naming practices. This confusion resulted
from type sections defined at localities where
the lower or upper contacts were obscured or
missing, difficulty in correlating dolomites,
and a lack of subsurface data which led to the
correlation of similar lithologies from
different depositional cycles (Woodson and
Bunker, 1989). The Iowa Geological Survey
Bureau’s current stratigraphic column for the
The Chapin type section of
fossiliferous crinoidal to oolitic limestone in
Franklin County (Van Tuyl, 1925; Laudon,
1931) may not be at the same stratigraphic
interval as the oolitic limestone referred to as
the Chapin in the LeGrand area (Woodson
and Bunker, 1989). However, in common
usage, the term Chapin is used to refer to the
thick
oolitic
limestone
in
the
LeGrand/Montour area.
This unit is
57
recommendation that the Maynes Creek
formation be expanded to include the entire
stratigraphic interval between the Prospect
Hill and Gilmore City formations. This
recommendation was made after Burggraf
(1981) concluded that the Maynes Creek type
section occupies a higher stratigraphic
position than originally proposed. The Iowa
Dept. of Transportation (Iowa DOT) views
the Maynes Creek type section as cherty
Iowa Falls dolomite, and they define the
Maynes Creek as the cherty dolomite
between the Chapin oolite and the Eagle City
limestone/dolomite (B. D. Gossman and M.
R. Dawson, pers. comm., 2005).
considered correlative with the skeletal to
oolitic grainstone Starrs Cave Formation of
southeast Iowa (Lawler, 1981; Glenister,
1987, Witzke and Bunker, 1996).
The Maynes Creek type section in
Franklin County was defined as the cherty
dolomite above the Prospect Hill siltstone
and Chapin oolite and below the Eagle City
limestones (Van Tuyl, 1935; Laudon, 1931).
In their bedrock geology map of northcentral Iowa, Witzke et al. (2001) locally
include the Chapin and Eagle City as
members of the Maynes Creek Formation,
following Woodson and Bunker’s (1989)
The Eagle City type area is in
Hardin County and includes fossiliferous
oolitic limestones and dolomites (Van Tuyl,
1925; Laudon, 1931). These early workers
described a lithographic limestone and
fossiliferous oolitic limestone in the upper
Eagle City at Iowa Falls. Woodson and
Bunker (1989) put the contact between their
expanded Maynes Creek and Gilmore City
formations between the lithographic unit
and the fossiliferous oolitic bed. However,
the Iowa DOT considers these limestones
the “Weldon Ledge” and assigns them as
the basal unit of the Gilmore City
Formation (B. D. Gossman and M. R.
Dawson, pers. comm., 2005).
Other
workers have grouped the Eagle City beds
with either the Maynes Creek Formation
(Witzke et al., 2001) or in an informal
stratigraphic grouping with the Gilmore
City Formation where regional stratigraphic
relationships are uncertain (Witzke et al.,
2003).
Dolomite overlying the Eagle City
beds was originally defined as the Iowa
Falls dolomite by Van Tuyl (1925). The
type section in the Iowa River gorge in the
Figure 2. Mississippian relative sea-level curve for
city of Iowa Falls is now recognized as a
southeast Iowa (Witzke et al., 1990; Witzke and Bunker, dolomitized lateral facies of limestones of
1996) (from Witzke and Bunker, 1996).
58
Figure 3. Schematic northwest-southeast stratigraphic cross section of Mississippian
strata across Iowa, spanning inner-shelf and middle shelf facies tracts (from Witzke
and Bunker, 2001).
the Gilmore City Formation (Thomas, 1960;
Woodson and Bunker, 1989). Consequently,
the Iowa Falls dolomite was included as a
member of the Gilmore City Formation by
Witzke et al. (2001). The Iowa DOT uses the
term Iowa Falls to refer to locally cherty, soft
dolomites that occur in the stratigraphic
interval
above
the
Eagle
City
limestones/dolomites
and
below
the
lithographic and oolitic limestones of their
“Weldon Ledge” (B. D. Gossman and M. R.
Dawson, pers. comm., 2005).
Keokuk Formation
Bentonsport Member
Montrose Chert Member
Burlington Formation
Cedar Fork Member
Haight Creek Member
Dolbee Creek Member
Sub-Augusta Group
Gilmore City Formation
Humboldt Oolite
Alden Limestone
Iowa Falls Dolomite
Maynes Creek Formation
Eagle City Member
Wassonville Member
North Hill Group
Chapin Formation
Starrs Cave Formation
Prospect Hill Formation
McCraney Formation
MISSISSIPPIAN (subsystem of
Carboniferous System)
Pella (Ste. Genevieve) Formation
St. Louis Formation
Waugh Member
Verdi Member
Yenruogis Sandstone
Croton Member
Spergen Formation
Augusta Group
Warsaw Formation
Figure 4.
The Iowa Geological Survey’s
stratigraphic column for the Mississippian in
Iowa.
http://www.igsb.uiowa.edu/gsbpubs/
Stratigraphy/iastratcolum2.asp
59
Bed #
Thickness
EAGLE CITY
20
19
18
17
16
15
14
13
12-8
7-6
5-4
3-1
Dolomite; reddish brown, fine to
0’-2.0’
medium grained.
Limestone; light gray to white,
1.7’-2.0’
oolitic, variable in thickness.
Limestone; light gray, medium to
1.3’-1.5’
coarse grained, crinoidal, as one or
two beds.
Dolomite; calcareous, gray-brown,
3.7’
dense, may become limestone in some
areas, a .9’ silty greenish-gray bed at top,
rip-up clasts.
Dolomite; brown to gray-brown, fine 2.2’-2.5’
to medium grained, alternates from
soft and powdery to hard and laminated
with occasional calcite nests, banded
with iron oxide staining throughout.
Dolomite; brown to gray-brown, fine
3.0-3.9’
to medium grained, top 2.4’ breaks
into irregular beds and may grade to
limestone, dense gray bed in middle,
chert band at base.
Dolomite; brown to light gray-brown,
4.5’
fine to medium grained, .8’ light
gray crinoidal limestone near middle
of bed, occasional chert near top.
Dolomite and limestone; light gray to 3.0’-3.9’
brown, fine to medium grained,
crinoidal lenses near base, may grade
to limestone in top 1.2’.
MAYNES CREEK
Dolomite; brown, gray-brown color
±24.0’
banding in some areas, fine to
medium grained, scattered gray chert
bands throughout but most abundant
in lower half, ripple marks at upper ledge.
CHAPIN
Limestone, gray to white oolite,
3.5’
flaggy-bedded, coarsely oolitic,
many brachiopods throughout.
Limestone; gray to white oolite,
4.5’
massive, may have two thin beds at
top, many brachiopods throughout.
Limestone, gray to white oolite,
7.5’
coarsely oolitic, as 3 beds, many
brachiopods throughout.
Figure 5. The Iowa Dept. of Transportation’s stratigraphic column for Montour Quarry (modified
from Gossman, 1985).
60
but that the ooids formed on migratory shoals
and were dispersed in large non-crossbedded
sheets.
The correlative Starrs Cave
Formation in southeastern Iowa contains
cross-bedded oolitic sands and is more
representative of an ooid factory area
(Lawler, 1981). Whether or not the Chapin
ooids originated in southeast Iowa, it is clear
that they do not have local origins.
STRATIGRAPHY AT MONTOUR
QUARRY
For the purposes of this field trip, we will use
the Iowa DOT designations for the
Mississippian units ledged and quarried at
Montour (Figs. 5 and 6).
Lawler (1981) noted a hardground
approximately in the middle of the Chapin
section in both Montour and LeGrand
Quarries (Fig. 7), and found evidence that
the contact between the Chapin and the
Maynes Creek dolomite is another probable
hardground. The contact is an abrupt,
styolitic and bored surface, and ooids are
truncated along the walls of vertical borings
where Maynes Creek sediments infilled
Chapin sediments (Fig. 8).
Figure 6. Annotated photo of Montour Quarry wall with
the Iowa Dept. of Transportation's stratigraphic divisions.
Prospect Hill
The basal Mississippian Prospect Hill
formation is a dolomitic siltstone here
(Glenister, 1987) and is not quarried.
The Chapin oolite is ledged at
Montour Quarry because it is an excellent
source of concrete aggregate. It has a 3i
Portland cement concrete durability class
rating from the Iowa DOT, making it
suitable for use in interstate construction (B.
D. Gossman and M. R. Dawson, pers.
comm., 2005). Dimension stone cut from
the Chapin in the 1990’s was used to restore
the old state historical building in Des
Moines.
Maynes Creek Dolomite
The Maynes Creek dolomite is really
more of a dolomitic limestone as chemical
analyses show only a small percentage of
MgO (B. D. Gossman and M. R. Dawson,
pers. comm., 2005). This unit was originally
deposited as shallow marine to supratidal
carbonate muds (Glenister, 1987). Few
fossils occur, which has been attributed to
original impoverishment and the effects of
dolomitization (Glenister, 1987), although
dolomitization is relatively minor. Abundant
Chapin Oolite
The Chapin oolite at Montour is a
thick
bedded,
fossiliferous
(mainly
brachiopods), oolitic grainstone (Lawler,
1981). The well-sorted bimodal texture in
the ooids and a general fining upward
sequence (Fig. 7), suggested to Lawler
(1981) that this was not an ooid factory area,
61
Because of the abundant chert
horizons, the Maynes Creek unit is
not approved for use as concrete
aggregate by the Iowa DOT, but is
approved for use as asphalt
aggregate, rip-rap, granular base and
surfacing, and other aggregate
products (B. D. Gossman and M. R.
Dawson, pers. comm., 2005).
Eagle City Dolomite
The Eagle City dolomite is
more of a dolomitic limestone,
averaging
a
25%
dolomite
component (B. D. Gossman and M.
R. Dawson, pers. comm., 2005).
Limestone beds include a crinoidal
limestone near the base and an oolitic
limestone near the top of the section.
Dolomitized crinoidal deposits are
often storm-derived hash beds.
Other storm deposits include a bed
with intraformational rip-up clasts
about two-thirds of the way up from
the base of the section (Fig. 11).
Well-preserved silicified brachiopods
Figure 7. Stratigraphic column of the Chapin oolite showing the
are abundant in the upper part of the
bimodal size distribution of ooids and fining upward trend. The
section. See Glenister (1987) for
arrow denotes the location of a hardground (from Lawler, 1981).
illustrations of the macrofossils
found in both the Chapin and Eagle
chert horizons are present. A shallowing
City sections.
upward sequence is indicated, capped off by
ripple marks in the top of the Maynes Creek
ledge (Fig. 9).
This area around LeGrand is perhaps
best known for the spectacular fossils that
have
been
recovered
from
Lower
Mississippian beds. Just north of LeGrand, a
world-famous nest of fully articulated
crinoids was discovered in 1933 (Laudon and
Beane, 1937) in a quarry that now serves as
the headquarters for Cessford Construction.
The crinoids came from a bed just above the
Maynes Creek – Eagle City contact (as
defined by the Iowa DOT). This same bed
also contains crinoids at Montour, but no
fossils of this quality have been found in the
area since the last great find in the 1930’s.
Native Americans from paleo times
to European contact used Maynes Creek
chert to make stone implements, likely
quarrying the chert at exposures that crop out
near the Iowa River from Franklin County
southeast to Tama County (Morrow, 1982).
Artifacts made from Maynes Creek chert
(Fig. 10) have been found over the entire
upper Midwest, indicating it was widely
traded.
62
Montour has a 3i Portland cement
concrete durability class rating from
the Iowa DOT, and is ledged wherever
it
remains
thick
enough
to
economically remove it (B. D.
Gossman and M. R. Dawson, pers.
comm., 2005).
Pleistocene Deposits
Montour
Quarry
offers
excellent exposures of the Pleistocene
stratigraphy in east-central Iowa (Figs.
12 and 13). Most studies of Iowa
Figure 8. Photomicrograph of the contact between the Chapin oolite Pleistocene stratigraphy are conducted
and Maynes Creek dolomite in a vertical boring where Maynes Creek using cores from bore-holes because of
has infilled into the Chapin. Note the truncation of oolite grains. Up the lack of natural exposure of these
is to the left, bar scale = 1 mm (from Lawler, 1981).
strata and rapid erosion when they are
exposed. Stripping of overburden at
quarries offers a unique and relatively rare
As you look around the quarry, note
perspective. The exposure at Wendling’s
the undulatory erosional surface through the
Montour Quarry tells a story of how glacial
Eagle City beds. Glacial striations are visible
and interglacial periods affect sediment
on this erosional surface where the
deposition, soil formation, erosion, and
overburden has been removed. While the
burial.
upper contact of the Eagle City has been
removed at Montour Quarry, it is preserved
After parking on 290th Street, walk
just a few miles away at LeGrand Quarry
down the slope (along the east edge of the
where the overlying Iowa Falls dolomite is
bean field) towards the quarry. Once you
present in places (B. D. Gossman and M. R.
have descended into the quarry you will be
Dawson, pers. comm., 2005).
walking on Eagle City dolomite. Examine
the upper surface for Pre-Illinoian striations.
These striations were formed when rocks
entrained in basal ice were dragged across
the bedrock surface, and can be used to
interpret local ice-flow direction. Because
of erosion in the intervening time between
Pre-Illinoian glaciations and the present,
striations of this age can only be observed in
settings where overburden has preserved
them.
The Eagle City dolomitic limestone at
Figure 9. Ripple marks at the top of the Maynes Creek ledge.
63
Directly overlying the striated
Mississippian bedrock is approximately 30
feet of Pre-Illinoian till (labeled 5 in Figs. 12
and 13). In Iowa, Pre-Illinoian tills range in
age from at least 1.6 Ma up to the onset of
Alburnett Formations on the basis of clay
content and type. Work by Hallberg (1980)
has shown that the Wolf Creek Formation is
dominated by expanding clays with minor
amounts of kaolinite and illite. The Wolf
Creek Formation is divided into three
separate till units based upon lithology and
associated paleosols. The type section for
the Wolf Creek is located in northwest Tama
County within 20 miles of the Montour
Quarry. The Pre-Illinoian till here is most
likely the Hickory Hills Member of the Wolf
Creek Formation, the upper-most till of the
eastern Iowa Pre-Illinoian succession.
Figure 10. Examples of points made from Maynes Creek chert.
The Table Rock point (left) was found in Dallas County, IA and
is 3000-5000 years old. The Scottsbluff point was made in 2000
by Professor John Whittaker of Grinnell College.
the interglacial preceding Illinoian glaciation,
pre-250 ka (Hallberg and Boellstorff, 1978).
Till of this age is extremely difficult to date
given that it is radiocarbon “dead” (>80 ka).
Figure 12. Annotated photograph of the Pleistocene section along
the northwest wall of Montour Quarry. 1 = modern Fayette and
Downs soils; 2 = interbedded late Wisconsinan loess and sand;
3 = late Sangamon paleosol; 4 = Yarmouth-Sangamon paleosol;
5 = pre-Illinoian till.
At least four tills of Pre-Illinoian age
are generally recognized in Iowa although
there are likely many more (Hallberg, 1980).
These are split into the Wolf Creek and
Examination of the Pre-Illinoian till
at Montour Quarry reveals a largely finegrained diamicton. Clay and silt predominate
with pebble to boulder size particles present
randomly throughout the till.
Clast
composition within the Wolf Creek is often
bimodal. Locally derived carbonate clasts
can sometimes be found near the base of the
till and reflect erosion of Paleozoic bedrock
in the immediate area of the quarry. Fartraveled clasts of igneous/metamorphic rocks
are also present and indicate a northern
source area. Many of the crystalline clasts
show extreme weathering and have decayed
Figure 11. Rip-up clasts near the top of the Eagle City dolomite.
64
1A
A horizon of Fayette-Downs soils; eroded in some
areas.
1B
Bt horizon of Fayette-Downs soils; Significant
accumulation of clay, angular, blocky structure.
1C
BC horizon of Fayette-Downs soils; transition between
parent material and mature soil.
2A
Fine to coarse sand; oxidized sand bodies, fairly
continuous with occasional cross bedding.
2B
Interbedded loess and sand; oxidatized loess and sand,
convoluted interbeds, some clay present.
3
Late Sangamon Paleosol-Pisgah Loess; highly
developed argillic horizon in LSP, Pisgah loess
occasionally present in some locations, preserved wood
fragments.
4
Yarmouth-Sangamon Paleosol; extremely well
developed argillic horizon, stone line near upper
boundary, sharp erosional contact with overlying LSP.
5
Wolf Creek Formation; Pre-Illinoian till, argillic
horizon near top, far traveled clasts in upper, locally
derived clasts in lower.
Figure 13. Figure 13. Pleistocene stratigraphic section from the northwest wall of Montour Quarry.
65
Soil formation resumed during the
Sangamon Interglacial after the Illinoian and
before the Wisconsinan glaciations. During
this time the Late Sangamon Paleosol (LSP)
was developed atop the YSP (bed 3 in Figs.
12 and 13). Here, the LSP is three feet thick
and is easily visible as it is often inset into
the outcrop face with the YSP protruding
below and Wisconsinan Loess extending
above it. At some locations within the
quarry it is possible to see the LSP extending
down into eroded areas of the YSP and
cutting across YSP structures.
Figure 14. Boulder-sized erratic weathered to grus in pre-Illinoian
Wolf Creek Formation.
On the west wall of the quarry the
LSP is overlain by sediments of late
to grus as a result of hundreds of thousands
of years of groundwater action (Fig. 14).
Grading into the Pre-Illinoian till
from above is approximately five feet of
Yarmouth-Sangamon Paleosol (YSP) (unit 4
in Figs. 12 and 13). In Iowa, the YSP is
widespread in areas not affected directly by
ice during the Illinoian glaciation. This soil
began to form during the Yarmouth
interglacial following Pre-Illinoian glaciation
and preceding Illinoian glaciation. In Tama
County, the Yarmouth interglacial lasted at
least 200 ka and resulted in a soil profile with
a strongly developed Bt horizon containing
up to 56% clay (J. Sandor, pers. comm.,
2005). The lower contact between the YSP
and the Pre-Illinoian till is gradational since
the till is the parent material for the YSP.
Note the stone line near the top of the YSP
(Fig. 15).
This stone lag is present
throughout the Iowan surface (Prior, 1976)
and represent erosion of the YSP during the
Illinoian time. Unlike the lower contact, the
upper contact of the YSP (Fig. 16) is
extremely sharp (less than ½”) and is marked
by a thin zone of Fe oxides and organic
matter. The contact is immediately above the
stone lag.
Figure 15. Stone lag indicating erosion surface between YarmouthSangamon paleosol (YSP) and late Sangamon paleosol (LSP).
Figure 16. Sharp contact between Yarmouth-Sangamon paleosol
(below) and late Sangamon paleosol (above).
66
deposits in order to observe their often
beautiful interbedded patterns (Fig. 17). In
some locations the relationship between the
loess and the sand suggests that soft sediment
deformation occurred (Fig. 17, lower left of
trowel). These forms are often referred to as
“load casts” or “flame structures” and
indicate downslope mass wasting of the
interbedded material while still soft.
Often a thick (12”-18”) sand body is
present at the top of the aeolian deposits at
this site and may represent localized fluvial
activity. In some areas of the quarry a thin
layer of earlier Wisconsinan loess (Pisgah) is
present between the LSP and the late
Wisconsinan aeolian deposits. Wood in this
loess was dated by George Hallberg in 1978
to 24,500 ± 820 Y.B.P (unpublished data).
To view this loess you will likely have to
make a short excursion to the southern end of
the current exposure area within the stripped
overburden area.
Figure 17. Interbedded late Wisconsinan loess and sand.
Wisconsinan in age that are interpreted to be
aeolian (bed 2 in Figs. 12 and 13). These
sediments were likely produced by adiabatic
winds that made dunes of loess and sand
derived from the Iowa River just south of the
quarry. At this outcrop the sand is often
cross bedded and is coarser than would
commonly be expected of aeolian deposits,
possibly due to proximity of the river source.
During the late Wisconsinan, the Iowa River
would have been a braided outwash stream
overloaded with sediment. Use a trowel or
paint scraper to make a clean cut of these
Topping the exposure are modern
soils from Fayette to Downs. These soils
often have strongly developed argillic
horizons and mirror in many ways the
paleosols beneath them. Examine the Bt
horizon of the modern soils to see the blocky,
angular structure common when clay
accumulation is present. As with many
Iowa soils, the A horizon has been largely
removed by erosion. The parent material
(C horizon) for the modern soil is the
interbedded late Wisconsinan loess and
sand. The contact between the two (Fig.
18) is gradational and in some areas the
modern Bt/BC horizon can be seen grading
into sand.
ACKNOWLEDGEMENTS
Figure 18. Contact between modern Bt/BC horizon (above)
and sandy late Wisconsinan aeolian parent material.
67
We are grateful to Marc Whitman
of Wendling, Inc. for providing access to
Montour Quarry. Numerous discussions
with Brian Gossman and Robert Dawson of
the Iowa DOT improved our understanding
of the Mississippian nomenclature problems
in Iowa.
Laudon, L. R., and Beane, B. H., 1937, The
crinoid fauna of the Hampton Formation
at LeGrand, Iowa, University of Iowa
Studies in Natural History, 17, 225-273.
REFERENCES
Lawler, S. K., 1981, Stratigraphy and
petrology
of
the
Mississippian
(Kinderhookian) Chapin Limestone of
Iowa [M.S. thesis]: Iowa City, Iowa,
University of Iowa, 118 p.
Burggraf, G. K., 1981, Clarification of the
stratigraphic position of the Maynes
Creek Member of the Hampton formation
(Mississippian): Geological Society of
America Abstracts with Programs, 13,
273.
Morrow, T. A., 1982, Maynes Creek Chert:
A common lithic material from central
Iowa, Office of the State Archaeologist
Research Papers, Iowa City, Iowa, 7(2)
306-319.
Glenister,
B.F.
1987.
Mississippian
carbonates of the Le Grand area: ancient
analogs of the Bahama Banks, Geological
Society of Iowa, Guidebook 47.
Prior, J. C. 1976. A regional guide to Iowa
landforms:
Iowa Geological Survey
Education Series 3.
Gossman, B.D., 1985. Stratigraphic column
for Montour Quarry, NW ¼ Sec 9 T83
R16W Tama County, Iowa Dept. of
Transportation.
Thomas, L. A., 1960, Guidebook for the 24th
Annual Tri-State Geological Field
Conference, North-Central Iowa, 28 p.
Hallberg,
G.R.
1980.
Pleistocene
Stratigraphy in East-Central Iowa: Iowa
Geological
Survey
Technical
Communication Series, no. 10.
Van Tuyl, F.M. 1925. The stratigraphy of the
Mississippian formations of Iowa, Iowa
Geological Survey, Annual Report, 30,
33-359.
Hallberg, G.R., and Boellstorff, J.D., 1978,
Stratigraphic “confusion” in the region of
the type areas of Kansan and Nebraskan
deposits: Geological Society of America
Abstracts with Programs, 10(6), 255.
Witzke, B. J., 1990, Paleoclimate constraints
for Palaeozoic paleolatitudes of Laurentia
and Euramerica, In McKerrow, W.S. and
Scotese,
C.R.,
eds.,
Palaeozoic
palaeogeography and biogeography:
Geological Society of London Memoir
12, 57-73.
Harris, S.E. 1947. Subsurface stratigraphy of
the Kinderhook and Osage Series in
southeastern Iowa, [Ph.D. thesis]: Iowa
City, Iowa, University of Iowa, 155 p.
Witzke, B. J. and Bunker, B. J., 1996,
Relative sea-level changes during Middle
Ordovician
through
Mississippian
deposition in the Iowa area, North
American craton, in Witzke, B. J.,
Ludvigson, G. A., and Day. J., eds.,
Paleozoic Sequence Stratigraphy: Views
from the North American Craton:
Laudon, L. R., 1931, The stratigraphy of the
Kinderhook Series of Iowa, Iowa
Geological Survey, Annual Report, 35,
333-451.
68
geologic map of Iowa, Phase 6: EastCentral Iowa, Scale 1:250,000, Iowa
Geological Survey Open File Map 03-2.
Boulder, Colorado, Geological Society of
America Special Paper 306, 307-330.
Witzke, B. J. and Bunker, B. J., 2001,
Comments
on
the
Mississippian
Stratigraphic Succession in Iowa, in
Heckel, P.H., ed., Stratigraphy and
Biostratigraphy of the Mississippian
Subsystem (Carboniferous System) in its
Type Region, the Mississippi River
Valley of Illinois, Missouri, and Iowa,
International Union of Geological
Sciences
Subcommission
on
Carboniferous Stratigraphy Guidebook
for Field Conference, St. Louis, Missouri,
63-75.
Witzke, B. J., Anderson, R. R., Bunker, B. J.,
Ludvigson, G. A., and Greeney, S., 2001,
Bedrock geology of north-central Iowa,
Digital geologic map of Iowa, Phase 3:
North-Central Iowa, scale 1:250,000,
Iowa Geological Survey Open File Map
01-3.
Woodson, F. J. and Bunker, B. J., 1989,
Lithostratigraphic
framework
of
Kinderhookian and Early Osagean
(Mississippian) strata, north-central Iowa
in Woodson, F. J., 1989, An excursion to
the historic Gilmore City quarries,
Geological Society of Iowa, Guidebook
50, 3-17.
Witzke, B. J., Anderson, R. R., Bunker, B. J.,
and Ludvigson, G. A., 2003, Bedrock
geology of east-central Iowa, Digital
69
Appendix A: Maps
70
Location map of field trip Stops 1 to 5
71
Location map of field trip Stop 6
72
Appendix B: 2005 Tri-State Road Log
#
1
2
3
Mile
0.0
0.1
0.2
4
1.6
5
4.1
6
7
4.2
4.8
8
5.2
9
5.5
Location
Start east of Scheman building
Turn left onto Center Dr.
Turn right onto Elwood (going south).
Stop #
Note
Upon entering Hwy 30 and
going east, note the
Hunziker Sports Complex
on the right (south) in the
floodplain of the South
Skunk River. Two city
wells here account for
about 20 percent of Ames'
drinking water. There is a
USGS gauging station on
the north side of the road.
Turn left to enter ramp for Hwy 30 (going
east). Upon entering Hwy 30, note sports
Complex on the right (south) and its 2 city
wells.
Exit Hwy 30 and turn north onto overpass
(going north).
Turn left onto 16th Ave. (going east).
Turn left onto S. Dayton Ave. (going north).
Stop at sign for Shady Grove Trailer Park
and turn right (going east)
Follow road to the end and turn left
Stop 1:
Whatoff's
Pit
10
5.7
Return to Shady Grove sign on Dayton and
turn right (going north)
11
7.2
Turn left onto 13th St. (going west).
12
8.7
Turn left on Duff (going south)
After descending into the
South Skunk River
floodplain, note River
Valley Park on the right.
This is the home of the lowhead dam that will be
discussed at Stop 2.
This is the main part of the
Downtown well field. All
wells in this area are
finished below ground
surface and draw water
from the buried channel
aquifer.
13
9.1
Turn right on 7th St. (going west).
14
9.6
Turn right onto Grand Ave. (going north)
15
11.5
Turn left onto Bloomington Rd. (going
west).
73
The Skunk River floodplain
is to the right. We are now
within the watershed for
Ada Hayden Heritage Park.
16
12.4
Turn right on Eisenhower by Stonebrook
church (going north)
17
12.9
Park on Harrison St. – walk to parking lot in
Ada Hayden Heritage Park.
Now entering the
Bloomington Heights
subdivision. Stormwater
from this area discharges
into the treatment wetlands
prior to entering the South
lake.
Stop 2:
Ada
Hayden
Park/Lake
.
18
13.8
19
14.6
20
15.7
21
16.0
22
16.2
23
22.8
Return to Bloomington Rd. (going south on
Eisenhower ) and turn left (going east)
Turn left on Grand Ave./US 69 (going
north).
Turn left into the entrance of the park.
Stop at parking lot north of lake (restrooms
nearby).
Exit the park to Grand Ave. (US Hwy 69)
and turn left (going north).
Turn left onto E18 (also called Hwy 221 or
130th St., going west).
24
31.8
Turn right onto 17 (going north).
25
38.7
Turn left onto Hwy 175 (also called 360th
St., going west).
26
45.1
27
52.5
28
59.2
29
61.2
30
31
61.5
62.4
Turn right onto R21 (east of Stratford, also
called Stagecoach Rd., going north).
Turn left on D46 (also called 290th St.,
initially going west, passing Brushy Creek
State preserve), follow road curving down
south into Des Moines River valley and
crossing bridge.
Turn right (directly after bridge) onto (old)
Hwy 50 (town of Lehigh) going west.
snack stop
This intersection is very
near the crest of the
Altamont Moraine (see
Stop 4).
Note hummocky
topography of the Altamont
Moraine.
possible pitstop to the right
before bridge (Riverside
tavern)
Convenience store on the
right, directly after the
bridge
Turn right onto D33 (also called Quall St.,
going north).
Entrance to Dolliver State Park.
Park here for Dolliver tour.
Lunch stop with benches to
Stop 3:
the left about 1 mile after
Dolliver
State Park the official entrance
Note: a tour inside the Park follows (9 pages) after which the road log continues
74
Parking lot is on alluvium covered bench
over bedrock. Trail to Copperas Beds is
along Prairie Creek. Start of trail is on a
dissected alluvial fan of Prairie Creek when it
flowed at the level of the parking lot terrace.
Dolliver Memorial State Park tour
(Park entrance – Miles counted from here)
This is an area of about 600 acres located on
the Des Moines River and is a memorial to
Jonathan P. Dolliver; orator, statesman, and
conservationists. Mr. Dolliver served in the
US House of Representatives from 1899 –
1900 and the US Senate from 1900 – 1910.
He formerly practiced law in nearby Fort
Dodge.
Stop 3A – Copperas Beds, Dolliver Park
(SW ¼, SE ¼, SE ¼, Sec. 34, T88N, R28W;
and NW ¼, NE ¼, Sec. 3, T87N, R28W).
From the north end of the parking lot, follow
the trail to the west along Prairie Creek to the
wooden foot bridge crossing the creek. (At
this point, a short introduction will be given.)
Of interest in the park are the deep ravines
with massive cliffs of Pennsylvanian
sandstone and wooded ridges and valleys
covered with almost every variety of tree,
shrub, wild flower and fern native to central
Iowa. The Pennsylvanian strata of central
Iowa are very poorly exposed. There are,
however, a few outcrops of the Lower Des
Moinesian Cherokee Group exposed along
the Des Moines River. One of these is in
Dolliver Park. The exposures here and in the
near vicinity were studied by Burggraf,
White and Lindsay in 1981 and delineated
into
six
lithofacies
representing
subenvironments of fluvially-dominated,
high constructive deltaic systems. The best
exposed is the lenticular fine to medium
grained, cross-stratified sandstone facies
which was interpreted to represent
distributary channel deposits.
We will
examine and discuss 3 outcrops of this facies
in Dolliver Park.
These sand bodies
correspond to the channels labeled #4 in
figure 8 of the attached article entitled
“Introduction and Regional Geology” by
Lemish, Chamberlain and Mason (1981).
The following comments, descriptions and
illustrations of outcrops are taken verbatim
from the road log published by Burggraf,
White, Palmquist and Lemish (1981).
Proceed south, along the eastern bank of
Prairie Creek, to view the sandstone outcrop
comprising the Copperas Beds. The section
(figure 2) includes a series of thin to very
thick beds of very fine to conglomeratic
sandstones with abundant carbonized, and in
some cases permineralized, branches and
leaves.
Fragmented carbonized and/or
pyritized woody material is very common in
the lower, coarser-grained deposits which
also include abundant subangular to wellrounded discoidal rip-up clasts of fine sandy
siltstone. These are usually armored by a 1/8
to 1/4 inch (3-6 mm) thick rind of iron oxide
minerals which causes the clasts to stand out
from the outcrop face and often to pluck out
as a single piece leaving only a partial mold
behind. The clasts are rarely imbricated and
usually very poorly sorted. They occur in
thick lenses or wedges with sharp basal
surfaces and sharp to gradational upper
contacts; in many cases these bodies are
interpreted to represent the cores of channel
bars with laterally gradational and
interfingering deposits of cross-stratified
fine- to medium-grained bar-side sands. The
strata of the Copperas Beds represent several
cycles of aggradation within a channel and
include vertically stacked and interfingering
bar forms and overbank fine-grained
carbonaceous sand and silt. Of special
interest at this locality is the occurrence of
Mile 1.2
Turn left (west into picnic grounds at
Copperas Beds. Park in lot at end of road.
75
Figure 1. Field trip stops, Dolliver Park.
sulfate efflorescences along the outcrop face.
These were originally identified as ferrous
sulfate, or copperas, from which the beds
derive their name. Later, analyses by the
Iowa Geological Survey identified the
hydrated
ferrous
sulfate
melanterite
76
grained, is interpreted to represent delta front
sedimentation. The lenticular sand bodies,
horizontally bedded and cross-stratified to
ripple-bedded are believed to represent a
crevasse channel while the thin-sheet
sandstone represents a crevasse splay.
Overlying the delta front interval is a thick
sand sequence, the lower half of which is
well exposed in the lower part of the cliff.
Notice the sharply erosional basal contact
with abundant rip-up clasts and carbonaceous
debris, the lenticular body geometry with
thinning to the north, and the abundant
primary structures exposed in the cliff face.
This locality includes abundant large-scale
planar cross-bed sets to greater than 2 feet
(0.6 m) thick which typically thicken in a
down stream direction. Each set includes an
erosional bse and is overlain either by
shallowly dipping stoss-side cross-strata.
Throughout Dolliver Park cross-beds are
overwhelmingly directed to the southwest
with a maximum of readings (143 or 249 or
57%) trending between S30°W and S70°W.
Individual beds commonly exhibit normal
grading on outcrops which are strongly ironstained but poorly indurated. To the north, a
wooden footbridge crosses the creek. If the
creek level is low enough, walk across the
bridge and along the outcrop to observe
details of bedding.
(FeSO4·7H2O). More recent work by Cody
and Biggs (1973) of Iowa State University
has shown that the efflorescences, consisting
of a layer (3/4 inch thick) of white, fibrous
crystals intermixed with equant very finegrained crystals, include at least 3 mineral
species:
halo-rozenite (FeAL2(SO4)4·22
H2O), szomolnokite (FeSO4·H2O), and
rozenite (FeSO4·4H2O). All are readily
soluble in water and during heavy rainstorms
may be completely cleared from the outcrop
face. Subsequent periods of low to moderate
humidity result in renewed precipitation of
the sulfate rind.
Return to the bus.
1.8
Ford Prairie Creek
2.0
Outcrop of channel sandstone to left (north);
notice large concretion with mammalary
morphology.
2.1
Stop 3B – Copperas Beds, Dolliver Park
(SW ¼, NE ¼, SW ¼, Sec. 35, T88N,
R28W). From the bus, cross the sand and
proceed east-ward to the outcrop exposed
along the east side of Prairie Creek. Stop
before crossing the creek.
Return to the bus.
This exposure (Fig. 3) provides an
opportunity to see deposits of the distributary
channel
and
floodbasin
depositional
environments.
At the very base are
arenaceous siltstone and laminated silty
claystones with small lentils of very finegrained sandstone and thin sheet sandstone.
This interval, carbon-rich and very fine-
2.3
Oxbow lke to right.
2.8
Structural terrace to left with camp ground.
(text continues after Fig. 2)
77
(5Y6/4); fine-grained; quartz with
silt galls, micaceous, clay mineral
cement; basal contact sharp, soured;
moderately friable; up to 18 ft.
Unit Description
J.
Sandstone, silty; grayish orange
(10YR7/4); very fine grained;
indistinct bedding; very friable, 7 ft.
I.
Claystone, silty; light gray (N6);
laminated; friable, 2 ft.
H.
Sandstone’ similar to unit
occasional claystone band; 5 ft.
G.
Alternating
sandstone
and
claystone’ similar to unit E; 6 ft.
F.
Sandstone, same as unit D; 5 ft.
B.
Conglomerate,
intraformational;
dark reddish brown (10R3/4) to
light brown (5YR5/6); very poorly
sorted, pebble to boulder size,
moderate sphericity and roundness,
quartz and siltstone clasts cemented
by ferroan dolomite; basal contact
sharp and erosional; thin-bedded
with slight imbrication; very well
indurated, cliff former; up to 20 ft.
A.
Conglomerate – sandstone: cgl;
quartz pebble and sandstone clasts;
very dusky red (10R2/2) to
moderate reddish brown (10R4/6);
pebbles well rounded, poorly
sorted; basal contact not exposed;
thin bedded; very well indurated
with dolomitic cement; sandstone’
light brown (5YR5/6); quartzose;
fine
grained;
small
scale
crossbedding; friable to moderately
indurated; unit up to 10 ft.
C;
E.
Sandstone – siltstone’ sandstone’
same as unit D; siltstone;
argillaceous and carbonaceous with
clay galls; friable’ 8 ft.
D.
Sandstone, similar to unit C; some
carbonaceous debris; horizontal
laminations to trough crossbedding;
lensatic, cobble-clast band at top;
10 ft.
C.
Sandstone, subarkose with thin
claystone beds; dusky yellow
78
Figure 2. Graphic Section of Dolliver Park – Copperas Beds.
79
Approximately 15 feed (4.5 m) above the
creek level at outcrops B and C deformed
cross-strata are exposed. These deformation
features are discussed earlier in the text
(figure 15) and have been referred to as
intraformational recumbent folding (Reineck
and Singh, 1975) or recumbent-folded
deformed crossbedding (Allen and Banks,
1972). The over-turning of the foreset
laminae and other soft sediment deformation
features suggests a saturated condition for the
sand beds shortly after deposition and may
reflect seismic activity which triggered the
deformation.
Continue along the creek
noting the abundant cross-stratification, both
undeformed and deformed. Leave the creek
bed, climb out along the north side of the
valley to a picnic area and the bus.
3.4
Stop 3C – Boneyard Hollow, Dolliver Park
(Park; NW ¼, NW ¼, NE ¼, Sec. 35, T88N,
R28W). From the bus proceed across the
road along the north bank of the creek
draining Boneyard Hollow
(figure 4).
Looking to the south-southwest, notice the
sandstone outcrop which comprises the south
bank of the drainage way. This distributary
channel sand overlies light gray mudstone
which is poorly exposed at the cliff base.
Notice also the recent rock fall and the
abundance of cross-stratification marked by
iron-oxide stained foreset laminae. The
sandstone/mudstone
contact,
though
indistinctly exposed, is marked by springs
which drain the overlying permeable sands.
Continuing to the northwest along the trail,
cross the creek and observe the vertical cliff
faces exposed adjacent to the creek.
(Figs. 3 and 4 follow):
80
Figure 3. Graphic section of Dolliver Park.
sand; same colors as unit B;
carbonaceous and micaceous;
clasts up to 6 in dia.; sandstone
clast conglomerate in upper
portion;
indistinct
bedding;
moderately friable; 20 ft.
Unit Description
I.
Sandstone-siltstone
repetitions;
sandstones, calcareously cemented;
similar to unit D below; 23 ft.
H.
Siltstone; sandy at base but very
poorly exposed; covered slope; 36
ft.
G.
Sandstone, subarkose, similar to
unit E; 9 ft.
F.
Alternating
layers
of
sandstone and siltstone-claystone;
as in Unit D; 9 ft.
E.
Sandstone’ same colors as unit B;
fine-grained; indistinct bedding
with scour surface and clay galls in
lower portion; basal contact sharp;
friable; 16 ft.
D.
Sandstone-siltstone;
sandstone,
carbonaceous
with
wood
fragments; yellowish gray (5Y7/2),
fine-grained, horizontal lamination;
friable; 4 ft. thick; siltstone,
argillaceous to arenaceous and
carbonaceous; grayish orange
(10YR7/4), to light gray (N7);
laminated; friable; 1.5 ft. thick;
total thickness 5.5 ft.
C.
Sandstone, abundant claystone
bands and rip-up clasts of silt and
81
B.
Sandstone;
subarkose;
dark
yellowish
orange
(10YR6/6),
weathers
moderate
brown
(5YR4/4); fine-grained, moderate
sphericity and roundness, well
sorted; quartzose with minor
feldspar, siltstone rip-up clasts at
base; basal contact sharp, irregular
with isolated coaly pods; small to
large scale trough and planar
crossbedding in sets to 2.5 ft. thick;
well indurated, cliff former; thins
laterally; 22 ft.
A.
Siltstone-sandstone’
siltstone,
argillaceous and carbonaceous;
medium dark gray (N4) to grayish
orange (10YR7/4) at the top;
thickly laminated to laminated;
friable;
sandstone;
moderate
yellowish brown (10R5/4) to
moderate brown (5YR5/5); very
fine-grained, well sorted; basal
contacts sharp; lenticular; faint
cross-bedding, ripple marks on
upper surfaces; moderate to well
indurated; lenses to 3 ft. thick;
basal contact of unit A not
exposed; up to 12 ft.
82
Figure 4. Generalized graphic section of Boneyard Hollow.
Unit Description
B.
Sandstone; pale yellowish orange
(10YR8/6) to pale brown (5YR5/2);
fine- to medium-grained; abundant
large-scale planar cross-strata;
strongly iron-stained in basal
portion; upper contact covered;
basal contact sharp, erosional;
greater than 35 feet.
A.
(Road log continues on the next page)
83
Sandy mudstone to claystone;
medium gray (N5) to dark
yellowish orange (10YR6/6); lower
contact covered; basal 4 feet is
laminated claystone; gradational
upward to sandy mudstone 7 feet
thick; highly iron-stained; 11 feet
exposed.
# Mile
32 64.6
33 66.3
68.5
34 72.1
35 81.1
37 84.1
38 88.0
Location
Return to Park entrance (going south)
From official entrance follow D33 south
Turn left onto Hwy 50 (also 290th St.) going back east
towards Lehigh
Follow Hwy 50 into Lehigh - do not go over bridge(!) but
keep going and turn left where road becomes P73, keep
going south after leaving Lehigh (then also called Samson
St.)
Turn left onto D54 (also called 330th St., first going east
then turning south and running west of Stratford)
Turn left onto Hwy 175 (also called 360th St., going east)
Turn right onto R27 (also called Fenton, going south)
39 91.0
Turn left onto 400th St. (gravel road, going east).
40 93.8
Stop at Bjorkboda Marsh .
Stop #
Note
Convenience
store on the
left
Stratford
Note
hummocky
topography
of the
Altamont
Moraine.
Stop 4:
Bjorkboda
Marsh
41 96.6
42 105.6
43 106.4
44 107.1
45 107.9
Going west on 400th St. return to R27 and turn left going
south.
Turn right onto E26 (W 22 St.), going west.
At stop sign go straight and stay on E26 (also called W 22
St. or Monarch Dr.).
Curving left, follow E26 (W 22 St.) now going south.
Cross rail road (still going south).
After 1/4
mile: pitstop
at gas
station/bar
(right)
46 108.3 Turn right onto W Mamie Eisenhower Rd.
47 109.0 Turn left and stay on E41 (216th Dr.).
48 111.0
Abandoned
gravel pit to
the south is
cut into a
lateWisconsinan
river terrace.
Turn right onto Laurel Lane for Rose Hill Cemetery (stay
right, up the hill).
49 111.2 Stop at cemetery.
Stop 5: Des
Moines
River Valley
93.8
From cemetery go south back to 216th Dr. and turn
right(!) (going west).
51 112.0 Turn left to R18 (also called L Ave., going south).
50 111.4
84
52
53
54
55
56
57
112.2
115.8
128.5
130.1
130.3
130.5
58 130.5
59 131.9
60 134.2
61 156.1
62 179.1
63 179.7
64 180.4
65 180.6
66 181.2
Turn left onto Hwy 30 (going east towards Ames).
Stop sign.
Exit Hwy 30 and turn left onto Elwood (going north).
Turn left onto Center Dr.(going west).
Turn right and follow to Scheman Bldg. (going north).
Stop at Scheman Bldg.
Turn right (south) on to Elwood Drive from the Iowa
State Center.
Turn left on to U. S. Highway 30 (eastbound).
South Skunk River.
Drive off the east edge of the Des Moines Lobe at
State Center.
Turn left on to C Avenue. As of press time, this was
an unmarked intersection in a construction zone.
Descend into the Iowa River floodplain.
Iowa River.
Entrance to Montour Quarry.
Turn left on to 290th Street.
Park in the private road on the south side of 290th
Street or along the side of 290th Street. Walk south
along the edge of the bean field to reach the quarry.
85
Stop 6:
Montour
Quarry