EARLY CRETACEOUS TO PALEOCENE PALEOGEOGRAPHY OF THE WESTERN INTERIOR SEAWAY: THE INTERACTION OF EUSTASY AND TECTONISM

Abstract

To address the limited temporal and spatial context of earlier paleogeographic studies of the Western Interior Seaway (WIS), we provide a broad overview of the paleogeographical history of this epeiric sea and the basin it filled by synthesizing available data from detailed analyses of lithostratigraphic sections, isopach maps, as well as biostratigraphic and biogeographic distributions. The Early Cretaceous to Paleocene WIS connected the Arctic Ocean with the Gulf of Mexico and was likely episodically linked to the Atlantic Ocean via the Hudson Seaway. The paleogeography of the WIS primarily reflects the interplay between sea level and physiography, which were controlled by eustasy, tectonics, and pre-existing topographic features. The Western Interior Foreland Basin (WIFB) initiated in the Middle Jurassic with the subduction of Panthalassan crust beneath the North American plate which resulted in a continental margin arc-trench system and the Cordilleran Orogenic Belt. The deformation front of the Cordillera migrated progressively eastward forcing a cratonward (i.e., eastward) shift in marine sedimentation throughout its history. Several punctuated Albian marine transgressions in the WIFB were restricted to the northern part of the basin due to low sea level compounded by its physiography. The first marine connection of the WIS linking the Arctic Ocean and the Gulf of Mexico occurred during the late Albian as a response to a eustatic rise, together with increased basin subsidence. The sea's north-south connectivity was lost during lower sea levels, but, once established in the middle Cenomanian, persisted at least until the early Maastrichtian and possibly into the Paleocene. The onset of the Laramide Orogeny that broke the WIFB into a series of smaller intermountane basins along with a drop in sea-level caused the final withdrawal of the WIS from North America during the Paleocene.

Full text

22 WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
EARLY CRETACEOUS TO PALEOCENE PALEOGEOGRAPHY OF THE WESTERN
INTERIOR SEAWAY: THE INTERACTION OF EUSTASY AND TECTONISM
Joshua S. Slattery1, William A. Cobban2, Kevin C. McKinney2, Peter J. Harries1, and Ashley L. Sandness3
1 School of Geosciences, University of South Florida, 4202 East Fowler Ave., NES 107, Tampa, FL 33620 2 U.S. Geological
Survey, MS 975, Denver Federal Center, Denver, CO 80225, 3 Tampa, FL 33613
ABSTRACT
To address the limited temporal and spatial context of earlier paleogeographic studies of the
Western Interior Seaway (WIS), we provide a broad overview of the paleogeographical history of
this epeiric sea and the basin it filled by synthesizing available data from detailed analyses of
lithostratigraphic sections, isopach maps, as well as biostratigraphic and biogeographic
distributions. The Early Cretaceous to Paleocene WIS connected the Arctic Ocean with the Gulf of
Mexico and was likely episodically linked to the Atlantic Ocean via the Hudson Seaway. The
paleogeography of the WIS primarily reflects the interplay between sea level and physiography,
which were controlled by eustasy, tectonics, and pre-existing topographic features. The Western
Interior Foreland Basin (WIFB) initiated in the Middle Jurassic with the subduction of
Panthalassan crust beneath the North American plate which resulted in a continental margin arc-
trench system and the Cordilleran Orogenic Belt. The deformation front of the Cordillera migrated
progressively eastward forcing a cratonward (i.e., eastward) shift in marine sedimentation
throughout its history. Several punctuated Albian marine transgressions in the WIFB were
restricted to the northern part of the basin due to low sea level compounded by its physiography.
The first marine connection of the WIS linking the Arctic Ocean and the Gulf of Mexico occurred
during the late Albian as a response to a eustatic rise, together with increased basin subsidence.
The sea's north-south connectivity was lost during lower sea levels, but, once established in the
middle Cenomanian, persisted at least until the early Maastrichtian and possibly into the
Paleocene. The onset of the Laramide Orogeny that broke the WIFB into a series of smaller
intermountane basins along with a drop in sea-level caused the final withdrawal of the WIS from
North America during the Paleocene.
INTRODUCTION
In the west-central portion of North America,
Lower Cretaceous through middle Paleocene strata
representing sedimentation in the Western Interior
Foreland Basin (WIFB) form a thick and complex
mosaic of interfingering marine and terrestrial
deposits recording successive transgressions and
regressions of the Western Interior Seaway (WIS).
Throughout most of its history, this epeiric sea (i.e.,
a sea covering the interior of a continent) connected
the Arctic Ocean (or Boreal Ocean) with the Gulf of
Mexico (or Tethys Sea) and was joined to the North
Atlantic Ocean, possibly episodically, via the
Hudson Seaway (Williams and Stelck, 1975;
Kauffman and Caldwell, 1993; Roberts and
Kirschbaum, 1995; Ziegeler and Rowley, 1996).
Since the publication of Schuchert’s (1910)
“Paleogeography of North America”, numerous
studies have dealt with aspects of the WIS’s
paleogeography, but most of these have only focused
on relatively specific time intervals and restricted
geographical areas (e.g., Schuchert, 1955; Reeside,
1957; Sloss et al., 1960; Birkelund, 1965; Sohl,
1967; Jeletzky, 1971; McGookey, 1972; Gill and
Cobban, 1973; Williams and Stelck, 1975; Witzke et
al., 1983; Cobban and Hook, 1984; Kauffman, 1984;

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WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
Lillegraven and Ostresh, 1991; Dixon, 1993;
Kauffman and Caldwell, 1993; Stott et al., 1993;
Cobban et al., 1994; Sageman and Arthur, 1994;
Ziegeler and Rowley, 1996; Kennedy et al., 1998;
Erickson, 1999; Roberts and Kirschbaum, 1995;
White et al., 2000; Stelck et al., 2007; Miall et al.,
2008; Nielsen et al., 2008; Boyd and Lillegraven,
2011; Landman et al., 2012; Blakey, 2013; Schröder-
Adams, 2014). The most detailed studies of the
seaway’s paleogeography and its transgressive/
regressive successions have been primarily
concentrated along its western margin in an area
spanning from Alberta south to west Texas where
strata of this age are thick, well-exposed, and
relatively biostratigraphically complete (Reeside,
1957; Sloss et al., 1960; McGookey, 1972; Gill and
Cobban, 1973; Williams and Stelck, 1975; Cobban
and Hook, 1984; Kauffman, 1984; Lillegraven and
Ostresh, 1991; Stott et al., 1993; Cobban, 1994;
Sageman and Arthur, 1994; Roberts and
Kirschbaum, 1995; Boyd and Lillegraven, 2011;
Landman et al., 2012). Relatively fewer studies have
attempted detailed reconstructions in its northern and
eastern parts because of substantial post-Mesozoic
erosion and/or rarity of exposures due to greater
cover (e.g., Jeletzky, 1971; Williams and Stelck,
1975; Witzke et al., 1983; Cobban et al., 1994;
Roberts and Kirschbaum, 1995; Ziegeler and
Rowley, 1996; Erickson, 1999; Stelck et al., 2007;
Nielsen et al., 2008; Schröder-Adams, 2014). Similar
problems hamper our understanding of the WIS’s
paleogeography throughout its final regressions and
transgressions during the late Maastrichtian to
middle Paleocene (e.g., Witzke et al., 1983;
Erickson, 1999; Boyd and Lillegraven, 2011). The
combination of these factors and paucity of studies
synthesizing paleogeographic knowledge have
resulted in an incomplete understanding of the
spatial evolution of the WIS and the WIFB.
In this paper, we provide a broad overview of the
Early Cretaceous to Paleocene paleogeographic
evolution of the WIS and the WIFB that it flooded.
This summary will also delve into the geological
setting, the methodologies used for paleogeographic
reconstruction, as well as the various eustatic
controls that modulated sea level and the location of
the seaway’s shorelines. A detailed documentation of
western North America’s physiography during this
interval is necessary to understand how the WIS
formed, transformed, and eventually ceased to exist.
This overall consideration of the seaway’s
paleogeographic evolution should provide insights
for a better understanding of the stratigraphic
architecture of Lower Cretaceous through Paleocene
strata, assist in elucidating the distinctive
sedimentation patterns in the WIFB, and provide a
framework against which to compare paleontologic
trends and patterns.
PHYSIOGRAPHIC AND GEOLOGICAL
HISTORY OF THE WESTERN INTERIOR
FORELAND BASIN
The WIS’s paleogeography primarily reflects the
interplay between sea level and the physiography of
North America, which were both influenced to
varying degrees by tectonics (Gill and Cobban, 1973;
Williams and Stelck, 1975; Jeletzky, 1980;
Lillegraven and Ostresh, 1991; Krystinik and
DeJarnett, 1995; Kauffman and Caldwell, 1993;
Miall et al., 2008). Tectonic forces shaped western
North America’s physiography during the Middle
Jurassic to the Eocene through various processes,
including volcanism, uplift, subsidence, and fold-
thrust-belt migration. These processes, in turn,
heavily influenced shoreline patterns and lithofacies
distributions across North America.
The most prominent physiographic feature of
western North America during this time interval, as
well as subsequently, has been and remains the
Cordilleran Orogenic Belt, which spans nearly 6,000
km from Alaska to southern Mexico, reaching its
maximum width of over 1000 km in the western
coterminous United States and southwestern Canada
(Figure 1; Monger, 1993; DeCelles, 2004). This
mountain belt and its associated basins (including the
WIFB) comprise approximately a seventh of the
40,000 km Circum-Pacific Orogenic Belt (more
colloquially the ‘Ring of Fire’) that fringes the
Pacific Ocean (DeCelles, 2004). Since the Middle
Jurassic, this geographic feature has acted as a
sig nific a nt oceano grap h ic, climat ic, and
biogeographic barrier between the Panthalassan or
Pacific Ocean and the Western Interior (Williams
and Stelck, 1975). Prior to the uplift of the
Cordilleran Orogenic Belt, most Mesozoic marine
transgressions the flooded the Western Interior came
from the Pacific Ocean (Williams and Stelck, 1975).
During the Late Triassic to Middle Jurassic,

24 WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
Figure 1. Tectonic map of North America showing key geological and physiographic features of the Cordilleran
Orogenic Belt, Western Interior Foreland Basin, Hudson Bay Basin, North American Craton, Paleozoic orogenic belts,
and Gulf and Atlantic Coastal plains (modified from Bally et al., 1989; DeCelles, 2004).

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WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
prior to the development of the Cordilleran Orogenic
Belt and the WIFB, western North America was the
site of numerous fringing arcs and interarc basins
(Harper and Wright, 1984; Wright and Fahan, 1988;
Saleeby and Busby-Spera, 1992; Dickinson et al.,
1996; DeCelles, 2004). The precursor events that
eventually resulted in the Cordilleran Orogenic Belt
and WIFB’s formation initiated with the westward
drift of North America relative to Europe (Monger,
1993; DeCelles, 2004; Miall et al., 2008). This
westward drift resulted in the subduction of
Panthalassan oceanic crust beneath the North
America Plate accompanied by regional shortening
and thickening of the continental crust (Monger,
1993; DeCelles, 2004). This movement resulted in a
coherent continental margin arc-trench tectonic
system (Monger, 1993; DeCelles, 2004) that, at least
in the United States, was tectonically very similar to
the modern continental margin arc-trench tectonic
system of western South America (DeCelles, 2004).
Like the modern Andes Mountains and Amazon
Basin in South America, this region included a
forearc system (including a forearc basin; Figure 2),
a magmatic arc (Figure 2), and a retroarc (Figure 2).
The retroarc included both the Cordilleran Fold-
Thrust Belt (Figure 2) and resulting foreland-basin
system (or WIFB; Figures 1, 2; Monger, 1993;
DeCelles and Giles, 1996; DeCelles, 2004) with the
stable North American craton situated to the east
(Figures 1, 2; Kauffman and Caldwell, 1993;
DeCelles, 2004; Miall et al., 2008). The spatial
position and extent of these various tectonic
subdivisions changed through time both
perpendicular and parallel to the Cordillera’s strike
(Monger, 1993; DeCelles, 2004; Miall et al., 2008).
These changes played a critical role in the evolution
of the WIFB and consequently the paleogeography
of the WIS. Broadly, within the Cordilleran
Orogenic Belt, different tectonic processes were
expressed diachronously from south to north, which
were reflected in the physiographic changes that
occurred from the Middle Jurassic to Eocene. In
Canada, the Cordilleran Orogenic Belt and WIFB
arose during the Middle Jurassic from the collision
of terranes that could not be subducted beneath the
North American plate (Coney et al., 1980; Ricketts,
2008; Fuentes et al., 2009, 2011). Each new terrane
collision, starting in the Jurassic and continuing into
the Eocene, increased tectonic loading on the
Canadian Shield and initiated a new cycle of fold-
thrust belt tectonism, uplift, and the formation of
clastic wedges in the WIFB (Stockmal et al., 1992;
Miall et al., 2008). The Cordilleran Orogenic Belt
and WIFB in the United States were initiated in the
Middle Jurassic due to the accretion of the fringing
arcs associated with the closure of marginal oceanic
basins (e.g., the Mezcalera plate) along with the start
of Panthalassan oceanic crust subduction (DeCelles
and Currie, 1996; DeCelles, 2004; Miall et al.,
2008).
The forearc region along the western margin of
North America was mainly composed of Mesozoic
Figure 2. Generalized tectonic, structural, and stratigraphic cross-section across the Cordilleran Orogenic Belt and
Western Interior Foreland Basin System in North America during the Late Cretaceous (modified from DeCelles and
Giles, 1996; Miall et al., 2008).

26 WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
accretionary terranes and thick deposits of arc- and
oceanic-slab-derived sediments deposited into
forearc basins (Figures 1, 2; Monger, 1993;
DeCelles, 2004; Ingersoll, 2008; Ricketts, 2008). A
magmatic arc (or the western Cordillera) was
positioned eastward of the forearc and was the site of
active melting and related arc magmatism above the
eastward-dipping subduction zone (Figures 1, 2;
Monger, 1993; Ducea, 2001; DeCelles, 2004). The
volcanoes produced by this arc have long since
eroded away, although evidence for their former
presence exists in the form of calc-alkaline granitoid
intrusions preserved as batholiths in the Peninsular
Range of Baja California, the Sierra Nevadas of
California, the Idaho Batholith of Idaho and
Montana, and within the Coast Range of British
Columbia (Figure 2; Ducea, 2001; DeCelles, 2004).
Additional evidence for this volcanism include the
numerous widespread ash beds found throughout the
stratigraphic sequences of the WIFB (i.e., Utah,
Colorado, Wyoming, Montana, as well as North and
South Dakota), which have been used to
radiometrically date and correlate Upper Jurassic
through Eocene strata in this region (Elder, 1988;
Kauffman and Caldwell, 1993; Obradovich, 1993;
Kowalis et al., 1995; Cobban et al., 2006).
The retroarc, which included the Cordilleran Fold
-Thrust Belt (or eastern Cordillera) and the adjacent
WIFB, were situated in southern Nevada, Arizona,
Utah, Wyoming, Idaho, Montana, eastern British
Colombia, Alberta, Saskatchewan, eastern Yukon
Territory, and Northwest Territories (Figures 1, 2;
Kauffman and Caldwell, 1993; Monger, 1993;
DeCelles, 2004; Miall et al., 2008). During most of
its history, deformation in the eastern Cordillera was
characterized by numerous folds and thrust faults,
now obscured by later Neogene and Quaternary
erosion, igneous events, structural developments, and
metamorphic processes related to extension as well
as migration of the Yellowstone Hot Spot (DeCelles
and Mitra, 1995; DeCelles, 2004). Paleobotanical
and associated geological evidence suggests that the
region between the Cordilleran magmatic arc and
Cordilleran Fold-Thrust Belt was the site of a high
plateau (i.e., “Nevadaplano”) similar to the Altiplano
in the central Andes (Figure, 2; DeCelles, 2004;
Colgan and Henry, 2009; Ernst, 2009; Chamberlain
et al., 2012). The retroarc region was the primary
location of deformational shortening in western
North America during the Middle Jurassic to Eocene
(Monger, 1993; DeCelles, 2004).
Further east, tectonic loading by the Cordilleran
Fold-Thrust Belt depressed the crust (i.e., load-
induced flexural subsidence) and resulted in the
formation of the WIFB (Figures 1, 2; Kauffman and
Caldwell, 1993; DeCelles and Giles, 1996; Miall et
al., 2008). At its maximum extent, the WIFB was
over 1,500 km in width, extending from modern-day
central Arizona, central Utah, western Wyoming,
western Montana, and western Alberta to eastern
Kansas, Iowa, western Minnesota, and eastern
Manitoba (Cross, 1986; Kauffman and Caldwell,
1993; DeCelles and Giles, 1996; Miall et al., 2008).
The basin’s overall width is much greater than would
be expected based on a typical load-induced flexural-
subsidence basin model (e.g., Kauffman, 1977) as
WIFB subsidence was likely enhanced by subduction
-induced subsidence involving “viscous coupling
between the base of the continental plate and
downward circulating mantle-wedge that is entrained
by the subducting slab” (DeCelles and Giles, 1996,
p. 105; Figure 2). This subsidence, along with load-
induced flexural subsidence from sediment
accumulation and the Cordilleran Fold-thrust Belt,
were the primary mechanisms for creating
accommodation space in the WIFB (Kauffman and
Caldwell, 1993; DeCelles and Giles, 1996; Miall et
al., 2008).
The WIFB system consisted of four distinct
depositional zones (Figure 2), including from east to
west: 1) the wedge-top depozone, filled with
sediment that accumulated above the frontal part of
the orogenic wedge and thickens toward the
foredeep; 2) the foredeep depozone, a cratonward
tapering, thick accumulation of sediment located
between the tip of the deformation front and the
proximal flank of the forebulge; 3) the forebulge
depozone, a broad region of potential flexural uplift
between the foredeep and backbulge basin; and 4)
the backbulge depozone, a mass of sediment
collected in a shallow, but expansive zone of
possible flexural subsidence cratonward of the
forebulge (McMechan and Thompson, 1993;
DeCelles and Currie, 1996; DeCelles and Giles,
1996). With the addition of the wedge-top depozone
by DeCelles and Currie (1996) and DeCelles and
Giles (1996) as part of the WIFB, the basin’s
geometry becomes a doubly tapered prism instead of

27
WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
Figure 3. Map of structural features that were active during the Late Jurassic to Paleocene in the Western Interior
Foreland Basin (modified from Stott et al., 1993; Stelck et al., 2007; Miall et al., 2008). Abbreviations as indicated:
BLT, Blow Trough; KT, Keele Trough, PT, Peel Trough; KA, Keele Arch; CMA, Coppermine Arch; PA, Punnichy
Arch; CMU Central Montana Uplift; CMT, Central Montana Trough; CCA, Cedar Creek Anticline; BS, Beartooth Shelf;
BU, Bighorn Uplift; PRB; Powder River Basin; BHU, Black Hills Uplift; BT, Bighorn Trough; MCA, Miles City Arch;
CA; Casper Arch; ChA, Chadron Arch; GRT, Green River Trough; LU; Laramie Uplift; AB, Alliance Basin; UU,
Uncompahgre Uplift; US, Utah Shelf; CCT, Central Colorado Trough; FRU, Front Range Uplift; DB, Denver Basin; PB,
Paradox Basin; BMB, Bisbee-McCoy Basin; CHT, Chihuahua Trough; MB, Marathon Basin.

28 WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
the classically asymmetric wedge suggested by
earlier studies of the basin (e.g., Kauffman, 1977;
Kauffman and Caldwell, 1993).
Latitudinally, the WIFB contains a number of
different, relatively minor pre-existing tectonic
elements or basement features that are thought to
have exerted substantial influence over the region’s
physiography, the WIS’s paleogeography, and the
lithofacies patterns within the basin. These tectonic
elements consist of various arches, block-uplifts, and
sub-basins that were hypsographically expressed
throughout the basin’s history or became expressed
because of tectonic activity (Figure 3; Weimer, 1978;
Stott et al., 1993; Stelck et al., 2007; Miall et al.,
2008). Most of these structural features reflect
reactivated basement elements that produced various
topographic highs and lows that acted as local
sediment sources, sediment sinks, and barriers to sea
-level rise (Weimer, 1978; Stelck et al., 2007; Miall
et al., 2008). In the northern part of the WIFB,
transverse structural features or highs, such as the
Dave Lord-Eskimo Lake Arch (Canada), Peace
River Arch (Canada), and Sweetgrass Arch (US-
Canadian border) limited or exerted substantial
controls on the initial marine transgression into the
WIFB from the Arctic Ocean (Figure 3; Stelck et al.,
2007). Further south in the United States, the
Transcontinental Arch isolated the northerly-
influenced sea from connecting with the Gulf of
Mexico during intervals characterized by low sea
level (e.g., early Albian-early Cenomanian; Figure 3;
Weimer, 1978; Carlson, 1999; Erickson, 1999; Miall
et al., 2008).
During the earliest (Middle Jurassic- early Late
Jurassic) and middle (late Early Cretaceous-middle
Paleocene) stages of its history, the WIFB was
predominantly a site of marine sedimentation,
whereas its early middle (Late Jurassic-Early
Cretaceous) and latest (late Paleocene-Eocene)
phases were dominated by terrestrial sedimentation
(i.e., the WIS; Kauffman and Caldwell, 1993;
DeCelles, 2004; Miall et al., 2008; Fuentes et al.,
2009, 2011). Regional tectonics had significant
control over the seaway’s shoreline position(s) as
well as on sedimentation rates and patterns across the
basin (Miall et al., 2008). Gill and Cobban’s (1973)
survey of the Montana Group in the northern Great
Plains is among the first studies to recognize the
influence of active tectonic controls on the
paleogeography and sedimentary record of the WIS.
Their study of Campanian and Maastrichtian
biostratigraphy and lithostratigraphy revealed that
transgressions in Montana were correlative with
regressions in Wyoming, which they interpreted as a
tectonically driven regression that was associated
with the initiation of Laramide-style uplift (Gill and
Cobban, 1973). Hancock and Kauffman’s (1979)
study of sea-level change in the WIS showed that
eustasy dominated sea level during the Cenomanian
to Santonian, but became increasingly overprinted by
tectonic controls during the Campanian through
Maastrichtian. Krystinik and Dejarnett’s (1995)
study of the Campanian and Maastrichtian sequence
architecture of the west-central part of North
America documented that tectonic controls were
more significant in the United States segment of the
WIS than in the Canadian part where eustatic
controls dominated its stratigraphic record and its
western shoreline position.
Throughout the region’s history, the deformation
front moved further eastward, and ultimately the
WIFB was subdivided into a series of smaller basins
by block uplift of basement rocks in the United
States segment of the Cordillera (Figure 1; Gill and
Cobban, 1973; Lillegraven and Ostresh, 1991;
DeCelles, 2004; Lawton, 2008; Miall et al. 2008).
Despite an alteration in shortening mechanisms in
the United States, thrust faulting and folding
remained dominant in Canada and Mexico (Figure 1;
DeCelles, 2004). This alteration in shortening
mechanisms can be attributed to a shallowing of the
subduction angle and to differences in the
composition of the overlying crust (DeCelles, 2004;
Lawton, 2008). This change in deformation styles
(i.e., thin-skinned to thick-skinned) marks the onset
of the Laramide Orogeny and the cessation of
significant compressional deformation in the United
States portion of western North America (DeCelles,
2004; Lawton, 2008). This increase in orography
combined with a drop in eustatic sea level forced the
final retreat of the WIS and initiated the renewal of
terrestrial deposition in the WIFB during the
Paleogene (Gill and Cobban, 1973; Lillegraven and
Ostresh, 1991; Miall et al., 2008; Boyd and
Lillegraven, 2011).
Situated east of the WIFB was the stable North
American Craton that was characterized by relatively
low topographical relief (Figures 1, 2). The eastern

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WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
and southern margins of this region were bounded by
the elevated Appalachian and Ouachita (e.g.,
Ouachita-Ozark Interior Highlands) orogenic belts,
which were formed during the Paleozoic (Figure 1;
Arbenz, 1989; Rast, 1989). During the Mesozoic and
Cenozoic, the Appalachian Mountains acted as
barrier that prevented marine inundations into the
continental interior of North America from the
Atlantic Ocean (Williams and Stelck, 1975). During
the Late Cretaceous and Paleogene, uplifted regions,
such as the Ouachita-Ozark Interior Highlands,
prevented marine flooding into the WIFB from the
extension of the Gulf of Mexico known as the
Mississippi Embayment. Marine flooding into the
continental interior was only possible west of the
Ouachita-Ozark Interior Highlands and north of the
Appalachians on the Canadian Shield and into the
Hudson Bay Basin where there was limited
topographical relief (Williams and Stelck, 1975).
RECONSTRUCTING WESTERN INTERIOR
SEAWAY PALEOGEOGRAPHY:
APPROACHES
The ten paleogeographic maps presented in this
paper show key changes in the distribution of land
and sea in both the Western Interior and North
America during the Early Cretaceous to Paleocene.
The paleogeographic reconstructions of the southern
portion (southern Alberta to Texas) of the WIS were
primarily constructed by W. A. Cobban over his
career examining the Cretaceous stratigraphy and
paleontology of the west-central United States.
Several of these reconstructions have been published
in earlier works (e.g., Gill and Cobban, 1973;
Cobban and Hook, 1984; Cobban et al., 1994).
Albian to Paleocene shoreline configurations of the
southern part of the WIS are also compiled from
detailed maps presented by many other studies of its
paleogeography (e.g., Reeside, 1957; Sloss et al.,
1960; Sohl, 1967; McGookey, 1972; Williams and
Stelck, 1975; Kauffman, 1984; Lillegraven and
Ostresh, 1991; Kauffman and Caldwell, 1993;
Sageman and Arthur, 1994; Erickson, 1999; Roberts
and Kirschbaum, 1995; Boyd and Lillegraven, 2011;
Landman et al.; 2012). Furthermore paleogeographic
reconstructions of the eastern and northern portions
of the WIS during the Cretaceous are largely based
on maps from various sources (e.g., Birkelund, 1965;
Jeletzky, 1971; Williams and Stelck, 1975; Witzke et
al., 1983; Kauffman, 1984; Cobban et al., 1994;
Sageman and Arthur, 1994; Ziegeler and Rowley,
1996; White et al., 2000; Stelck et al., 2007; Nielsen
et al., 2008; Schröder-Adams, 2014), as are those
used for the Pacific Coast, Arctic Coast, Gulf and
Atlantic Coastal plains, and Mexico (e.g., Jeletzky,
1971; Alencaster, 1982; Scott, 1984; Owens and
Gohn, 1989; de Cserna, 1989; Sohl et al., 1991;
McFarlan, Jr., and Menes, 1991; Galloway et al.,
1991; Landman et al., 2004; Umhoefer and Blakey,
2006; Galloway, 2008; Blakey, 2013; Schröder-
Adams, 2014).
These WIS paleogeographic reconstructions
through time and controls on relative sea level in the
WIFB are based on detailed documentation of both
global and regional tectonic reconstructions, paleo–
Figure 4. Upper Jurassic to Lower Cretaceous
biostratigraphic chart for the Canadian Western Interior (1)
Gradstein, 2012; 2) Stott et al., 1993).

30 WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
Figure 5. Summary chart of Lower Cretaceous to Paleocene chronostratigraphy, biostratigraphy, magnetostratigraphy,
and second-order sea-level fluctuations in the US and Canadian Western Interior (1) Gradstein, 2012; 2) Cobban et al.,
2006; 3) Scott, 2007 4) Stott et al., 1993 5) Braunberger and Hall, 2001; 6) Lillegraven and Ostresh, 1991; 7) Cifelli et
al., 2004; 8) Hicks et al., 1999; 9) Kauffman, 1969).

31
WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
biogeography, lithostratigraphy, biostratigraphy, and
chronostratigraphy. The most common method
employed to reconstruct WIS paleogeography has
been through detailed measuring and interpretation
of lithostratigraphic sections across the region as
well as the collection of biostratigraphic data derived
from various organisms, especially ammonites and
inoceramids (e.g., Gill and Cobban, 1973; Stelck et
al., 2007). These data are used to correlate and
constrain different lithofacies, distinguish regional
from more local changes that influenced the WIS/
WIFB, and facilitate recognition of underlying
controls (e.g., eustatic vs. tectonic) on the sea-level
record preserved in the Western Interior. The Lower
Cretaceous to Paleocene marine and terrestrial strata
of this basin has one of the most continuous and
highly resolved biostratigraphic frameworks in the
world, which is constrained by radiometric dates of
ash beds, magneto-, and cyclostratigraphy (Figure 4,
5; Fox and Ross, 1942; Fox and Olson, 1969;
Lillegraven and Ostresh, 1991; Obradovich, 1993;
Kauffman et al., 1993; Kowalis et al., 1995; Hicks et
al., 1999; Cifelli et al., 2004; Anderson et al., 2006;
Cobban et al., 2006; Scott, 2009; Scott et al., 2009;
Wilson et al., 2010). The biostratigraphic
frameworks for the northern and southern parts of
the WIFB differ slightly due to a paucity of modern
biostratigraphic work in the former and because the
northern part has a more continuous Lower
Cretaceous biostratigraphic record due to its longer
interval of marine deposition (Figure 5; Stott et al.,
1993).
Another important method used to reconstruct the
paleogeography of the WIS is to map the thicknesses
of Cretaceous to Paleogene rocks in the Western
Interior and use the resulting isopach maps to
interpret the changing geometry of the WIFB
through time (e.g., Cross, 1986; Roberts and
Kirschbaum, 1995; DeCelles, 2004). Isopach maps
have been also used to document the migration of
different depozones through time and to depict the
WIFB breakup into smaller fault-bounded basins
during the Laramide Orogeny (e.g., Cross, 1986;
Roberts and Kirschbaum, 1995; DeCelles, 2004).
The biogeographic ranges of fossil occurrences
within as well as outside of the Western Interior have
revealed critical information on how and when the
WIS was connected with open-shelf seas (e.g., Arctic
Ocean, proto-Gulf of Mexico) along the margins of
the North American continent (e.g., Birkelund, 1965;
Sohl, 1967; Jeletzky, 1971; Williams and Stelck,
1971; Cobban, 1993; Cvancara and Hoganson, 1993;
Boyd and Lillegraven, 2011). This is important for
certain intervals and parts of North America where
there is a poor stratigraphic record due to extensive
erosion and/or a lack of exposure (i.e., Canadian
Shield and eastern North American Craton; Ziegler
and Rowley, 1998; White et al., 2000). This type of
data can reveal paleobiogeographic connections,
which, in the absence of direct stratigraphic data,
would remain obscure (e.g., Birkelund, 1965; Sohl,
1967; Jeletzky, 1971; Boyd and Lillegraven, 2011).
CONTROLS ON SEA LEVEL IN THE
WESTERN INTERIOR SEAWAY
As noted above, the paleogeography of the WIS
was primarily controlled by the interaction of sea
level with physiography (Williams and Stelck, 1971;
Gill and Cobban, 1973; Jeletzky, 1980; Lillegraven
and Ostresh, 1991; Krystinik and DeJarnett, 1995;
Caldwell and Kauffman, 1993; Miall et al., 2008).
Sea level in the WIFB was modulated by eustasy,
long considered an important driver of sea-level
change and, in turn, a control on the paleogeography
of the WIS (Kauffman, 1977, 1984; Weimer, 1984;
Hallam, 1992), and tectonic controls that altered the
basin’s physiography through uplift, subsidence, and
sedimentation; the interplay of which resulted in
localized transgressions and regressions. These
different mechanisms interacted both individually
and simultaneously, at times in concert and at other
times in opposition, on different spatial and temporal
scales as primary controls on the seaway’s spatial
extent, its paleoshorelines, and on its lithofacies (Gill
and Cobban, 1973; Jeletzky, 1980; Lillegraven and
Ostresh, 1991; Caldwell and Kauffman, 1993;
Krystinik and DeJarnett, 1995; DeCelles, 2004;
Miall et al., 2008). Several different authors have
named and described the various second-order sea-
level cycles or transgressive-regressive phases of the
WIS that are represented by the history of its
shoreline positions and on the stratigraphic record of
the WIFB (Figure 5; e.g., Greenhorn Cycle of
Kauffman, 1967; Fox Hills Regression of Gill and
Cobban, 1973). Each sea-level cycle is characterized
by differing degrees of influence from both eustatic
and tectonic controls.
In terms of the overall sea-level history, the WIS

32 WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
is characterized by a rise in eustatic sea level during
the early Albian to Paleocene accompanied by
increased basin subsidence (Figure 6; Kauffman and
Caldwell, 1993). The most-accepted explanations for
this elevated eustatic record during the Early
Cretaceous to Paleocene are: 1) a higher rate of sea-
floor spreading that increased the volume of mid-
ocean ridges (Pitman, 1978; Pitman and
Golovchenko, 1983; Arthur et al., 1985, 1991;
Lillegraven and Ostresh, 1991), and 2) the
“greenhouse” phenomenon, resulting from higher
CO2 concentrations degassed from elevated volcanic
activity and, therefore, the lack of significant ice
sheets (e.g., Arthur et al., 1985, 1991; Miller et al.,
2003; Hay, 2008). In this hypothesis, these
mechanisms could have elevated eustatic sea level to
a maximum of approximately 300 m higher than
present day and resulted in the expansion of marine
environments into continental interiors (Kauffman
and Caldwell, 1993).
In contrast, the Miller et al. (2005) analysis of
Phanerozoic sea level argued for variation in CO2
concentration as the primary cause and driver of sea-
level change during the Cretaceous and contended
that eustatic sea level during the Cretaceous peaked
at 100 ± 50 m, which would indicate much slower
rates of ocean-crust production than previously
estimated. They went on to suggests that short-term
variations in sea level were potentially controlled by
changes in the volume of Antarctic ice sheets, which
in turn would have modulated temperature changes
that were associated with tectonically sourced CO2
input. A number of studies have argued that short-
term variations in sea level that would have
modulated shoreline movement of the WIS during
the Cretaceous were controlled by Milankovitch
cycles (e.g., Elder et al., 1994; Sageman et al., 1997).
Despite recognition of the importance of tectonic
controls on sea level in the WIFB, most studies have
viewed eustasy as the primary driver of cyclic sea-
level change and shoreline movement in the WIS
(e.g., Weimer, 1984; Kauffman, 1977, 1984).
Beginning in the 1970’s, this view was criticized,
and the importance of tectonics along with eustasy in
controlling sea level and shoreline movement
became increasingly more recognized (e.g., Gill and
Cobban, 1973; Jeletzky, 1979).
WESTERN INTERIOR SEAWAY
PALEOGEOGRAPHIC EVOLUTION
During the Early Jurassic, the future location of
the WIFB was characterized by predominantly
terrestrial passive margin sedimentation (Miall,
2009). Sediments representing this time span were
derived from the Appalachian and Ouchita orogenic
belts to the east and southeast, respectively (Blakely
and Ranney, 2008). Throughout this interval,
western North America was virtually tectonically
neutral with transtension governing the plate margin
and future retroarc (DeCelles, 2004). By the Middle
Jurassic, a westerly derived sediment source had
initiated as a result of the amalgamation of the
continental-margin arc-trench tectonic system and
the onset of Cordilleran orogenic activity along the
western margin of North America (DeCelles, 2004;
Fuentes et al., 2009; 2011). As this orogenic uplift
proceeded and continued to deform the North
American craton to the east, it led to the formation
the WIFB during the Bajocian (DeCelles and Currie,
Figure 6. Phanerozoic global sea level curves and
temperature curve (1) Gradstein, 2012; 2) Frakes et al.,
1992; 3) Miller et al., 2005). The shaded area represents the
interval characterized by marine flooding in the Western
Interior Foreland Basin.

33
WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
Figure 7. Generalized middle Valanginian (Buchia inflata time) paleogeographic map of North America showing marine
inundation from Pacific Ocean into the Western Interior Foreland Basin (Western Interior shorelines based on Jeletzky,
1971; Williams and Stelck, 1975; shoreline outside Western interior modified from Alencaster, 1984; Owens and Gohn,
1985; McFarlan and Menes, 1991; Goldhammer, 1999; Blakey, 2013). Shaded areas represent land.

34 WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
1996; DeCelles, 2004; Fuentes et al., 2009; 2011).
The basin reflected increased subsidence as a result
of the initiation of subduction of the Panthalassan
crust beneath western North America (i.e.,
subduction-induced subsidence) and formation of the
Cordilleran Fold-Thrust Belt (i.e., load-induced
subsidence; Fuentes et al., 2009, 2011).
Westerly derived sediment sources became
dominant during the Middle Jurassic and persisted as
such until the latest Eocene (Fuentes et al., 2009,
2011). This interval (i.e., Bajocian) is characterized
by the transgression of the shallow, epeiric Sundance
Sea into the WIFB (Imlay, 1980; 1984). This sea was
an extension of Panthalassa and covered most of
modern-day Utah, Idaho, Wyoming, Montana,
Alberta, and southern Saskatchewan (Imlay, 1980).
During its history the Sundance Sea underwent
several different transgressive and regressive phases,
which formed numerous, but significant
unconformities throughout the Middle to Upper
Jurassic strata in the Western Interior (McMullen et
al., 2014).
During the Late Jurassic (i.e., Kimmeridgian), the
shoreline of the Sundance Sea was prograding
northward into present-day Alberta in association
with a drop in eustatic sea level (Figure 6; Imlay,
1980; 1984). In northern Montana and southern
Alberta, this regression is characterized, by marine
shales and sandstones of the Upper Jurassic Fernie
Formation being overlain by the deltaic and
terrestrial coal-rich lithofacies of the Kootenay
Formation (Stockmal et al., 1992; Miall et al., 2008).
The deltaic deposits of the Kootenay Formation
represent the final marine incursion by the Sundance
Sea from the Pacific Ocean into the United States
portion of the Western Interior during the Jurassic
(Imlay, 1980, 1984). In southern Montana,
Wyoming, Colorado, Utah, New Mexico, and South
Dakota, the post- Sundance stratigraphic record is
characterized by riverine and lacustrine deposits of
the Morrison Formation (DeCelles and Currie, 1996;
Demko et al., 2004; Foster, 2007). As the foredeep
depozone was located in central Utah and Idaho
during this interval, most of the Morrison Formation
was deposited in the WIFB’s back-bulge depozone
(s) (DeCelles and Currie, 1996; Currie, 1998).
Sediments deposited into the wedge-top, foredeep,
and forebulge depozone(s) were likely eroded during
subsequent Cretaceous thrust-fault uplifts (DeCelles
and Currie, 1996; DeCelles, 2004).
Terrestrial environments persisted across most of
the WIFB (except for the most northern part)
throughout the latest Jurassic into the Early
Cretaceous (i.e., Kimmeridgian to early Albian;
DeCelles, 2004). In British Columbia and
northwestern Alberta, Berriasian and Valanginian
marine strata deposited in the Peace River
Embayment represent the final Mesozoic marine
incursion into the Canadian portion of the WIFB
from the Pacific (Figure 7; Williams and Stelck,
1975). Post-Valanginian marine inundations into the
Western Interior from the Pacific Ocean were most
likely prevented by the hypsographic rise of the
Cordilleran Orogenic Belt in the western part of
Canada (Williams and Stelck, 1975). This rise also
shifted the principal drainage of the Western Interior
from the Pacific to the Arctic Ocean or Western
Interior Sea (i.e., the extension of the Arctic Ocean
into the northern part of WIFB that was isolated
from the Gulf of Mexico by various topographical
highs) until the end of the Early Cretaceous when the
WIS formed (Williams and Stelck, 1975).
In the United States, it has been traditionally
thought that a widespread unconformity separates the
top of the Tithonian or Berriasian-Valangian (Upper
Jurassic to Lower Cretaceous) Morrison Formation
from the overlying Aptian-Albian (upper Lower
Cretaceous) strata (DeCelles and Burden, 1992;
Dyman et al., 1994; Foster, 2007, Zaleha and
Wiesemann, 2004; Zaleha, 2006; Elliot et al., 2007).
This long-term depositional hiatus primarily reflects
inadequate chronostratigraphic sampling from Lower
Cretaceous st rata and use of out dated
biostratigraphic data (Zaleha, 2006; Sames et al.,
2010). Recent studies utilizing new biostratigraphic
data have revealed that this unconformity represents
a much shorter time span than previously thought,
varies in duration across the basin, and is represented
by multiple unconformities bounding a number of
Lower Cretaceous sequences (Zaleha, 2006; Sames
et al., 2010). Zaleha’s (2006) use of new
chronostratigraphic data identified three Lower
Cretaceous unconformities in the United States that
formed prior to the initiation of widespread marine
deposition during the Albian. Both Zaleha (2006)
and Sames et al. (2010) have also shown that strata
formerly attributed to Aptian and Albian actually
represents Berriasian to Albian deposition within the

35
WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
Figure 8. Generalized early Albian (Lemuroceras cf. L. indicum time) paleogeographic map of North America showing
initial marine transgression into the Western Interior Foreland Basin from Arctic Ocean (Clearwater Sea shorelines
modified from Stelck et al., 2007; shorelines outside Western interior modified from Alencaster, 1984; Owens and Gohn,
1985; McFarlan and Menes, 1991; Goldhammer, 1999; Blakey, 2013). Shaded areas represent land.

36 WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
WIFB.
Fluvial, flood plain, and lacustrine deposition
persisted in the southern half of the WIFB up to the
start of the early Albian marine transgression that
would form the WIS (DeCelles and Currie, 1996;
Zaleha, 2006; DeCelles, 2004). Most of the
paleochannels found within these strata were
deposited as point bars in meandering rivers that
show east to north paleoflow direction within the
WIFB (Zaleha, 2013). It is even possible that some
rivers along the eastern margin of the basin were
flowing westward from the vicinity of the
Transcontinental Arch (Zaleha, 2013).
Examination of the strata overlying these Berri-
iasian to Aptian beds indicates that load-induced
flexural subsidence associated with increased
thrusting was active during the late Albian
(DeCelles, 2004; Miall et al., 2008). This subsidence
coincided with the first large-scale marine incursions
by the Western Interior Sea or Clearwater Sea, as it
is also known, into the WIFB from the Arctic Ocean
(Figure 8; Jeletzky, 1971; Williams and Stelck,
1975; Kauffman and Caldwell, 1993; Brenner et al.,
2000). This marine transgression is characterized by
several punctuated, southward-directed, early to
middle Albian flooding events that were primarily
restricted to Canada (Williams and Stelck, 1971;
Stelck et al., 2007). The limited spatial extent of
these transgressions to the Canadian Western Interior
and punctuated nature of these transgressions was
primarily due to low eustasy and pre-existing,
topographically elevated structural features such as
the Dave Lord-Eskimo Lakes, Tathlina-Liard River,
Peace River, and Sweetgrass arches (Figure 3;
Kauffman and Caldwell, 1993; Stelck et al., 2007).
As sea level rose, the advance of the Clearwater Sea
was temporally impeded by each of these elevated
highs, which would eventually be transgressed with
continued sea-level rise (Stelck et al., 2007). Further
south, the WIFB was the site of numerous north-
flowing fluvial systems that drained into the
Clearwater Sea (Stelck et al., 2007; Miall et al.
2008).
During the latest Albian, eustatic-sea-level rise
culminated in the first marine connection between
the Gulf of Mexico and the Clearwater Sea
somewhere in the vicinity of Colorado (i.e., Kiowa-
Skull Creek Transgression; Figures 9; Reeside, 1957;
Sloss, 1963; Williams and Stelck, 1975; Kauffman
and Caldwell, 1993; Brenner et al., 2000; Oboh-
Ikuenobe et al., 2009). This transgression formed the
Skull Creek Seaway, which divided North America
into the isolated landmasses of Appalachia and
Laramidia (Figure 9; Archibald, 1996). It was the
first transgression to extend past the intracratonic
Transcontinental Arch (Figure 3, 9; Williams and
Stelck, 1975; Kauffman and Caldwell, 1993;
Carlson, 1999; Miall et al., 2008; Oboh-Ikuenobe et
al., 2009). This unification of northern and southern
water masses is indicated by marine sediment
deposition on the Transcontinental Arch and by the
occurrence of mollusks with strong affinities to the
Gulf of Mexico in Canada (Jeletzky, 1971; Williams
and Stelck, 1975). Faunal evidence also supports a
connection with the Arctic Ocean during this interval
(Williams and Stelck, 1975). The location of the
eastern and western shorelines of the Skull Creek
Seaway are well delineated in the southern portion of
the Western Interior based on nearshore strata from
this interval preserved on both sides of the basin
(Figure 9; Witzke et al., 1983; Cobban et al., 1994).
Stratigraphic and faunal evidence from Quebec,
Ontario, and Labrador suggests that this
transgression also resulted in the first marine
connection between the Skull Creek Seaway (or
WIS) and the Atlantic Ocean via the Hudson Seaway
(Figure 9; White et al., 2000).
Eustatic sea level briefly fell during the earliest
Cenomanian, and the connection between the Gulf of
Mexico and Arctic Ocean was temporarily disrupted
(i.e., Kiowa-Skull Creek Regression; Figures 10;
Miall et al., 2008; Oboh-Ikuenobe et al., 2009; Scott
et al., 2009). During this interval, the Mowry Sea (or
Western Interior Sea) extended from the Arctic
Ocean southward to the Transcontinental Arch
(Figure 10; Williams and Stelck, 1975; Cobban et
al., 1994; Scott et al., 2009). This structural feature
separated the Arctic-Ocean-influenced Mowry Sea
from a northern extension of the Gulf of Mexico to
the south, which extended northward up to southern
Colorado, and Kansas (Williams and Stelck, 1975;
Weimer, 1978; Miall et al., 2008; Scott et al., 2009).
A connection with the Arctic Ocean is indicated by
an endemic fauna with strong affinities to Arctic
faunas (Reeside and Cobban, 1960; Williams and
Stelck, 1975; Yacobucci, 2004). However, the
endemic nature of the fauna suggests that the
opening between the Arctic Ocean and Mowry Sea

37
WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
Figure 9. Generalized late Albian (Inoceramus comancheanus time) paleogeographic map of North America showing
first full-marine connection between the Arctic Ocean and Gulf of Mexico in Western Interior Foreland Basin (Skull
Creek Seaway shorelines based on unpublished maps of W.A. Cobban; Jeletzky, 1971; shorelines outside Western
interior modified from Alencaster, 1984; Owens and Gohn, 1985; McFarlan and Menes, 1991; Goldhammer, 1999;
White et al., 2000; Blakey, 2013). Shaded areas represent land.

38 WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
Figure 10. Generalized early Cenomanian, (Neogastroplites cornutus time) paleogeographic map of North America
showing loss of marine connection between the Arctic Ocean and Gulf of Mexico in the Western Interior Foreland Basin
(Mowry Sea shorelines based on unpublished maps of W.A. Cobban; Jeletzky, 1971, Williams and Stelck, 1975;
shorelines outside Western interior modified from Alencaster, 1984; Owens and Gohn, 1985; McFarlan and Menes,
1991; Goldhammer, 1999; White et al., 2000; Blakey, 2013). Shaded area represents land.

39
WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
was likely periodically restricted or even closed
(Williams and Stelck, 1975; Schröder-Adams, 2014).
This biogeographic restriction between the Arctic
Ocean and Mowry Sea may have been caused by
exceptionally low sea levels or the emplacement of
the Mackenzie Salient along the border between the
Yukon and Northwest Territories, which resulted in
an eastward shift in the western shoreline in this
region (Jeletzky, 1971; Williams and Stelck, 1975).
The location of the Mowry Sea’s western shoreline
was situated in western Alberta, western Montana
and along the Wyoming-Idaho border, whereas the
locations further to the south and along the eastern
margin are speculative (Figure 10). The poor
stratigraphic record on the Canadian Shield makes it
difficult to determine if a marine connection between
the Mowry Sea and the Atlantic Ocean was
maintained through the Hudson Seaway during this
drop in eustatic sea level. The onset of a new
transgressive phase and reunification with the Gulf
of Mexico is first evident by the appearance of
ammonites and foraminifers with affinities to faunas
from the Gulf of Mexico in the Neogastroplites
maclearni ammonite biozone of Alberta, Montana,
and Wyoming (Williams and Stelck, 1975;
Yacobucci, 2004).
A eustatic rise resulted in the ‘Great
Transgression’ (or more regionally known as the
Greenhorn Transgression) during the middle
Cenomanian through the early Turonian, which re-
established the interconnection between the Gulf of
Mexico and the Arctic Ocean (Figure 11; Hancock
and Kauffman, 1979; Kauffman and Caldwell, 1993;
Cobban, 1993; Oboh-Ikuenobe et al., 2009; Schröder
-Adams, 2014). This continuity persisted at least
until the early Maastrichtian and possibly into the
Paleocene (Williams and Stelck, 1975; Lillegraven
and Ostresh, 1991; Kauffman and Caldwell, 1993;
Boyd and Lillegraven, 2011). This connection is
represented by a continuous lithostratigraphic record
across the basin (including the Transcontinental
Arch) and the first occurrence of diverse taxa in the
WIS with strong affinities to Gulf of Mexico faunas
(Cobban, 1993; Kauffman and Caldwell, 1993). The
occurrence of ammonites in Canada with affinities to
Alaskan faunas indicates the persistence of a
connection with the Arctic Ocean (Jeletzky, 1971).
There is strong paleontologic and geologic evidence
on the Canadian Shield for the existence of the
Hudson Seaway during this time (Williams and
Stelck, 1971; Ziegler and Rowley, 1998; White et al.,
2000; Schröder-Adams, 2014).
Throughout this interval, the seaway underwent a
number of small transgressions and regressions
(Jeletzky, 1971; Williams and Stelck, 1975; Cobban
and Hook, 1984; Roberts and Kirschbaum, 1995),
likely Milankovitch controlled, that strongly
influenced the basin’s stratigraphic architecture and
shoreline positions (Barron et al., 1983; Meyers et
al., 2001). The location of the western shoreline
during the Cenomanian varied over 600 km in the
southwestern part of the United States (i.e., south
central Utah to south central New Mexico) and in
northern Alberta to less than 100 km in western
Wyoming, western Montana, and southwestern
Alberta (Roberts and Kirschbaum, 1995). In
northern Alberta, the Dunvegan Delta prograded
eastward across the basin as a result of increased
sediment supply associated with uplift caused by the
accretion of the Insular Superterrane (Stockmal et
al., 1992; Plint, 2000). During the same interval, the
Woodbine Delta prograded westward (in the vicinity
of Oklahoma) into the WIFB and southward (in the
vicinity of northeast Texas) into the Gulf of Mexico
(Roberts and Kirschbaum, 1995). Further north, the
positions of the eastern shoreline are poorly
constrained, but it most likely varied from western
Minnesota and Iowa during peak transgressions to
the eastern Dakotas, eastern Nebraska, and eastern
Kansas during regressions. This mid-Cretaceous
expansion of the WIS was one of the most extensive
transgressions in North American history and was
characterized by deposition of widespread offshore
siliciclastic and carbonate muds across most of the
basin and thick successions of nearshore clastic
sediments along the coastlines (Reeside, 1957;
Jeletzky, 1971; Williams and Stelck, 1975;
Kauffman and Caldwell, 1993).
During the early Turonian Greenhorn Cycle, the
WIS reached its greatest geographic extent (Figure
11; McDonough and Cross, 1991; Kauffman and
Caldwell, 1993; Sageman and Arthur, 1994). For the
duration of this interval, the sea extended from
central Utah to possibly Wisconsin and from Texas
to the Arctic Ocean (Figure 11). There is evidence
that the sea expanded over the Mackenzie Salient
that previously restricted connection between the sea
and Arctic Ocean (Jeletzky, 1971). Faunal

40 WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
Figure 11. Generalized late Cenomanian (Neocardioceras juddii time) paleogeographic map of North America (WIS
shorelines based on unpublished maps of W.A. Cobban; Jeletzky, 1971; Roberts and Kirschbaum, 1997; shorelines
outside Western Interior modified from Alencaster, 1984; Owens and Gohn, 1985; de Cserna, 1989; McFarlan and
Menes, 1991; Goldhammer, 1999; Blakey, 2013). Shaded areas represent land.

41
WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
similarities between the WIS and Pacific coastal
fauna led Lang and McGugan (1987) to tentatively
suggest that a marine connection formed along the
Alberta-Montana border between the Pacific Ocean
and WIS during the Turonian peak transgression.
Despite some faunal similarities, relatively little
additional research has been devoted to testing this
hypothesis and, in light of current geological (e.g.,
tectonic, stratigraphic) evidence, it seems unlikely.
The Turonian is the youngest interval with direct
stratigraphic and paleontological evidence preserved
on the Canadian Shield for the existence of the
Hudson Seaway (Figure 11; Williams and Stelck,
1975; Ziegler and Rowley, 1998). The areal extent of
this sea over the Canadian Shield during this interval
is still poorly constrained due to post-Mesozoic
erosion. However, it could have been much more
extensive than the area depicted on typical paleogeo-
graphic maps (e.g., Jeletzky, 1971; William and
Stelck, 1975; Ziegler and Rowley, 1998; White et al.,
2000). Limestone-marl couplets were deposited
across northern Texas to northern North Dakota and
from western Colorado to eastern Kansas (Elder et
al., 1994; Sageman and Arthur, 1994; Roberts and
Kirschbaum, 1995; Sageman et al., 1998; Meyers et
al., 2001; Keller et al., 2004). These carbonate
deposits can be correlated across the basin to clastic
sequences along its margins (Elder et al., 1994).
Carbonate deposition ended during the middle
Turonian in association with a drop in sea level that
was controlled by tectonics and eustatic changes
(Greenhorn Regression; Reeside, 1957; Williams and
Stelck, 1975; Roberts and Kirschbaum, 1995;
Merewether et al., 2007; Nielsen et al., 2008; Miall
et al., 2008). The surface area of the WIS shrunk
drastically during this interval, and it is possible that
the connection between the WIS and the Atlantic
Ocean was lost or restricted (Figure 12; Nielsen et
al., 2008). During the middle to late Turonian, the
location of the western shoreline varied over 300 km
in areas south of the Montana-Wyoming border to
less than 150 km north of this border (Roberts and
Kirschbaum, 1995). This regression is associated
with the deposition of major sand complexes across
the basin during the late Turonian, such as the upper
Frontier Formation in Wyoming and the Cardium
Formation in Alberta (Williams and Stelck, 1975;
Roberts and Kirschbaum, 1995).
Relative sea level rose rapidly during Coniacian
and Santonian time (i.e., Niobrara Transgression),
resulting in a resurgence of carbonate deposition in
the center of the basin, widespread deposition of
muds across much of Alberta, Montana, Wyoming,
Colorado, and Utah, as well as the development of
clastic wedges along the basin margin in central
Utah, eastern Idaho, Montana, Alberta, and
northeastern British Columbia (Figure 13; Roberts
and Kirschbaum, 1995; Merewether et al., 2007;
Nielsen et al., 2008). The location of the western
shoreline during the Coniacian and Santonian
remained fairly static, but probably varied within a
150 km range along most of its length (Roberts and
Kirschbaum, 1995). The location of the eastern
shoreline during this interval is poorly known, but
was probably located in central Iowa and Minnesota
(Witzke et al., 1983). Isopach maps of this interval
also show a significant change in the basin’s
geometry reflecting the demise of the well-organized
flexural-foreland-basin system (Cross, 1986;
DeCelles, 2004).
The Coniacian and Santonian represent the
greatest marine inundation of the Canadian Western
Interior and Arctic Archipelego during the
Cretaceous (Figure 13; Jeletzky, 1971). This fact
lends support for a possible connection between the
WIS and the Atlantic Ocean via the Hudson Seaway
during this interval (Figure 13; Williams and Stelck,
1975). The presence of Western Interior ammonite
species in western Greenland during the Late
Turonian to Santonian supports this hypothesis
(Birkelund, 1965; Williams and Stelck, 1975).
However, it is possible, as suggested by Birkelund
(1965), that these faunas were using an Arctic route
since a connection (i.e., the Teichert Strait of Ziegler
and Rowley, 1998) was likely made between the
Sverdrup Basin and Baffin Bay during this interval
(Jeletzky, 1971; Williams and Stelck, 1975) forming
the Labrador Seaway (Figure 13).
Tectonic activity, closely associated with the
accretion of the Insular Superterrane and uplift of the
Purcell Anticlinoform (Stockmal et al., 1992),
increased along the Cordilleran Fold-Thrust Belt
during the Campanian. Furthermore, Lawton (1994)
suggested that this activity was amplified in the
United States portion due to an increase in the rate of
orthogonal convergence between the Farallon and
North American plates. Isopach maps of Campanian
-Maastrichtian strata indicate a broad area of thick

42 WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
Figure 12. Generalized late Turonian (Prionocyclus germari time) paleogeographic map of North America (WIS
shorelines based on unpublished maps of W.A. Cobban; Jeletzky, 1971; Witzke et al., 1990; Roberts and Kirschbaum,
1997; Nielsen et al., 2008; shorelines outside Western interior modified from Alencaster, 1984; Owens and Gohn, 1985;
de Cserna, 1989; Sohl et al., 1991; McFarlan and Menes, 1991; Goldhammer, 1999; Blakey, 2013). Shaded areas
represent land.

43
WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
Figure 13. Generalized middle Coniacian (Scaphites ventricosus time) paleogeographic map of North America (WIS
shorelines based on unpublished maps of W.A. Cobban; Jeletzky, 1971; Witzke et al., 1990; Roberts and Kirschbaum,
1997; Nielsen et al., 2008; shorelines outside Western interior modified from Alencaster, 1984; Owens and Gohn, 1985;
de Cserna, 1989; Sohl et al., 1991; McFarlan and Menes, 1991; Goldhammer, 1999; Blakey, 2013). Shaded areas
represent land.

44 WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
strata in central Wyoming and Colorado, eastward of
the foredeep, indicating a strong influence of
subduction-induced subsidence (Cross, 1986). This
change in tectonic activity coincided with a
permanent shift from carbonate to siliciclastic mud
deposition in the basin center (Niobrara Regression;
Miall et al., 2008). This interval is stratigraphically
represented in the central part of the Western Interior
by the transition from the Niobrara Formation to the
Pierre Shale (Gill and Cobban, 1973; Roberts and
Kirschbaum, 1995). The locations of the western and
eastern shorelines in the United States during this
time were probably similar to the Coniacian and
Santonian (Figure 14; Witzke et al., 1983). However,
in Canada the western shoreline began to migrate
back and forth within distances spanning 250-650
km (Roberts and Kirschbaum, 1995). The shared
occurrence of ammonites in Greenland and the WIS
supports a northern connection into the middle
Campanian (Figure 14; Birkelund, 1965). However,
whether this biogeographic route was through the
Arctic Ocean or Hudson Seaway is unclear in the
absence of direct evidence (Birkelund, 1965).
Schröder-Adams (2014) argued that biogeographic
similarities between the WIS and Labrador Seaway
were most likely maintained by the Hudson Seaway.
Beginning in the middle Campanian, a major
tectonic reorganization associated with the onset of
the Laramide Orogeny resulted in a cratonward shift
in marine sedimentation and the expansion of a
broad coastal plain that persisted into the
Maastrichtian (Clagget and Bearpaw marine cycles;
Figures 14, 15; Gill and Cobban, 1973; Roberts and
Kirchsbaum, 1995; Miall et al., 2008). The western
shoreline shifted eastward to central Montana,
central Wyoming, and western Colorado during the
peak transgressions of the middle Campanian (Gill
and Cobban, 1973; Lillegraven and Ostresh, 1990;
Roberts and Kirschbaum, 1995). During regressions
it was situated in eastern Montana, eastern
Wyoming, central Colorado, and central to eastern
New Mexico (Gill and Cobban, 1973; Lillegraven
and Ostresh, 1990; Roberts and Kirschbaum, 1995).
The location of the eastern shoreline for the
Campanian, Maastrichtian, and Danian is completely
unknown because of erosion and a lack of available
exposures representing nearshore strata along the
cratonic side of the basin (Witzke et al., 1983;
Erickson, 1999). However, the location of the
eastern shoreline for the latest Cretaceous and
earliest Paleogene was probably similar to earlier
intervals. Williams and Stelck (1975) suggested that
the seaway might have been connected to the
northern part of the Mississippi Embayment along its
eastern side and that the Ouachita-Ozark Interior
Highlands might have been multiple islands (or an
island) during the Campanian and Maastrichtian.
Despite being possible, this hypothesis remains
untested and without further study will remain
tentative. Lithostratigraphic evidence for a
connection to the Arctic Ocean ends at this time due
to a substantial regression; however, faunal data
provide support for the retention of an Arctic
connection throughout this interval (Cobban, 1993;
Erickson, 1999). There is currently no paleontologic
or stratigraphic evidence for a connection to the
northern Atlantic Ocean through the Hudson Seaway
in the late Campanian, but it is possible that faunas
with affinities to the Arctic Ocean were entering into
the WIS via the Hudson Seaway, as the connection
that opened up during the Coniacian between the
Sverdrup Basin and Baffin Bay was maintained into
the Maastrichtian (Ziegler and Rowley, 1998).
Throughout the Maastrichtian (Figure 15), the
western shoreline migrated basinward as eustatic sea
level fell and the WIFB started to segregate into
nonmarine sub-basins with the onset of the Laramide
Orogeny (Bearpaw Regression; Waagé, 1968; Gill
and Cobban, 1973; Dickinson et al., 1988; Kauffman
and Caldwell, 1993; Lillegraven and Ostresh, 1991;
Miall et al., 2008). The western shoreline retreated
during the early to late Maastrichtian in northeastern
Wyoming, Montana, and in the western Dakotas due
to the Sheridan Delta’s progradation across the basin
(Figure 15; Gill and Cobban, 1973; Krystinik and
Dejarnett, 1996, Kennedy et al., 1998; Pyles and
Slatt, 2007). The western shoreline’s location is not
clearly resolved for the late Maastrichtian; however,
recent studies of the preserved stratigraphic and
paleontological marine record have revealed that it
extended from southwestern South Dakota to south
central North Dakota (Landman et al., 2012). Based
on the available evidence, the seaway was most
likely located to the east of the current outcrop belt
in the vicinity of the Great Plains (Williams and
Stelck, 1975). Occurrences of marine taxa in upper
Maastrichtian strata with affinities to both the Gulf
Coast and Arctic regions also strongly support a

45
WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
Figure 14. Generalized middle Campanian (Baculites obtusus time) paleogeographic map of North America (WIS
shorelines based on unpublished maps of W.A. Cobban; Jeletzky, 1971; Roberts and Kirschbaum, 1997; shorelines
outside Western interior modified from Alencaster, 1984; Owens and Gohn, 1985; de Cserna, 1989; Sohl et al., 1991;
McFarlan and Menes, 1991; Goldhammer, 1999; Umhoefer and Blakey, 2006; Blakey, 2013). Shaded areas represent
land.

46 WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
Figure 15. Generalized early Maastrichtian (Baculites clinolobatus time) paleogeographic map of North America (WIS
shorelines based on unpublished maps of W.A. Cobban; Jeletzky, 1971; Roberts and Kirschbaum, 1997; shorelines
outside Western interior modified from Alencaster, 1984; Owens and Gohn, 1985; de Cserna, 1989; Sohl et al., 1991;
McFarlan and Menes, 1991; Goldhammer, 1999; Landman et al., 2004; Umhoefer and Blakey, 2006; Blakey, 2013).
Shaded areas represent land.

47
WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
Figure 16. Generalized Danian paleogeographic map of North America (Cannonball Sea shorelines based on Catuneanu
and Sweet, 1999; Catuneanu et al., 2000; Boyd and Lillegraven, 2011; shorelines outside Western interior modified from
Owens and Gohn, 1985; de Cserna, 1989; Galloway et al., 1991; Blakey, 2013). Shaded areas represent land.

48 WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
through-going connection with both areas during the
late Maastrichtian either through the McKenzie
Basin or Hudson Seaway (Erickson, 1999; Ziegler
and Rowley, 1998).
As mentioned above, it has been traditionally
postulated that a complete marine withdrawal and
continental draining occurred prior to end of the
Maastrichtian and that subsequent Paleocene marine
deposition was part of a separate transgression into
the Western Interior (e.g., McGookey, 1972; Gill and
Cobban, 1973; Williams and Stelck, 1975;
Lillegraven and Ostresh, 1990). This interpretation,
however, has been questioned by recent studies that
suggest the WIS persisted through the latest
Maastrichtian and into the earliest Paleocene
(Erickson, 1999; Hoganson and Murphy, 2002;
Hartman and Kirkland, 2002; Wroblewski, 2004,
2008; Boyd and Lillegraven, 2011). There is
evidence for marginal marine condition during the
late Maastrichtian occurring as far west as eastern
Wyoming and along the Montana-North Dakota
border (Schlaikjer, 1935; Boyd and Lillegraven,
2011). These data suggest that what remained of the
WIS during the latest Maastrichtian and subsequent
Paleocene was probably located in the Great Plains
or along the cratonic side of the basin where
subsequent Cenozoic erosion has removed any
record of its existence or is covered by younger
strata and vegetation (Boyd and Lillegraven, 2011).
Additional evidence, for the presences of a
continuous seaway dividing North America is
provided by a lack of mixing of Laramidian and
Appalachian dinosaur faunas up to the K/Pg Mass
Extinction Boundary (Holtz Jr., pers. comm., 2014).
If a land connection had been made between
Laramidia and Appalachia near the end of the
Maastrichtian then a mixing of terrestrial vertebrate
faunas from the two regions should have occurred,
which has not been documented.
Lithostratigraphic as well as biostratigraphic data
on foraminifers and mollusks indicate that marine
deposition in the WIS persisted into the Danian and
Selandian ages of the early and middle Paleocene
(66.0-59.2 Ma) as the Cannonball Seaway (Figure 5;
Fox et al., 1942; Fox and Ross, 1969; Cvancara,
1976; Garvie, 2013). This age designation is also
supported by mammalian index fossils found in
terrestrial strata that directly under- and overly
marine tongues in the Cannonball Formation in
North Dakota (Anderson et al., 2006; Boyd and
Lillegraven, 2011). This relationship indicates that
marine incursions occurred during at least the
Puercan, Torrejonian, and Tiffanian Land Mammal
Ages of the lower to middle Paleocene (Figure 5;
Anderson et al., 2006; Boyd and Lillegraven, 2011).
Although the paleogeographic extent of the seaway
is poorly constrained during the Paleocene, faunal
data from the WIFB have been cited to support
expansions of both the Gulf of Mexico and the
Arctic Ocean into the Great Plains region of the
United States and southern Canada (Figure 16;
Cvancara, 1976; Boyd and Lillegraven, 2011). There
is even strong faunal evidence to suggest full
connections between the Gulf of Mexico, Arctic
Ocean, and Atlantic, although this was possibly
ephemeral (Figure 16; Cvancara and Hoganson,
1993; Boyd and Lillegraven, 2011). The location of
the western shoreline(s) of the Cannonball Seaway is
poorly known during transgression (Boyd and
Lillegraven, 2011). However, stratigraphic and fossil
evidence suggests that the sea may have extended at
least as an ephemeral water body as far west as
Saskatchewan, eastern Montana, and south Central
Wyoming (Belt et al., 1997, 2000; 2004; Catuneanu
and Sweet, 1999; Catuneanu et al., 2000; Kroeger
and Hartman, 1997; Boyd and Lillegraven, 2011).
There is some evidence to suggest that the seaway
was split into an arm of the Arctic Ocean and an arm
of the Gulf of Mexico by the Sheridan Delta in the
location of the Dakotas (Hartman, pers. comm.,
2013). As with the Campanian and Maastrichtian,
the location of the eastern shoreline is completely
unknown.
The final withdrawal of the Cannonball Seaway
from the interior of North America most likely
occurred no earlier than the middle Paleocene,
however, the timing is poorly constrained due to a
paucity of preserved strata from this age across most
of the Great Plains. Mapping upper Paleocene strata
in the Gulf Coastal Plain indicates that the Late
Paleocene shoreline closely mirrors the modern
shoreline offset by ~150 km inland (Galloway et al.,
1989; Ziegler and Rowley, 1998). This shoreline
data shows no deviations extending up into the
direction of the Western Interior for upper Paleocene
strata, which does not hold for lower to middle
Paleocene strata (Galloway et al., 1989; Ziegler and
Rowley, 1998). This shoreline data lends strong

49
WYOMING GEOLOGICAL ASSOCIATION GUIDEBOOK
support for the complete retreat of the sea from the
interior during the latest Paleocene.
CONCLUSIONS
The WIS is one of the largest post-Paleozoic
epeiric seas and covered most of west-central North
America for ~46 Mya. The seaway’s location within
this actively subsiding foreland basin and its
proximity to a continually uplifting sediment source
resulted in a thick, but complicated mosaic of
interfingering marine and terrestrial deposits that
record sea-level fluctuations within the seaway. This
thick stratigraphic record on the tectonically actively
subsiding portion of the WIFB has made it possible
through the detailed analysis of lithofacies and
fossils to paleogeographically reconstruct the
successive changes in the position of the sea's
western shoreline at a fine scale of temporal
resolution. In contrast, the evolution of the northern
and eastern shorelines remains poorly resolved due
to condensed sections, higher erosion rates, and
limited exposures along the cratonic side of the
basin.
The seaway’s well-documented paleogeography
along with its detailed lithostratigraphic,
biostratigraphic, and chronostratigraphic frameworks
make it a perfect laboratory to examine a range of
questions related to paleobiology, sequence
stratigraphy, sea level change, paleoceanography,
and paleoclimatology within a highly refined
temporal and spatial framework. These detailed
paleogeographic reconstructions will aid in
developing and testing new hypotheses concerning
evolution, faunal dynamics, and biogeography
during greenhouse climatic intervals. However, it
should also be noted that future research utilizing
paleogeographic reconstructions of the WIS must
take into account the inherent biases that are found in
the preserved or available stratigraphic record of the
WIFB.
ACKNOWLEDGEMENTS
The authors would like to thank M. Bingle-Davis,
JP Cavigelli, D. Boyd, B. Andres, N. Landman, R.
Scott, J. Hartman, J. Hoganson, J. Lillegraven, M.
Zaleha, R. Blakey, G. Plint, J. Cicarelli, and P.
Wetmore for their helpful discussions and
suggestions during the preparation of this
manuscript. The authors would also like to
acknowledge M. Bingle-Davis, B. Andres, and D.
Boyd for editing early versions of this manuscript.
Partial funding for this research has been provided
by National Science Foundation through Grant EAR
1053517 to PJH.
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