4: Getting to the bottom of the High Plains aquifer: New insights into the depositional history, stratigraphy, and paleoecology of the Cenozoic High Plains
Published:September 07, 2016
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J.J. Smith, A.L. Layzell, W.E. Lukens, M.L. Morgan, S.M. Keller, Robert A. Martin, David L. Fox, 2016. "Getting to the bottom of the High Plains aquifer: New insights into the depositional history, stratigraphy, and paleoecology of the Cenozoic High Plains", Unfolding the Geology of the West, Stephen M. Keller, Matthew L. Morgan
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This field trip is tied to a GSA 2016 joint Pardee Keynote Symposium (P5) and related topical session on the Cenozoic geology of the Great Plains (T201) that integrate new breakthroughs in scientific coring and sedimentary geoscience with the long-term demand for better characterization of the High Plains aquifer. The aquifer is the primary source of groundwater for all uses on the Great Plains, and is experiencing historic declines in water levels that could seriously compromise sustainability in areas under increasing developmental stress. Understanding the depositional history, stratigraphic framework, and hydrogeologic properties of the Cenozoic sediments comprising the aquifer is important for developing water management strategies. This guide describes a three-day field trip to examine several ongoing and increasingly collaborative projects investigating the depositional history, stratigraphy, and paleoecology of the Cenozoic High Plains. The field trip is focused on two main areas: (1) the Denver Basin, where synorogenic sedimentation associated with the Laramide orogeny deposited sediments during the late Eocene; and (2) the High Plains of western Kansas, where during the Miocene–Pliocene uplift of the Rocky Mountains, a series of clastic wedges prograding eastward formed a nearly contiguous aggradational surface covering most of the western Great Plains. Late Cenozoic uplift of the region initiated deep incision into Neogene strata and the subsequent exposure of sediments and paleosols of the Ogallala Formation. During the trip, we will visit sites and landscapes that have a long and rich history of geologic investigation. In addition to previously published results, we also present new data on the sedimentology, geochronology, paleontology, and paleoichnology of Cenozoic High Plains deposits.
The High Plains aquifer, one of the largest freshwater aquifers in the world, is the primary source of potable and irrigation waters in the central Great Plains and underpins much of the region’s economy (Fig. 1). The aquifer provides ~30% of irrigation waters for the United States and drinking water for 82% of the population living within the aquifer boundary (Dennehy, 2000). Intensive development of the aquifer for agriculture began in the 1940s, and now supports an estimated 15.5 million acres of irrigated land in the states of Kansas, Colorado, Nebraska, New Mexico, Oklahoma, South Dakota, Texas, and Wyoming (McGuire, 2014). In some areas, groundwater withdrawals greatly exceed local rates of recharge and have produced dramatic declines in water levels and growing concerns for long-term sustainability of the aquifer (Butler et al., 2015). Measured water-level declines of >60 m from predevelopment levels are common in southwestern Kansas (McGuire, 2014), where the usable lifetime of the aquifer at present rates of withdrawal ranges from less than 10 yr to greater than 250 yr (Wilson, 2007).
The aquifer is under increasing developmental stress, and predicting the effects of various groundwater management scenarios will depend on continuing improvements in the characterization of the hydrologic properties of aquifer sediments and their stratigraphic framework. These properties control hydrogeologic relationships and aquifer geometries, and were derived from the depositional histories and formative processes of water-bearing and confining strata (Anderson and Woessner, 1992). In order to better understand the depositional history of aquifer sediments, paleoclimatic and paleoecological studies are becoming increasingly important. In particular, such studies can help elucidate the relative effect of late Cenozoic tectonic and climatic forcing mechanisms on the deposition of aquifer sediments.
Acquiring good quality stratigraphic information from the aquifer has been challenging because of the complex terrestrial stratigraphy, limited surface exposures that represent only a small fraction of the total thickness of the aquifer, and the great difficulty in retrieving intact drill cores from unconsolidated and saturated intervals (Frye et al., 1956; Macfarlane, 2009). The High Plains-Ogallala Drilling Program (HPODP) at the Kansas Geological Survey is carrying out a multi-faceted scientific coring program to improve understanding of the chronostratigraphy of sediments and hydrostratigraphic architecture comprising the High Plains aquifer. Scientific drilling and outcrop investigations in western Kansas were inaugurated in 2010 with the help of grants provided by the National Science Foundation, the Kansas Water Office, and the U.S. Geological Survey. To date, 26 long continuous cores totaling over 1000 m have been collected from 10 study areas in central and western Kansas (Fig. 2) and also a site in Nebraska. Geophysical and geochemical analyses of the core, conducted mostly at the University of Kansas, have included particle size analyses; stable isotope geochemistry and hydrochemistry of organic and inorganic mineral constituents and in situ pore fluids; gamma ray spectrometry, magnetic susceptibility, and magnetostratigraphy; optically stimulated luminescence (OSL) dating; and dating of detrital zircons from selected intervals via laser-ablation inductively coupled plasma mass spectrometry (LA-ICP-MS).
The objectives of this field trip are to highlight several ongoing projects investigating the depositional history, stratigraphy, and paleoecology of the Cenozoic High Plains (Fig. 3). These projects, in large part, began independently of each other, but more recently have involved an increasing level of collaboration. The trip consists of two main areas: (1) the Denver Basin, where synorogenic sedimentation associated with the Laramide orogeny deposited sediments during the Late Cretaceous, Paleocene, and Eocene (ca. 70–34 Ma); and (2) the High Plains of western Kansas, where Miocene-Pliocene uplift of the Rocky Mountains resulted in sediments being shed eastward, forming a nearly contiguous aggradational surface that covered most of the western Great Plains. Late Cenozoic uplift of the region initiated deep incision into Neogene strata and the subsequent exposure of sediments and paleosols of the Ogallala Formation. During the trip, we will visit sites and landscapes that have a long and rich history of geologic investigation, including sites studied by John Frye and Claude Hibbard.
This field trip is being offered in conjunction with a Pardee Symposium (P5) on the High Plains aquifer and a topical session (T201) on the Cenozoic geology of the Great Plains at the 2016 Geological Society of America Annual Meeting. The material in this guide is mostly derived from our own published work, including Smith et al. (2011, 2014), Layzell et al. (2015, 2016), Keller and Morgan (2013, 2016), Fox et al. (2012a, 2012b), and Martin et al. (2000, 2008). In addition to previously published results, we also present new data on the sedimentology, geochronology, paleontology, and paleoichnology of Cenozoic High Plains deposits. We hope that this trip will facilitate important discussions and foster future collaborations to improve our understanding of Cenozoic geology in the Midcontinent and implications for sustainability of the High Plains aquifer.
This section presents a general overview of the geology of the Denver Basin and the High Plains. More specific details are presented in the background information for each stop. The Denver Basin lies east of the Front Range, extending to eastern Colorado, southeastern Wyoming, and southwestern Nebraska, and records the uplift and denudation of the central and southern Front Range of the Rocky Mountains (e.g., Thorson, 2011). The majority of deposition in the Denver Basin occurred during the Late Cretaceous, Paleocene, and Eocene (ca. 70–34 Ma), during a period of Front Range uplift and exposure. East-flowing braided stream systems carried clastic sediments from the uplift area (west) into the basin (east), and in the Castle Rock area deposited the Denver Formation (TKd) and the Dawson Arkose (Tda). Renewed uplift of the Tda, probably during middle Eocene time, caused fluvial deposition of the Larkspur Conglomerate (Tlc) in valleys of the Tda paleotopography. In the late Eocene, a volcanic eruption ~100 km to the west propelled ignimbrite of the Wall Mountain Tuff (Twm) eastward to the Castle Rock area, and draped it on the surface of the Tda and Tlc. Near the end of the Eocene, the Castle Rock Conglomerate (Tcr) was deposited mostly by southeast flow in a large paleovalley incised into the Tda, Tlc, and Twm. The Tcr is an important precursor deposit to the Miocene-Pliocene sediments comprising the High Plains aquifer, and likely is a close analog for deeply buried paleovalley fills underlying the Ogallala Formation farther to the east.
The High Plains consists of the Neogene Ogallala Formation and overlying, hydraulically connected Quaternary deposits. These units consist mostly of alluvial gravel, sand, silt, and clay, together with calcareous paleosols, and eolian silt and clay; local lenses of volcanic ash and lacustrine limestones are also present (e.g., Frye et al., 1956; Seni, 1980; Gustavson and Winkler, 1988; Diffendal et al., 1996; Gustavson, 1996). Calcic paleosols occur with high stratigraphic frequency throughout the Miocene-Pleistocene succession, and are characterized by abundant carbonate nodules, root traces, and interbedded paleophreatic calcretes. The clastic sediments were eroded and transported eastward during Miocene-Pliocene uplift of the Rocky Mountains; likely by braided, high-energy, ephemeral streams and by eolian processes (Gustavson and Winkler, 1988; Fielding et al., 2007). These sediments were deposited initially in broad paleovalleys incised into pre-Cenozoic strata by regional uplift of the Great Plains region during the Late Cretaceous to Eocene Laramide orogeny (Heller et al., 1988). The paleovalleys were eventually filled and overlain by coales-cent alluvial fans that formed a nearly contiguous aggradational surface covering most of the western Great Plains (Seni, 1980). More recent studies have highlighted differences in local rates of aggradation and uplift, as well as sedimentary provenance, suggesting that streams feeding the High Plains depositional system likely experienced different depositional histories (e.g., Chapin, 2008; Eaton, 2008). Late Cenozoic uplift of this region halted Ogallala deposition and initiated widespread erosion and deep incision of Neogene strata by the Platte, Republican, Arkansas, and Canadian Rivers that continues to the present day (Leonard, 2002).
Tectonic forces are a major mechanism controlling High Plains deposition, but the influence of paleoclimatic evolution is poorly known. Climatic changes that occurred during High Plains deposition suggest a warming and humid period in the early-middle Miocene, shifting to a cooler, arid period associated with the global spread of grasslands later in the Miocene, and, eventually, transitioning to the modern climatic regime (Zachos et al., 2001). The maximum thickness of the High Plains succession is up to several hundred meters; however, regional thickness is highly variable due to the inherited paleotopography and the more recent incision and erosion associated with late Cenozoic uplift (Leonard, 2002).
Chronostratigraphy of the Ogallala Formation and overlying units has been constrained using biostratigraphy of fossil mammal and floral assemblages and tephrochronologic analyses of unaltered volcanic air fall tuffs. Fossil mammal assemblages of the Barstovian, Clarendonian, and Hemphillian North American Land Mammal Ages (NALMAs) and biostratigraphic zonations of fossil grass seeds indicate that Ogallala deposits in Kansas range from middle Miocene to possibly earliest Pliocene in age (e.g., Thomasson, 1990; Zakrzewski, 1990; Liggett, 1997; Bartley, 2005; Martin et al., 2008). Boellstorff (1976, 1978) proposed that Pliocene strata were largely missing from central and southern High Plains Neogene strata based on a 5.0 ± 0.2–2.8 ± 0.3 Ma gap in volcanic ash bed dates; possibly due to a regional episode of increased rainfall and fluvial incision beginning in the late Miocene associated with the opening of the Gulf of California (Chapin, 2008). There are, however, local exceptions to this proposed regional unconformity. For example, volcanic ash bed chronologies, paleomagnetic data, and biostratigraphic studies indicate Pliocene-aged strata in the Meade basin (Meade County) of southern Kansas (e.g., Martin et al., 2008).
Logistics and Hazards
For the most part, the stops on this field trip do not require off-trail hiking or clambering. We will briefly depart <100 m from the trail at a few of the stops, but we will remain in open country and will not be out of sight of where we departed. There will be some non-strenuous off-trail hiking at stops 2-2 (Borchers Badlands), 2-3 (Keefe Canyon), and 3-1 (Cimarron River). Stops 2-1, 2-2, 2-3, and 3-1 are on private property and future visitors will need to ask landowners for permission for access. Rattlesnakes are occasionally encountered at Castlewood Canyon State Park (CCSP), so keep an eye on the ground there.
Day 1. Castle Rock Conglomerate and the Neogene Ogallala Formation
Day 1 of this field trip will examine the Castle Rock Conglomerate (Tcr) at Castlewood Canyon State Park (CCSP) south of Franktown, Colorado; followed by the Neogene Ogallala Formation in and around Lake Scott State Park in Ladder Creek Canyon, Scott County, Kansas.
Background: Castle Rock Conglomerate
The Castle Rock Conglomerate (Tcr) is a late Eocene fluvial deposit flanking the east side of the Colorado Front Range and lying within the Colorado Piedmont. It may reach ~70 m in thickness (Morse, 1985), is nearly flat-lying, and is discontinuous. The unit lies in a northwest-southeast-trending swath ~63 km in length and varying between 3–10 km in width. It extends from near Rueter-Hess Reservoir, ~8 km southwest of Parker, Colorado, to the buttes at Fremont Fort, ~15 km southeast of Elbert. The width of the unit increases southeastward. Near the end of the Eocene, the Tcr was deposited mostly by southeast flow in a large paleovalley incised in the Dawson Arkose (Tda), Larkspur Conglomerate (Tlc), and Wall Mountain Tuff (Twm). The Tcr incorporated numerous clasts and slump blocks of Twm, and also some rip-up clasts of Tda. The Tcr dips very gently (several meters per kilometer) northwest, in the opposite direction from the generally southeast paleocurrent direction in the unit.
Thorson (2011) described the Tcr as follows:
The Castle Rock Conglomerate is a pebble, cobble, and boulder arkosic conglomerate composed predominantly of sub-round to round fragments of pink and gray granite, pink feldspar, and quartz, with subordinate amounts of gneissic metamorphic rocks, quartzite, red sandstone, welded tuff, and chert in a coarse to very coarse sand matrix of quartz and feldspar grains. The distinguishing characteristic of this unit is the presence of angular to subangular pebble- to boulder-sized clasts of gray, brownish-gray, maroon, or lavender-gray welded tuff, which have been eroded from deposits of the Wall Mountain Tuff. Outcrops of this unit usually are very strongly cross-bedded.
Morse (1985) performed the most comprehensive study of the Tcr to date. He recognized three fluvial facies in the unit: (1) planar cross-bed facies, (2) trough cross-bed facies, and (3) massive gravel bed facies. Large-scale cross-bedding, of both the planar and trough types, is characteristic of the Tcr and is prominent in many exposures. Based upon 24 clast surveys of the gravel fraction of the upper Tcr (Keller and Morgan, 2013, 2016), the majority of clasts >2 cm are granitics (~60%), followed by Twm (~30%), and then by quartz and quartzite (each ~5%). Sedimentary clasts are rare and most of these are red-brown sandstone superficially resembling the Pennsylvanian Fountain Formation, which is exposed in the hogback along the Front Range (Evanoff, 2007; Keller and Morgan, 2013, 2016). The Tcr is younger than the Twm, as evidenced by Twm clasts in the Tcr, and the tuff has been dated at ~36.7 m.y. (Thorson, 2011). Titanothere bones in the Tcr show that it is Eocene or older (Thorson, 2011).
Field observations indicate that a large southeast-trending paleovalley was eroded into the Tda and Twm, and that a southeast-flowing fluvial system, possibly consisting of braided channels, filled the paleovalley with Tcr deposits (Morse, 1985; Evanoff, 2007; Keller and Morgan, 2013, 2016). In the area of CCSP and also southeast toward Elbert, there are two major northeast-flowing tributaries to the southeast main paleochannel belt. These are shown in Figure 4 as the JA Ranch paleochannel and the Bucks Mountain trend (Keller and Morgan, 2013, 2016; Koch, 2013a). Based upon studies of distinctive blue-gray quartzite clasts in the Tcr, the Tcr fluvial system may have emerged from the Front Range near present-day Coal Creek Canyon, between the sites of present-day Boulder and Denver (Evanoff, 2007; Koch, 2013b). The Tcr between Coal Creek Canyon and Parker has been removed due to erosion by the South Platte River and its tributaries.
After deposition of the Tcr during the late Eocene, there were periods of erosion and deposition in the southwest Denver Basin. The Miocene Ogallala Formation, a unit of alluvial and eolian sediments deposited over much of eastern Colorado, could have extended into the southwest Denver Basin and afterward been removed by erosion (Thorson, 2011). The youngest sediments in the area are unconsolidated Quaternary sands and gravels found in paleochannels, alluvial plains, and higher erosion surfaces (Thorson, 2011).
Evanoff (2007) supported the concept of a large, southeast-flowing paleovalley in which the Tcr was deposited by a southeast-flowing stream system. He noted, however, that the Tcr currently is ~300 m higher at its southeast end than at its northwest end, and proposed that this is the result of “significant northward tilting” of the southern Denver Basin during the Neogene. Thus, the Tcr appears to have been tectonically reversed to a northwest dip from its original position on a southeast-sloping paleovalley.
Stop 1-1. Castlewood Canyon State Park (Castle Rock Conglomerate) [Park Entrance GPS Coordinates: N 39°19.404’, W 104°44.086’]
Directions: From downtown Denver, take Speer Blvd. ~2 km (1 mi) west and get on I-25 south at the Speer Blvd. exit (exit 111). Take I-25 south for ~43 km (~27 mi) to the Founders Pkwy. exit (exit 184). Turn left (east) and follow Founders Park-way/Hwy. 86 for ~8 km (~5 mi) to a signal light bearing a small green sign pointing left to continue on Hwy. 86. Turn left (east) here and follow Hwy. 86 for ~8 km (~5 mi) to Franktown and the signal-light junction with Colorado Hwy. 83. From Franktown, turn right (south) on Hwy. 83 and drive for ~8 km (~5 mi) to the entrance to Castlewood Canyon State Park (CCSP), which is on the right-hand (west) side of the highway. Turn in here, follow the entrance road for ~0.8 km (0.5 mi), then turn right (north) into the parking lot for Bridge Canyon Overlook, and park. Walk the Bridge Canyon Overlook trail for ~90 m (300 ft) toward the overlook structure. Before reaching the structure, go east around the end of the rail and walk to the ledge adjacent to the structure.
The purpose of this stop is to outline the stratigraphy and bedforms of the Tcr. Across Cherry Creek from the overlook is a steep exposure of the upper portion of the Tcr, ~20 m in thickness, and here exceptional examples of cross-bedding can be observed in three dimensions. Note the jutting outcrop with a weathered hole in it. On the ledge behind and to the right of the outcrop, between the trees above and a less resistant layer below, are two layered, concave-upward features (Fig. 5A). Each is roughly 4-5 m wide and ~2 m in height, and both belong to a set of northeast-trending trough cross-beds. We are looking approximately down-paleocurrent and along the axes of these troughs. In examining trough cross-bedding, the paleocurrent direction cannot readily be discerned from vertical outcrop faces. In the present example, however, the northeast paleocurrent direction was measured from trough exposures in the roughly flat-lying outcrop along the ledge; open ends of the troughs face northeast and the trough axes lie southwest-northeast. (A reliable method of measuring paleocurrent direction is to record the trough axis azimuth while facing the open end of the trough.) Looking down and to the right, in the exposure lying between the valley floor and the highway, we can view a lower horizon also displaying trough cross-bedding. At this location, the cross-beds are characteristically truncated at the top and have gently curved, tangential contacts at the base. We are looking northeast and roughly transverse to paleocurrent flow, and the flow direction is southeast to south. Shifting our attention up-section, we can see a relatively thin horizon of nearly flat-lying planar beds occurring at the top of the ledge exposure and just left of the bridge (Fig. 5B); the beds appear to dip very gently northward. These might be either lower-phase or upper-phase flat beds, according to the bedform stability diagram of Nichols (2009). Occurrences like this are rare in the upper Tcr.
Before leaving the overlook structure, note that it is built of blocks of Wall Mountain Tuff (Twm), which a century ago was a well-known building material in the Denver area. The moderately welded Twm was erupted as an ignimbrite ca. 36.7 Ma, and originated ~100 km (~60 mi) west of here. Twm clasts are an important constituent of the gravel fraction of the Tcr, and they range from subrounded pebbles to angular boulders >0.5 m in diameter. Very large Twm fragments (as great as 1 χ 2 m in plan view) possibly are slump blocks from paleochannel banks. Walking back along the trail to the parking lot, and pausing just beyond the left-hand (east) railing, are two tuff boulders in the flat exposure of conglomerate. A little further along, and also on the left-hand side of the trail, there is a prominent cobble of the distinctive blue-gray quartzite mentioned earlier; this quartzite is believed to have been transported from Coal Creek Canyon south of Boulder.
After heading back to the parking lot, we will walk for ~60 m along the Canyon View Nature Trail to view the north-northeast-trending trough cross-beds (on the right [east] side of the trail), which may correlate with the northeast-trending horizon we viewed at the overlook. This is a good place to see the lithology and bedforms of the Tcr in plan view. Fortunately, the Tcr has many flat-lying and gently sloping occurrences like this, making possible the efficient collection and mapping of large numbers of paleocurrent measurements, as well as collection of clast data.
After returning to the vehicles and driving to the west end of the parking lot for the Lake Gulch Trail, we will walk along the trail for ~0.5 km (~0.3 mi). Approximately 300 m (1000 ft) after the trail staircase, we will view a bluff of Tcr outcrop with a spur at its western end, adjacent to the trail. Here we can view a well-exposed horizon of the main paleochannel and its southeast to south-southeast-trending trough cross-beds, which are observable in three dimensions (Fig. 5C). The outcrop at the spur suggests a fining-upward sequence as presented in Morse (1985). There is a massive, cobbly deposit at its base overlain by finer, bedded sediments (Fig. 5D); note the red-brown sandstone boulder resembling the Pennsylvanian Fountain Formation.
Morse (1985) used Tcr clast size to estimate stream velocity at 6–10 m/s (13–22 mph). An alternative approach suggested by Abbott (2016, personal commun.) is to use the method of Paola and Mohrig (1996), in which median grain size and bank-full channel depth are measured to estimate paleoslope of the channel bed. This method cannot be used for the flood deposits of the Tcr (i.e., strata containing bedforms), but potentially can be used on fining-upward sequences like the outcrop at the spur. Using this method, it is quite likely that the channel paleodepth will be underestimated, owing to the fact that the channel top will typically be removed when the subsequent channel migrates over it. Underestimating the channel depth results in an overestima-tion of paleoslope. However, this effect is countered by the fact that it is equally likely that the channel’s median grain size will be underestimated. The estimated paleoslope can be used in the Manning equation to estimate paleovelocity (Fetter, 2001). This is a possible avenue of continued research in the Tcr.
Returning back along the trail to the top of the staircase, we will leave the trail and follow the base of the Tcr ledge southeast for ~50 m (~170 ft). The ledge is a trough cross-bedded horizon in which paleocurrents were east-northeast to east-southeast. The deposits probably belong to the JA Ranch paleochannel, one of the two major northeast-flowing tributaries to the southeast-flowing main paleochannel. After returning to the vehicles, we may stop to use the rest rooms at the CCSP visitor center.
Background: Ogallala Formation
For the second part of Day 1, we will travel east to Lake Scott State Park in west-central Kansas to examine the Neogene Ogallala Formation in outcrops of Ladder Creek Canyon. Underlying much of the North American High Plains, the Ogallala Formation consists mostly of interbedded conglomerate, sandstone, mudrock, loess, their uncemented equivalents, and calcretes (e.g., Moore et al., 1944; Frye et al., 1956; Seni, 1980; Gustavson and Winkler, 1988; Diffendal et al., 1996; Gustavson, 1996; Fielding et al., 2007; Joeckel et al., 2014). Deposits of the Ogallala Formation accumulated from ca. 18 Ma to around 5.5 Ma as a series of wedges of clastic sediments shed eastward during Miocene-Pliocene uplift of the Rocky Mountains; likely by braided, high-energy, bedload-dominated streams and as eolian deposits (Fielding et al., 2007; Gustavson and Winkler, 1988). Calcareous paleosols occur with stratigraphic frequency (Gutentag et al., 1984), though individual beds often grade laterally from one lithology to another, and dramatic changes in thickness and bed continuity over relatively short distances are not uncommon (Waite, 1947). The formation is up to several hundred meters thick in parts of western Kansas, but regional thickness varies considerably due to the uneven bedrock upon which sediments were deposited and erosion associated with post-Ogallala uplift.
Biostratigraphy of fossil mammal and floral assemblages and tephrochronologic analyses of unaltered volcanic ash beds have provided some age constraints on strata of the Ogallala Formation (Fig. 6). Fossil mammal assemblages of the Barsto-vian, Clarendonian, and Hemphillian NALMAs and biostratigraphic zonations of fossil seeds indicate that Ogallala deposits in Kansas range from middle Miocene to possibly earliest Pliocene in age (e.g., Thomasson, 1990; Zakrzewski, 1990; Liggett, 1997; Bartley, 2005; Martin et al., 2008). Volcanic ash beds from the High Plains are air fall tuffs that have been sorted to fine-grained micron-scale rhyolitic glass bubble-wall shards by long-distance atmospheric transport (Rose et al., 2003). At least six such deposits in the Ogallala Formation of northern Kansas have been matched with volcanic tuffs from the Bruneau-Jarbidge, Twin Falls, and Picabo volcanic fields in the Snake River Plain in southern Idaho by chemical fingerprinting (e.g., Perkins, 1998) and five by U/Pb dating of volcanogenic zircons (Hallman et al., 2015).
Regional lithostratigraphic correlations of the Ogallala Formation are complicated due to local differences in tectonic history, basin hypsography, aggradation rates, and sedimentary provenance. In southwestern Kansas, the Ogallala Formation and undifferentiated Quaternary sediments are up to several hundred meters thick, but regional thicknesses vary greatly due to uneven bedrock surfaces associated with dissolution of underlying Permian evaporites (Young and Buddemeier, 2002) and ongoing exhumation of the central High Plains surface (Leonard, 2002). Ogallala deposits have been characterized as complex, heterogeneous, or even as random by some investigators (e.g., Waite, 1947; Gutentag et al., 1984). Sedimentary beds have been interpreted to vary widely in facies, thickness, and layer continuity over short distances. Lithofacies associations originally assigned member status in the type area (Valentine, Ash Hollow, and Kimball Members; Lugn, 1939) are recognizable in outcrops throughout Nebraska (e.g., Diffendal, 1982; Diffendal et al., 1996; Fielding et al., 2007; Joeckel et al., 2014). Though similar facies are recognized in Kansas (e.g., Moore et al., 1944; Waite, 1947; Frye et al., 1956; Thomasson, 1979), they do not appear to occur in any consistent stratigraphic order and show little regional lithostrati-graphic utility (Ludvigson et al., 2009). Similarly, the Ogallala “cap rock” in Kansas colloquially refers to a regionally persistent, thick terminal calcrete previously thought to be defining the High Plains surface. Petrocalcic horizons, however, occur in numerous stratigraphic positions in the Ogallala Formation, and the concept of a unique terminal calcrete has not been supported by more recent stratigraphic studies (Diffendal, 1982).
Stop 1-2. Lake Scott State Park, Kansas (Neogene Ogallala Formation) [GPS Coordinates: N 38.676703’, W 100.917031’]
Directions: From the entrance of Castlewood Canyon State Park, turn left (north) onto Colorado Hwy. 83 for ~7.9 km (~5 mi). At Franktown, turn right (east) onto Colorado Hwy. 86 for ~83.7 km (~52 mi). Merge right onto I-70 and travel east for ~251 km (~156 mi) to Kansas exit 70. Merge right onto U.S. Hwy. 83 and drive south for ~7.7 km (~4.8 mi). Turn left (east) onto U.S. Hwy. 40/U.S. Hwy. 83 for ~1.6 km (~1 mi), then turn right (south) onto U.S. Hwy. 83 for ~48.2 km (~30 mi). Turn right (west) onto Kansas Hwy. 95 for ~5.6 km (~3.5 mi) to the entrance of Lake Scott State Park, which is on the right-hand (west) side of the highway. Enter the park by turning right onto West Scott Lake Drive and travel ~1.3 km (~0.8 mi) to the El Quartelejo parking area. This stop in the middle of the state park affords an excellent view of the massive, ledge-forming calcretes that buttress the canyon rim, and a chance to use the nearby rest rooms in preparation for a closer examination of Ogallala outcrops during the next two stops.
Approximately 23 km north of Scott City in west-central Kansas, up to 53 m of the Neogene Ogallala Formation is exposed around Lake Scott State Park along the bluffs and in tributary draws of Ladder Creek Canyon (Fig. 7). Spring-fed Ladder Creek, formerly known as Punished Woman’s Fork or Beaver Creek, is dammed to the north to create Lake Scott. The canyon bluffs have been the focus of long-term studies on the stratigraphy, calcrete development, and paleoecology of the Ogallala Formation (e.g., Gutentag, 1988; Smith et al., 2011; Platt et al., 2012, 2015). In addition, a 19.7-m-deep core (Lake Scott core) was retrieved on the western rim of the canyon to refine our understanding of the local stratigraphy and to sample pore fluids to analyze their isotopic composition (Day 2’s “Background: HP1A Core site in Haskell County, Kansas” section for more details).
Just to the east and closer to the shore of Lake Scott sit the ruins of El Quartelejo, the only known pueblo in Kansas. This seven-room building was constructed sometime around 1696 CE by Taos and Picuris Pueblo peoples fleeing Spanish rule in New Mexico (Shearmire, 2010). Native peoples would not occupy El Quartelejo again following their forced return to New Mexico by Spanish soldiers in 1706. During the eighteenth century, Spanish and French forces alternated use of El Quartelejo as a frontier outpost as they contested dominion over the central High Plains. The site was abandoned following French retreat from the area in 1763 and largely forgotten until archaeological excavations in 1898 by S.W. Williston and H.T. Martin of the University of Kansas. The site was designated a National Historic Landmark in 1964.
Stop 1-3. Devils Backbone (Neogene Ogallala Formation) [GPS Coordinates: N 38°38.430’, W 100°54.816’]
Directions: From the entrance of Lake Scott State Park, turn right (south) onto Kansas Hwy. 95 for ~2.9 km (~1.8 mi) to a roadcut that dissects an east-west-trending ridge called the Devils Backbone. There are flat areas to pull over on the left (east) shoulder of the highway just north of and just south of the roadcut.
In one of the freshest exposures within the canyon area, ~23 m of the Ogallala Formation are visible in the Devils Backbone roadcut (Fig. 8; Smith et al., 2011). The outcrops are composed mostly of tan- to reddish brown, moderately sorted, silty, fine- to very coarse-grained beds of structureless to cross-bedded arkosic sandstone with angular to prismatic structure and resistant calcrete horizons. Minor lithologies observed in local outcrops, though not necessarily at this location, include thin beds of brown claystone with angular blocky structure and conglomeratic sandstone. These deposits are interpreted as relatively high-energy fluvial deposits and proximal flood-plain sediments on which composite soils with thick pedo-genic and phreatic calcrete horizons developed (Gardner et al., 1992). Such soils form in overbank deposits during relatively long periods of landscape stability and low sedimentation rates between major inundations of the floodplain (Gustavson and Winkler, 1988). Thick calcrete horizons imply arid to subhumid paleoclimatic conditions with low, seasonal rainfall and high evapotranspiration rates; however, very thick petrocalcic horizons (>2 m thick) in the Ladder Creek Canyon study area suggests these calcretes are polygentic and modified by later phreatic carbonate precipitation (e.g., Wright, 2007).
Calcium carbonate pervades the section as fine-grained cement, but also as powdery stringers along cracks, rounded cobble-sized and smaller nodules, pipy concretions, and rhizoliths and burrow casts (Smith et al., 2011). Calcic paleosols are characterized by resistant horizons showing multiple developmental stages including laminar fenestral stringers, honeycomb calcretes, platy to massive calcretes, and prismatic calcretes (Fig. 9; Platt et al., 2012). Many calcretes contain pockets of friable, grain-supported silty sand, suggesting that carbonate precipitation was displacive and was initiated in voids likely created by biotic activity. Laminated and fenestral stringers are mm-thin planes of powdery to solid carbonate that penetrates and accentuates bedding and cracks (Fig. 9A). These calcretes are intimately associated with rhizoliths and many are interpreted as root mats—an r/z-horizon in the terms of Wright et al. (1995). Honeycomb calcretes and massive to platy calcretes are found near the tops of paleosol profiles and are interpreted as soil Bk-horizons. Honeycomb structures are composed of well-cemented, amalgamated networks of trace fossil assemblages exposed by the weathering of less indurated, interstitial sandstone matrix (Fig. 9B). Massive calcretes, the most prominent ledge-former in the canyon, appear to have originated as pedogenic calcretes, but were likely subsequently overprinted by pervasive carbonate cementation under phreatic conditions (Fig. 9C). Prismatic calcretes are found at the top of the measured section, are not associated with trace fossils, and are likely the pet-rocalcic products of more recent surface processes (Fig. 9D).
Honeycomb calcrete structures have been attributed to algal structures, rhizoliths, coprolites, or as recent dissolution features (e.g., Barbour, 1897; Schultz, 1942; Gutentag, 1988; Retallack, 1990). These structures are preserved within and below calcrete beds as carbonate-filled casts weathering in full relief, or are composed of powdery carbonate showing no relief. The stratigraphic placement and weathering patterns of calcretes in Ladder Creek Canyon, the honeycomb structures in particular, appear to be controlled by a diverse assemblage of carbonate-cemented trace fossils (Fig. 10; Platt et al., 2015). These ichnofossils include the lattice-like Daimoniobarax isp. interpreted as ant nests (Figs. 10A-10C); flask-shaped Cel-liforma isp. interpreted as bee cells (Figs. 10D-10F); largeand small-diameter vertical to subhorizontal tubes interpreted as mammal burrows; vertically flattened subhorizontal tubes interpreted as scorpion burrows; curved, finger-shaped tubes interpreted as fossorial vertebrate burrows; and a variety of rhi-zoliths (Smith et al., 2011; Platt et al., 2012). The vertical distribution of traces within individual paleosols suggests relatively long periods of landscape stability, and is tiered in a manner reminiscent of modern soil communities, i.e., there is a greater concentration of traces toward the tops of paleosols (Platt et al., 2015). In size, basic structural elements, and architectural morphology, Daimoniobarax isp. most closely resembles the nests of extant seed-harvester ants such as the New World genus Pogonomyrmex (Smith et al., 2011). Such ants are highly specialized granivores and prefer arid to semi-arid regions and sandy soils (Whitford and Jackson, 2007). Weathered chambers of this trace fossil often contain mass accumulations of fossil seeds and other vegetative structures of angiosperms (Thomasson, 2003, 2005, 2009), including grasses (Poaceae), borages (Boraginaceae), sedges (Cyperaceae), and hackberries (Ulma-ceae). Fossil plant taxa recovered suggest subhumid to subtropical savanna conditions without periods of freezing weather and with higher annual precipitation than experienced in this region currently (Thomasson, 1990).
Stop 1-4. Battle Canyon [GPS Coordinates: N 38°38.610’, W 100°55.668’]
Directions: Turn right (west) onto the dirt road immediately south of the roadcut that runs parallel with the Devils Backbone ridge and drive ~0.5 km (~0.3 mi). There is a small wooden kiosk with visitor information. At the kiosk, turn right over the cattle grid and drive on the dirt road for ~0.97 km (~0.6 mi) to the Battle Canyon monument. This stop is technically on private property; however, the current landowner has given visitors open access to the monument area.
This monument marks the site of the Battle of Punished Woman’s Fork, the last armed conflict of the American Indian Wars in Kansas between the Northern Cheyenne and the U.S. Cavalry (Fig. 11A). Most of this site is on private property, although the monument is maintained by the Scott County Historical Society. In 1878, Chiefs Dull Knife and Little Wolf led a group of Northern Cheyenne that included 92 warriors, 120 women, and 141 children out of a reservation near Fort El Reno, Oklahoma, in an attempt to return to their home in Montana Territory (Maddux, 2003). The commander at Fort Dodge, Lieutenant Colonel William H. Lewis, was dispatched to capture and return them. In late September, the Cheyenne took refuge in the canyon lands of Punished Woman’s Fork (modern Ladder Creek) to replenish supplies and construct defensive positions to meet the advancing cavalry troops. Their strategically positioned rifle pits are still visible on the hilltops surrounding the monument (Fig. 11B). On 27 September, the U.S. Cavalry and Cheyenne warriors clashed through the day, while the women and children hid in a natural rock shelter called Squaw’s Den Cave at the head of Battle Canyon (Fig. 11C). The Cheyenne escaped the cavalry on the evening of the 27th, continuing to travel to the northwest in what has become known as the Cheyenne raid. Lieutenant Colonel Lewis died from a thigh wound in the days following the battle, becoming the last Kansas military casualty of the Indian Wars.
The monument is situated stratigraphically near the top of the section measured at the Devils Backbone roadcut, and provides a closer view of natural exposures of the Ogallala Formation in the study area. While carbonate-rich strata are common in Ogallala sediments, the calcretes in Battle Canyon and other natural draws and bluffs of Ladder Creek Canyon appear thicker (>2 m) and more massive (less structured) than those exposed in the Devils Backbone roadcut. Thick calcretes are only rarely encountered in core of the Ogallala Formation and do not appear representative of petrocalcic horizons in most of the formation’s subsurface. Such massive units suggest a polygenetic formation—initially pedogenic, but now preserved as possibly phre-atic calcretes thickened and indurated by precipitation of calcite from solution at or just below the water table (e.g., Wright, 2007). Localized massive calcrete development is likely the result of phreatic carbonate precipitation at or below a slowly falling and fluctuating water table concomitant with the incision of Ladder Creek, and more recent case hardening of the outcrop face in some cases. Low-lying hills to the north and west are mantled by Quaternary loess deposits that thicken to over 8 m within a few kilometers of the canyon rim.
An ~1-m-thick lens of grayish-white, re-worked volcanic ash crops out on the north side of Landon Draw on private property, ~1.4 km (~0.9 mi) south of Battle Canyon (Fig. 11D). The ash is stratigraphically located ~2 m from the base of the Devils Backbone measured section, and has a maximum depositional age of 11.41 ± 0.39 Ma based on a U-Pb age of the youngest single zircon analyzed (Brian Sitek, 2014, personal commun.). In addition, fragmentary vertebrate fossil material—including molars, limb elements, and an ankle bone—were collected from the ash and identified as Calippus sp. (Larry Martin, 20 June 2012, personal commun.). Calippus was a common late Miocene (15-7.5 Ma) equid with several recognized Barstovian to Hemp-hillian NALMA species (MacFadden and Hulbert, 1988).
Day 2. Pliocene and Pleistocene Localities of Meade County, Kansas
The first half of Day 2 will be a brief stop at the HP1A core site in Haskell County, Kansas. For the second part of Day 2, we will visit Pliocene and Pleistocene localities in Meade County, followed by an opportunity to examine selected intervals of the HP1A core and tour the Meade Historical Museum and Dalton Gang Hideout in Meade, Kansas, courtesy of the Meade County Historical Society.
Background: HP1A Core Site in Haskell County, Kansas
HP1A is the first and deepest intact core of the High Plains aquifer ever attempted (HP1A core; Fig. 2). Previous investigations of High Plains subsurface sediments in southwestern Kansas were on drill cuttings rather than intact core segments (McMahon et al., 2003, 2004). The Kansas Geological Survey (KGS) commenced coring operations at the HP1A site on 12 April 2011. Part of the rationale for choosing this location was to complement ongoing groundwater investigations in the area. The nearby Haskell County index well has collected a continuous record of water level and barometric pressure measurements since 2007 (Fig. 12; Butler et al., 2015). In addition, this is an area of high-density agricultural water use—drawdowns of 40 m (130 ft) or more have been recorded during the irrigation season—and water levels appear to respond as if there are two separate aquifer units: a lower semi-confined aquifer and an upper unconfined aquifer (Butler et al., 2013). Drilling and core retrieval were accomplished using an Acker AD-II hollow-stem auger with a split spoon core sampler in the unsaturated zone, and an S-27 rotary-vibratory drill rig with an Aqualock piston sampling system core barrel for sediments below the water table. Approximately 98 m of core was retrieved with ~89% recovery.
Discrete interval sampling of pore waters from the HP1A core provided some of the first pristine groundwater samples to evaluate potential sources and timing of aquifer recharge. Colleagues at the University of Kansas and Kansas Geological Survey recently conducted isotopic analyses (52H and 518O) of pore fluids from five cores (including Lake Scott core [Stop 1-2], HP1A [Stop 2-1], and Borchers core [Stop 2-2]) collected along a north-south transect in western Kansas, to assess patterns of recharge to the High Plains aquifer (e.g., Katz et al., 2015; Stotler et al., 2015). Nearly all samples plotted near the global meteoric water line, indicating that waters had not been affected by evaporation and therefore, little, if any, water within the aquifer was irrigation return water. Isotopic values also increase from south to north across the state, which could be related to variations in average annual temperatures or different water sources within the aquifer in Kansas (Fig. 13). Radiocarbon dating of groundwater yielded ages of 11,474 and 4223 14C yr B.P. for samples from southwestern and northwestern Kansas, respectively (Stotler et al., 2012). The younger age from northwestern Kansas indicates that modern recharge is mixing with older water in this part of the High Plains aquifer.
Stop 2-1. HP1A Core Site in Haskell County, Kansas [GPS Coordinates: N 37°39.409’, W 100°39.877’]
Directions: From the Devils Backbone roadcut, drive south on Kansas Hwy. 95 ~2.1 km (~1.3 mi) to the intersection with U.S. Hwy. 83. Turn right (south) on U.S. Hwy. 83 and drive ~16.1 km (~10 mi) to Scott City, Kansas. Continue south on U.S. Hwy. 83 ~55 km (~34 mi), then turn left (east) to merge onto U.S. Hwy. 400 East/U.S. Hwy. 50 East/U.S. Hwy. 83 South ramp and continue to follow U.S. Hwy 83 South for ~44 km (~27.3 mi). Turn left (east) onto County Road 70 and drive for ~18 km mi), then turn left (north) onto County Road XX and follow it for ~0.64 km (~0.4 mi). The HP1A core site is located in the field to the east on private property (permission must be obtained before accessing it)—the core itself is housed at the Kansas Geological Survey in Lawrence, Kansas.
The HP1A core is composed of predominantly uncemented coarse-grained sand and gravel interbedded with fine-grained sand and silt beds, showing varying degrees of pedogenic modification (Fig. 14; Smith et al., 2014). Pedogenic carbonate in the form of rhizoliths and nodules is common throughout the core, while indurated calcrete beds are generally thin (<0.5 m thick) and only stratigraphically common in the upper 35 m of the core. Based on lithofacies and grain-size differences, the core can be divided into four sections (Harlow, 2013). Section 1 at the base of the core (98-86 m) consists of approximately equal portions of very fine- and medium-grained sand beds and sandy silt and clay with low chroma colors and redoximorphic features, suggesting poorly drained soil conditions. Section 2 (86-37 m) is the thickest and coarsest-grained interval, dominated by very coarse-grained sand and gravels with abrupt lower contacts and with very few fine-grained intervals (Figs. 15A-15D). Sedimentary features include armored mud balls, rare carbonate nodules, rare root traces, and common cobble clasts. Section 3 (37-12 m) is similar to the first in that it consists dominantly of interbedded fine- to medium-grained sands and calcareous silt beds. Paleo-sols in this section, however, suggest better drainage conditions and increased pedogenic modification, with higher chroma colors and upsection increases in carbonate nodules, calcrete layers, root traces, bioturbation, and clay skins on paleosol ped faces (Figs. 15E-15F). Section 4 (12-0 m) is composed mostly of silt and clay-dominated strata. Six units were identified and strati-graphically correlated to a similar core (CMC core, Cimarron River; Fig. 2), collected from an upland setting just north of the Cimarron River valley, as well as to more regionally extensive loess units (see Layzell et al., 2016). Optically stimulated luminescence (OSL) dating of six samples from the upper 12 m of the HP1A core yielded ages ranging from ca. 76.8 ka at ~12 m depth to 44.3 ka at ~4 m depth (Fig. 14). OSL ages from the CMC core ranged from ca. 84.0 ka at 18.7 m depth to 52.2 ka at 4.4 m depth. In addition, two radiocarbon samples from soil organic matter at 3.2 and 2.8 m depth in the CMC core yielded ages of ca. 34.6 and 29.2 ka, respectively. Chronostratigraphic relationships indicate that (1) loess and eolian sands were episodically deposited between ca. 84-70 ka, (2) the Sangamon paleosol formed from ca. 70-52 ka, and (3) soil development in the Gil-man Canyon Formation began at ca. 44 ka and continued until at least 29.2 ka. This overall chronology is consistent with evidence for late Loveland Loess aggradation (ca. 90-74 ka) and renewed Sangamon pedogenesis (ca. 74-54 ka) in the Mississippi River valley (Markewich et al., 2011), as well as the range of published ages for the Gilman Canyon Formation (ca. 45-24 ka) from the central Great Plains (e.g., Johnson et al., 2007).
Below the late Quaternary loess sequence (>12 m), the succession of lithofacies suggests a dominantly high-energy fluvial system, possibly that of a prograding megafan (Harlow, 2013). This type of depositional model has long been proposed for the High Plains (e.g., Seni, 1980; Diffendal et al., 1996; Gustavson and Winkler, 1988, Gustavson, 1996). Interbedded fine-grained to medium-grained sands and gleyed hydromorphic paleosols in section 1 suggest low sediment aggradation by a dominantly suspended load river system and poorly drained overbank conditions. Down-gradient increases in water content of soils and depositional environments due to a groundwater spring line at the leading edge of the fan are a distinctive characteristic of many megafans (Weissmann et al., 2010). As a megafan progrades, less saturated and thicker, coarser-grained facies build out over the wetter depositional settings if a spring line is present. Cobble clasts, armored mud balls, and a lack of fine-grained overbank deposits suggest that the depositional environment of section 2 was a high-energy, bed-load-dominated fluvial system. The abrupt superposition of hydromorphic strata in section 1 by these coarse sands and gravels supports the interpretation of a prograding megafan with a discrete spring line (Harlow, 2013). Section 3 transitions from a bed-load-dominated to a suspended-load-dominated fluvial system with a more stable landscape on which formed stratigraphically common, relatively mature, cumulic paleosols. This may represent a pause in basin subsidence or a downdip progradation of the bed-load-dominated megafan system and establishment of an alluvial plain.
Zircons from five fine-grained intervals were sampled below the Quaternary section and analyzed for U/Pb ages by LA-ICP-MS at the University of Kansas Geochronology Laboratory under the direction of Professor Andreas Möller. Volcanic ash beds are an excellent source for datable zircon grains and are common in High Plains deposits (Hallman et al., 2015; Turner et al., 2015), though such beds are typically of a very limited geographic extent (Ludvigson et al., 2009, and references therein). Mature paleosols, as condensed time-rich terrestrial sections in overbank facies, are likely to incorporate air fall zircons via piping through root channels, desiccation cracks, and animal burrowing activity and, thus, preserve a time-averaged record of air fall tephra events. Maximum depositional age (MDA) based on the weighted mean of the U-Pb ages of the youngest cluster of three or more overlapping concordant zircon grains at ~16 m and ~24 m intervals were 27.9 ± 1.3 Ma and 27.9 ± 0.7 Ma, respectively (Fig. 14). The MDA of samples from ~34 m and ~52 m were 35.3 ± 0.3 Ma and 35.6 ± 0.6 Ma, while zircons from ~87 m yielded an MDA of 36.2 ± 1.4 Ma. Importantly, all of these five MDAs from the HP1A core occur in a serial order that obeys the stratigraphic law of superposition. Also of note, Miocene age grains are absent. The lack of Miocene age grains is consistent with 13 previous samples from HP1A and four other cores from southwest Kansas, now totaling over 3000 zircons grains analyzed. Zircons that match the MDAs from southwestern Kansas (ca. 35-28 Ma) likely originate from explosive volcanism associated with the subduction-related mid-Tertiary ignimbrite flare-up (36-18 Ma) in western and southwestern North America, which blanketed large parts of Nevada, Utah, Arizona, Colorado, New Mexico, and northwestern Mexico in vast ash-flow tuffs and fallout deposits (Best et al., 2013). The absence of middle Miocene and younger zircons from any of the cores in southwestern Kansas is striking, given that such grains, likely derived from the Snake River Plain volcanic province (16.1-0.6 Ma), are readily identified in cores and outcrops of the Ogallala Formation in northern Kansas and Nebraska using the same techniques (Field et al., 2015; Hallman et al., 2015). Taken on their face value, MDAs in southwestern Kansas suggest Eocene to Oligocene age deposits equivalent to the White River Group in Nebraska; these ages are previously unknown from Cenozoic strata in Kansas. The lack of Miocene zircons, however, may alternatively indicate that putative Neogene (?) sediments in southwestern Kansas were derived from completely different paleo-drainage areas than those in northern Kansas and Nebraska (e.g., Fan et al., 2011, 2015).
Background: The Meade Basin: Miocene-Recent Paleoenvironments in Southwest Kansas
The Meade Basin of southwest Kansas consists of a number of highly productive fossil localities that have been the focus of paleontological study since the early to mid-twentieth century. Fluvial facies of the Miocene Ogallala Formation are exposed at the base of several Meade Basin sections, though most areas contain fluvial and lacustrine deposits from the Pliocene to Recent. No single stratigraphic section covers the entirety of geologic time preserved in the Meade Basin. Rather, numerous individual sections spread across a moderately constrained region (~50 χ 150 km, Fig. 16 inset) have been pieced together using a combination of tephro-, litho-, bio-, magneto-, and chronostratigraphy (see Martin et al., 2008).
Paleontologic and stratigraphic investigations in the Meade Basin were founded on the productive careers of Claude Hib-bard (e.g., Hibbard, 1938, 1949, 1951, 1955, 1967) and John Frye (e.g., Frye and Hibbard, 1941; Frye, 1942; Frye and Schoff, 1942; Frye and Leonard, 1952). Hibbard’s work focused extensively on mammalian communities from the middle Pliocene to Recent, and more recent efforts have examined rodent communities over the same interval due to their abundance in the fossil record (Martin and Fairbanks, 1999; Martin and Peláez-Cam-pomanes, 2014). This line of research culminated in the Meade Basin Rodent Project (MBRP), which began in 1997 and sought to address questions regarding the occurrence of populations and the drivers in evolution and community change (e.g., Marcolini and Martin, 2008). Since the inception of the MBRP, many of Hibbard’s quarries were relocated and re-excavated and, with the addition of new fossil localities, a modern database of rodent communities from the Pliocene to Recent has been amassed (Martin et al., 2000; Martin and Peláez-Campomanes, 2014). The MBRP continues today, currently as part of a multi-investigator, multi-disciplinary project funded by the Earth-Life Transitions track of the Sedimentary Geology and Paleobiology program at the U.S. National Science Foundation. The goals of the current iteration of the project are to: (1) compile a comprehensive history of rodent communities from the Pliocene to Recent in southwest Kansas, including community composition and diet among members; (2) reconstruct climate in the same interval from paleosol geochemistry, pedogenic carbonate isotopes (standard and clumped), and occurrence of temperature-sensitive taxa; and (3) reconstruct vegetation in the same interval using isotopic methods in paleo-sol carbonates, organic matter, compound-specific isotopes, and tooth enamel.
We will visit a selection of localities that have historical paleontological significance and showcase the MBRP’s latest results in the understanding of the sedimentology, paleontology, paleoecology, and paleoclimatology of the Pliocene-Recent strata of the Meade Basin.
Stop 2-2. Borchers Badlands [GPS Coordinates: N 37°10.109’, W 100°22.149’]
Directions: From the HP1A core site, drive south on County Road XX ~0.65 km (~0.4 mi) and turn left (east) onto County Road 70. After ~1.6 km (~1 mi), County Road 70 turns into Y Road, continue driving for ~27 km (~17 mi) on Y Road. Turn right (south) onto Kansas Hwy. 23 and continue for ~40 km (~25 mi) to the city of Meade, Kansas. Continue south of Meade for ~13 km (~8 mi); where Kansas Hwy. 23 makes a sharp turn right (west), drive another 2 km (1.3 mi) west to the shoulder on the south side of Kansas Hwy. 23. This stop is on private property; permission must be obtained before accessing it.
The outcrops on the south side of Hwy. 23 east of Crooked Creek are referred to as Borchers Badlands (Figs. 17 and 18). A composite stratigraphic section, including paleosol features and recent stable isotope analyses of pedogenic carbonates is shown in Figure 18. An ~28 m core was collected from the north side of Hwy. 23 to refine our understanding of the local stratigraphy, and to collect pore fluid samples for isotopic analyses (see discussion in Day 2’s “Background: HP1A Core Site in Haskell County, Kansas” for more details). The Ogallala Formation crops out at the base of the section, forming a bluff at the margin of the flood-plain of Crooked Creek. Age control for the Ogallala Formation at this locality is poor, given by a limited mammalian fauna of Clarendonian NALMA (12.5-9.0 Ma; Zakrzewski, 1988; Tedford et al., 2004). The Ogallala Formation is relatively coarsegrained at this locality, consisting of stacked, cross-bedded gravels of braided stream channels (Figs. 19A-19B). A good view of these deposits may be found in the nearby gravel pit.
The Ogallala Formation is capped by an unconformity of unconstrained duration and is overlain by the Pliocene-Pleistocene Crooked Creek Formation (Hibbard, 1949; Martin et al., 2008). At Borchers Badlands, the Stump Arroyo Member (?early Pleistocene) forms the base of the Crooked Creek Formation and consists of orange clayey sandstone with common gravel. This unit in the Borchers Badlands may correlate with sediments in the Seger gravel pit in Meade County referred to as the Stump Arroyo Member, from which a few large mammals have been recovered (Hibbard, 1951). Martin et al. (2008) tentatively placed the Seger gravels in the early Pleistocene. Paleosols of the Stump Arroyo Member in the Borchers Badlands consist of coarse-grained, massive units with varying degrees of carbonate accumulation, clay translocation, and slickenside development (Figs. 19C-19D), consistent with high-energy stream deposits subaerially exposed in semiarid to dry subhumid environments.
Gray-green clayey sandstones and mudstones of the Pleistocene Atwater Member overlie the Stump Arroyo. Age control is given by the Huckleberry Ridge Ash (HRA, 2.113-2.135 Ma; Ellis et al., 2012), which crops out in the middle of the Atwa-ter Member in Borchers Badlands, and the Cerro Toledo B Ash (1.23-1.47 Ma; Izett and Honey, 1995), which crops out in laterally correlative strata on the north side of Hwy. 23 (the Nope! section of Martin et al., 2008; Fig. 16 inset). The Atwater Member marks a gradual change in hydrologic regime from the bed-load-dominated, braided stream deposits of the underlying Stump Arroyo Member to a suspended load fluvial system (Figs. 19D-19F). Stacked paleosols, commonly capped by erosional surfaces, formed on distal floodplain deposits and emergent lacustrine clays. The HRA plugs a fluvial channel (Fig. 19E), and was later reworked into overbank paleo-Vertisols (Figs. 19F-19G).
Paleosols of the Atwater Member below the HRA are generally coarser-grained, clayey sandstones, and contain abundant carbonate nodules, occasional calcretes, calcareous rhizoliths, drab halo root traces, and iron mottles (Fig. 19E). These features are consistent with semiarid moisture regimes in proximal to medial floodplain environments. The finer-grained paleosols above the HRA show a marked transition to distal floodplain and emergent lacustrine clays (Fig. 19F). Carbonate nodules are less abundant and pedogenic slickensides are common, due to either an increase in precipitation, decrease in temperature, or simply a shift in depositional environment. Groundwater-derived carbonate seams are also present and cross-cut bedding features (Fig. 19H). The Margaret, Borchers, Short Haul, and Nash 72 fossil accumulations (close to the Aries localities in Fig. 16 inset) were likely initially deposited in shallow ponded environments before subsequent pedogenic modification. It should be noted that Margaret and Short Haul accumulations are located within ~3 km (~1.9 m) to the north of Hwy. 23 in laterally equivalent, clay-rich strata.
The fossil accumulations in the Atwater Member are renowned for rich rodent assemblages that have been studied since the early twentieth century (reviewed by Zakrzewski, 1975; Martin et al., 2008). Changes in faunal assemblages, particularly rodent communities, show evidence of climatic cooling associated with the onset of Pleistocene glaciation, which include: (1) significant dwarfing then extinction of the diminutive cotton rat (Sigmodon minor;Peláez-Campomanes and Martin, 2005) in the Borchers site, (2) dramatic body size increase then extinction of the gopher Geomys quinni (Martin et al., 2011), and (3) disappearance of the giant tortoise Hesperotestudo (Bayne, 1976). Additionally, a significant faunal turnover occurred just after deposition of the HRA (40% turnover of the entire rodent community, including the Microtus immigration event onto the Central Great Plains; Martin et al., 2008).
Carbon isotopic signatures of pedogenic carbonates were used to estimate the abundance of above-ground C3 versus C4 biomass (Fig. 18) following the methods of Fox et al. (2012a). Vegetation using the C3 photosynthetic pathway fractionates atmospheric 13C to a greater extent than C4 vegetation, resulting in disparate isotopic signatures between the two groups: -25.7%e and -10.8%c on average for C3 and C4 Pliocene-Pleistocene vegetation, given a pre-industrial atmospheric 513C of -6.5%e. Using a simple linear mixing model, the abundance of C4 vegetation was found to consistently vary between 40%-60% throughout the Borchers stratigraphic section (Fig. 18). Younger deposits in the Meade Basin reveal a gradual increase in C4 grass toward the ~78% C4 biomass of modern Oklahoma and Kansas (Martin et al., 2008; Fox et al., 2012a).
Oxygen isotopic signatures of pedogenic carbonates are influenced by the 518O of meteoric water and the temperature of carbonate precipitation, where greater values indicate warmer, drier conditions and less winter recharge. Several extreme, positive values of 518O are present at the onset of the Atwater Member (Fig. 18), either due to a change in local hydrology through increased evapotranspiration, or through localized diagenesis of carbonates. Overall, the decrease in values from ~-25%o to ~-22%o across the Pliocene-Pleistocene boundary is interpreted to reflect the termination of the warm Pliocene climate and the onset of major Northern Hemisphere glaciation (Haug et al., 2005; Martin et al., 2008).
Stop 2-3. Keefe Canyon [GPS Coordinates: N 37°2.307’, W 100°34.457’]
Directions: From the Borchers Badlands stop, continue west and then south on Kansas Hwy. 23 for ~30.6 km (~19 mi). Turn right (west) onto FF Road, a dirt road, and drive for ~8.5 km (~5.3 mi). Turn right (north) onto 5 Road for ~3.5 km (~2.2 mi), then turn right (east) onto Windmill Road. After ~1.3 km (~0.8 mi) on Windmill Road, take the first right, an unnamed road, and drive south for ~2.4 km (~1.5 mi). This stop is on private property and visitors must get landowner permission to access the outcrops in the canyons.
Numerous Pliocene-Recent localities exist north of the Cimarron River in canyons eroding into an escarpment of the Ogallala Formation (Honey et al., 2005). Of particular significance are Hibbard and Riggs’ (1949) locality (Loc.) 22 (also known as “Keefe Canyon” in the literature) and Raptor 1C, which have produced the most diverse rodent assemblages. They are part of a series of important early Pliocene vertebrate assemblages recovered from sediments along the north bank of the Cimarron River in Meade County (Honey et al., 2005; Martin et al., 2008), all of which appear to be older than another set of classic Pliocene assemblages farther north in Meade County, the latter including Rexroad Loc. 3 and Deer Park (Martin et al., 2002).
Because all of the sediments tested in the Cimarron River sequence have generated only reversed paleomagnetic signatures (Lindsay et al., 1975; Martin et al., 2008), we must rely on biostratigraphic correlations for the placement of fossil sites in and around Keefe Canyon. Index species and correlational logic were provided by Marcolini and Martin (2008), with critical taxa including the early Pliocene archaic arvicolid Ogmodontomys sawrockensis (an extinct vole), recovered from sites older than those in Keefe Canyon, such as Saw Rock Canyon and Fallen Angel B in Seward County, and its likely descendant O. poapha-gus from younger sites such as Raptor 1C and Loc. 22. Based on comparisons with other early Pliocene arvicolids, Marcolini and Martin (2008) suggested a range in time of 4.90-4.63 Ma for assemblages with O. sawrockensis. Another useful species for correlational purposes is the cotton rat Sigmodon holocuspis, which was originally described from paleomagnetically reversed sediments between the Cochiti and Nunivak Subchrons of the Gilbert Chron in the Verde Formation of Arizona (Czaplewski, 1987). S. holocuspis has only been recovered from Keefe Canyon (Loc. 22) and Raptor 1C, and is tentatively allocated in Kansas to the period between 4.48 and 4.29 Ma.
The main outcrop on the west side of Keefe Canyon (Fig. 20) consists of stacked fluvial channels and weakly developed paleosols (Fig. 21). These were previously referred to as the “Rexroad Formation,” but Honey et al. (2005) refrained from applying a formal stratigraphic name to sediments along the north bank of the Cimarron River due to unresolved correlations to the Rexroad Formation type section to the north. The fluvial channels are mixed energy, dominated by sand, but with gravel lags and internal fining-upward sequences capped by paleosols. Significant quantities of calcium carbonate are present in the outcrop, though some of the carbonates are associated with fluvial sandstones and are potentially either phreatic in origin or the result of modern case hardening. Paleosol analysis is ongoing.
Carbonates recovered from Keefe Canyon and correlative early Pliocene/Blancan localities nearby and south of the Cimarron River have mean δ13C signatures of -4.9 ± 5.7%c VPDB (Vienna Peedee belemnite) (n = 106), consistent with 20.158.5% C4 biomass (Fox et al., 2012a). This interval precedes a monotonic rise in C4 grass abundance based on pedogenic carbonate 513C values in the Borchers Badlands and the correlative outcrops to the north of Hwy. 23, which culminated in the modern ~78% C4 biomass grassland in the region (Martin et al., 2008; Fox et al., 2012a).
Day 3. Cimarron River Valley and Return to Denver
Day 3 of the field trip includes travel back to the Colorado Convention Center with one stop to examine alluvial fills in the Cimarron River valley. The Cimarron River originates on the High Plains in northeastern New Mexico and joins the Arkansas River near Tulsa, Oklahoma. A detailed stratigraphic, geomor-phic, and paleoenvironmental study was conducted at this stop by Layzell et al. (2015). Ongoing investigations at the University of Kansas and Kansas Geological Survey are seeking to ascertain whether alluvial fills, including those in the Cimarron River valley, function as areas of focused recharge to the High Plains aquifer.
Stop 3-1. Cimarron River [GPS Coordinates: N 37°25.0’, W101°3.50’]
Directions: From Meade, Kansas, head west on U.S. Hwy. 160/U.S. Hwy. 54 for ~22.5 km (~14 mi), then turn right onto U.S. Hwy. 160 for ~26.9 km (~16.7 mi). Turn right (north) onto U.S. Hwy. 160/U.S. Hwy. 83 and drive for ~11.3 km (7 mi). Turn left (west) onto Kansas Hwy. 190 for ~17.4 km (~10.8 mi). Turn left (west) onto U.S. Hwy. 56 at the City of Satanta, Kansas, and drive for ~8.4 km (~5.2 mi). Turn right onto County Road CC; the quarry stops are all on private property along County Road CC (Fig. 22).
This stop consists of exposures of alluvial valley fills in three sand and gravel quarries on the north side of the Cimarron River in Haskell County (Fig. 22). The lower two quarries (quarries 1 and 2) expose valley fills underlying two late Quaternary terraces (T-1 and T-2; Fig. 23), whereas quarry 3 exposes sediments tentatively assigned to the Neogene Ogallala Formation. Multiple paleosols are exposed in each of the quarries, all indicating periods of landscape stability or slow sedimentation in the Cimarron River valley.
This quarry is located ~1.5 km north of the Cimarron River. We will stop for a panoramic view of the quarry and the adjacent landscape before descending to observe the stratified sandy and gravelly alluvium exposed in the pit walls (Fig. 24). We pose the following question at this stop: Does the alluvium exposed in this pit represent the Ogallala Formation? In southwestern Kansas, ascertaining the difference between Neogene (or older?) and Pleistocene sediments is hindered because the lithologies of different-aged sediments are commonly virtually identical.
As we descend into the quarry, we will stop to observe an ~3.5-m-thick pedocomplex, consisting of two welded paleosols (Fig. 24B). Both paleosols are well developed with thick Btk horizons, prismatic structure, silty clay loam textures, and common to many rounded carbonate nodules. The lower paleosol is more strongly indurated with carbonate. Here we will pose the question: Does this pedocomplex correlate with the paleosols of the Stump Arroyo Member observed in the Borchers Badlands at stop 2-2? Although these paleosols are morphologically similar, Al/Be cosmogenic nuclide dating is currently being undertaken to determine the burial age of these paleosols and to shed further light on this question.
Exposures in this quarry are of the T-2 fill, which primarily consists of vertically stacked channel-fill sequences that fine upward and range in texture from fine sand to fine gravel (Fig. 23). Note the horizontal bedding and trough cross-stratification evident in outcrop. Sedimentological characteristics indicate deposition by a high-energy braided stream system and OSL ages indicate that this depositional environment prevailed in the study area during late Marine Isotope Stage (MIS) 5 and MIS 4 (ca. 77.0-58.6 ka). Three paleosols occur in the T-2 fill and are best observed at the north wall of the quarry. Of particular interest is paleosol S1, which has cumulic properties with a 1.25-m-thick, moderately developed Ak-Bk profile. A suite of radiocarbon (14C) and OSL dates indicates that this paleosol formed during MIS 3 (ca. 48-28 ka). Paleoenvironmental reconstructions from stable carbon isotope analyses (513C) of soil organic matter show increasing 513C values from -17.8%c at the bottom of the paleosol to -14.6%c VPDB at the top, which indicates the presence of a C4-dominated environment (up to 89% C4 biomass) during paleo-sol formation. Temperature reconstructions (after Nordt et al., 2007) indicate a mean July temperature of 25.0 °C (513C value of -14.6%c VPDB) at ca. 28 ka, which is similar to modern July temperatures in the study area.
As we travel from quarry 2 to quarry 1, note the amount of relief (~15 m) between the T-2 and T-1 surfaces (Fig. 23A). Layzell et al. (2015) inferred that the Cimarron River incised over 25 m either shortly after 28 ka or perhaps as early as ca. 50 ka. This estimate is based on the difference between the terrace treads and assumes that the river incised at least below the 28 ka OSL sample interval, at 10 m depth in the T-1 fill (if not the 53 ka OSL sample interval at 17 m depth) before subsequently aggrading (Fig. 23). It is debatable whether climatic forcing alone could cause this magnitude of observed valley deepening, and therefore, the authors consider the effect of local base-level fall on fluvial systems, which often results in channel incision and terrace generation through the creation and upstream migration of knickpoints. In the Cimarron River valley, evidence exists for local base-level control in the Meade Basin. The Meade Basin is believed to have been formed by the coalescence of dissolution-subsidence depressions that became integrated during the late Pleistocene (Frye and Schoff, 1942; Frye, 1950). Layzell et al. (2015) hypothesized, therefore, that local base-level lowering, produced by drainage integration during the late Pleistocene, resulted in significant headward valley deepening in the Cimarron River valley, and caused the abandonment of the T-2 surface either shortly after ca. 28 ka or possibly around ca. 50 ka.
The portion of the T-1 fill exposed in this small sand pit has sandy loam, loamy sand, and sand textures in which five paleosols have developed (Fig. 23B). Paleosols typically have Btk horizons, prismatic structure, and silty clay loam to silty loam textures. Of particular note is paleosol S4, which has cumulic properties with a 50-cm-thick, organic rich Ak horizon and a 1.15-m-thick Btk horizon. 14C dates indicate that this paleosol formed between ca. 13-12.5 ka, and 513C data indicate the presence of a C4-dom-inated environment (513C value of -16.3%o VPDB, up to 77% C4 biomass) at this time. Temperature reconstructions indicate a mean July temperature of 23.8 °C at ca. 12.5 ka, which is ~2.2 °C cooler than modern July temperatures in the study area.
We gratefully acknowledge support from the National Science Foundation Grant EAR-1023285, Kansas Water Office Grant KAN0066607, and the U.S. Geological Survey National Cooperative Geologic Mapping STATEMAP program under the guidance of the Kansas Geological Mapping Advisory Committee. We thank Greg Ludvigson (Kansas Geological Survey), Jesse Korus (University of Nebraska-Lincoln) and Dennis Terry (Temple University) for their technical reviews and thoughtful critiques of the manuscript. Special thanks to Jay and Jarvis Garetson of Garetson Brothers Farms and the staff of the KGS Exploration Services. We thank colleagues Rolfe Mandel, Andreas Möller, Matt Joeckel, John Doveton, Joe Thomasson, Tammy Rittenour, Brian Platt, Randy Stotler, Laura Murphy, Mike Petronis, and Kathryn Snell; and research assistants R. Hunter Harlow, Adrienne Duarte, Brian Sitek, Elijah Turner, and the staff of the KGS Geoarchaeology and Paleoenvironment Laboratory. We are also deeply indebted to the many landowners who have not only graciously allowed us access to sites on their property but often have shown a keen interest in our research findings.
Figures & Tables
Unfolding the Geology of the West
Prepared in conjunction with the 2016 GSA Annual Meeting in Denver, Colorado, this volume contains sixteen guides to field trips in this rich geologic region. The four “Great Surveys” of the late 1800s ventured west to explore and document the region’s unknown natural resources and collect valuable geologic information. Many of the field guides in this volume, aptly titled Unfolding the Geology of the West, will cover the same hallowed ground as the early geologic expeditions. Organized into four sections, this volume spans some of the major subdisciplines of geology: (1) stratigraphy, sedimentology, and paleontology; (2) structure and metamorphism; (3) Quaternary landscape evolution; and (4) engineering and environmental geology.