Middle to Late Pleistocene loess record in eastern Nebraska, USA, and implications for the unique nature of Oxygen Isotope Stage 2

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Middle to Late Pleistocene loess record in eastern Nebraska, USA, and implications for the unique nature of Oxygen Isotope Stage 2 J.A. Masona,, R.M. Joeckelb, E.A. Bettis IIIc a

Department of Geography, University of Wisconsin-Madison, 160 Science Hall, 550 N. Park St., Madison, WI 53706, USA Conservation and Survey Division, School of Natural Resources, University of Nebraska-Lincoln, Lincoln, NE 68588-0517, USA c Department of Geoscience, University of Iowa, Iowa City, IA 52242-1319, USA

b

Received 10 March 2006; received in revised form 11 September 2006; accepted 5 October 2006

Abstract New subsurface data reveal a nearly continuous stratigraphic record of Middle to Late Pleistocene loess sedimentation preserved beneath upland summits in eastern Nebraska, USA. Thickness and grain size trends, as well as pedologic evidence, indicate significant changes in loess sources, accumulation rates, and depositional environments. The newly defined Kennard Formation accumulated in the Middle Pleistocene, and may represent multiple thin increments of distal loess from nonglacial sources on the Great Plains. The overlying Loveland Loess, up to 18 m thick and deposited during Oxygen Isotope Stage 6 (OIS 6) (Illinoian glaciation), probably records the emergence of the Missouri River valley as a major glaciogenic loess source. The prominent Sangamon Geosol formed through long-term pedogenic alteration of the upper Loveland Loess during OIS 5 and 4. Thin loess of the Gilman Canyon Formation records slow loess accumulation and pedogenic alteration in OIS 3. The Peoria Loess (OIS 2) is similar in thickness to Loveland Loess, but may have accumulated more rapidly in an environment less favorable to bioturbation. More importantly, comparison of Peoria and Loveland loess thickness trends indicates much greater influx of nonglaciogenic loess from the Great Plains during OIS 2 than in OIS 6, suggesting colder and/or drier conditions in the Midcontinent during OIS 2 than in earlier glacial stages. r 2007 Elsevier Ltd. All rights reserved.

1. Introduction Loess deposits of North American mid-continent are a key terrestrial record of environmental change (Bettis et al., 2003). The especially thick loess of the Great Plains and Missouri River valley was derived in part from glaciofluvial sources, such as rivers draining the Laurentide Ice Sheet, but a significant portion of it is nonglaciogenic dust derived from semi-arid or periglacial landscapes (Mason et al., 1994; Aleinikoff et al., 1999; Muhs and Bettis, 2000; Mason, 2001a; Bettis et al., 2003). Two characteristics appear to distinguish the Midcontinent loess from other major Late Cenozoic loess sequences: (a) little preservation of loess older than the Late Middle Pleistocene (subdivisions of the Pleistocene used here follow Richmond and Fullerton (1986)); and (b) an apparent upward transition toward much thicker, more Corresponding author.

E-mail address: [email protected] (J.A. Mason). 0277-3791/$ - see front matter r 2007 Elsevier Ltd. All rights reserved. doi:10.1016/j.quascirev.2006.10.007

widely preserved, and less pedogenically altered loess units. Only three loess units are widely recognized across the North American Midcontinent: (1) Late Middle Pleistocene Loveland Loess ( ¼ Loveland Formation, Loveland Silt); (2) a thin lower Late Pleistocene unit referred to as the Gilman Canyon Formation (Nebraska, Kansas), the Pisgah Formation (Iowa), and the Roxana Silt (Mississippi and Ohio valleys); and (3) upper Late Pleistocene Peoria Loess ( ¼ Peoria Formation, Peoria Silt) (Bettis, 1990; Leigh and Knox, 1994; Muhs et al., 1999a; Grimley et al., 2003). Peoria Loess is by far the most extensively exposed of these units, and is typically much thicker than any underlying loess. Pre-Loveland loesses have been reported at only a small number of localities east of the Missouri River, where they are pedogenically altered throughout (Porter and Bishop, 1990; Jacobs and Knox, 1994; Markewich et al., 1998; Grimley et al., 2003). Multiple Middle Pleistocene loess units are present on the Great Plains (e.g. Feng et al., 1994), but have not been studied in detail. Much of the total thickness of Loveland Loess is

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pedogenically altered, as is the entire lower part of Late Pleistocene loess (Gilman Canyon Formation or equivalents), except at a few sites near the Missouri and Mississippi rivers (Bettis, 1990; Markewich et al., 1998; Grimley et al., 2003). In contrast, thick Peoria Loess is typically massive and largely unweathered, except for the surface soil formed in it. Currently available dating methods do not provide the resolution needed to directly compare accumulation rates within the mid-continent loess sequence, but Peoria Loess clearly was deposited at exceptionally high rates (Roberts et al., 2003). Other loess sequences around the world accumulated over much more time, and display much less extreme longterm changes in thickness and pedogenic alteration. In Central Asia, Central Europe, Alaska, and the Palouse region of the northwestern US, loess sequences began accumulating before the Brunhes–Matuyama magnetic reversal at 790 ka (Kukla, 1975; Busacca, 1989; Westgate et al., 1990; Dodonov and Baiguzina, 1995). In northern China, the loess record extends into the Miocene (Liu and Ding, 1998). The thickness of individual loess-paleosol packages in the Chinese Loess, generally attributed to glacial–interglacial cycles, increases from the Early to Late Pleistocene (Liu, 1988; Liu and Ding, 1998), but this transition is not comparable to the order-of-magnitude change in depositional unit thickness between Peoria Loess and pre-Loveland loess. Possible hypotheses to explain the distinctive characteristics of loess in central North America include: (1) a Middle to Late Pleistocene acceleration of loess accumulation, at least when averaged over full glacial–interglacial cycles, because of greatly increased dust production and transport during the Late Pleistocene; and (2) unusually effective and widespread post-depositional erosion, which has clearly truncated the loess sequence in many parts of the Midcontinent (Ruhe and Cady, 1967; Leigh and Knox, 1994; Jacobs et al., 1997). If the first hypothesis is valid, the loess sequence may record previously unrecognized contrasts between successive glacial–interglacial cycles in the mid-continent, which should be explored using other paleoenvironmental data. The second hypothesis implies a need to identify more complete loess sections. In this paper, we address these two hypotheses using extensive new observations of the loess stratigraphy in eastern Nebraska. This study area received both glaciogenic loess from the Missouri River valley and nonglaciogenic loess from sources on the Great Plains. A key aspect of this study is the incorporation of abundant data from cores, other drill holes, and geophysical logs, which has significantly changed our understanding and interpretation of the loess succession. 2. Study area The study area is the region surrounding Omaha, Nebraska, USA (Fig. 1b), with emphasis on broad interfluves between the Missouri River and Platte River

valleys (Area A, Fig. 1b). The study area, at the eastern edge of the Great Plains, has a subhumid climate (mean annual temperature ¼ 10–11 1C; mean annual precipitation ¼ 750–850 mm yr1). Unconsolidated Quaternary sediment up to 120 m thick overlies Pennsylvanian and Cretaceous bedrock. Relief is generally10–30 m km2, locally reaching 100 m km2. The study area is fluvially dissected, with a fully integrated drainage network. Despite this dissection, loess on wide ridgetops is often saturated below a depth of 5–10 m. The study area was glaciated during the Pleistocene, although no glacial landforms are preserved. Work by Boellstorff (1978a, b) indicated that the most recent glaciation occurred after 780 ka (the Brunhes–Matuyama paleomagnetic boundary), but before about 640 ka (deposition of the Lava Creek B tephra from the Yellowstone volcanic center). Roy et al. (2004a) revised Boellstorff’s stratigraphic framework, but also found no evidence of glaciation in eastern Nebraska after the Lava Creek B eruption. 3. Background: loess stratigraphy and chronology Loess has received much less attention in the study area than in nearby southwestern Iowa, the location of numerous studies of loess chronology, sedimentology, pedology, and paleopedology (Simonson and Hutton, 1954; Ruhe, 1969; Ruhe et al., 1971; Muhs and Bettis, 2000). Pleistocene loess units formally recognized in Nebraska include Peoria Loess, Gilman Canyon Formation, Loveland Loess and several older formations that include both loesses and fluvial deposits (Reed and Dreeszen, 1965). Paleosols appear throughout this sequence. Loess units older than Loveland Loess have occasionally been reported from the study area, but were poorly understood (Boellstorff, 1978a; Mandel and Bettis, 1995). The Loveland Loess is 5–7.3 m thick in its paratype section (Loveland Section, Fig. 1b), in Iowa just across the Missouri River valley from the study area (Daniels and Handy, 1959; Bettis, 1990). Ruhe and Cady (1967) demonstrated that Loveland Loess thins systematically eastward from the Missouri River valley. Thermoluminescence (TL) ages from Loveland Loess at the paratype section range from 165 to 125 ka (Forman et al., 1992), while newer infrared-stimulated luminescence (IRSL) results range from 165 to 146 ka (Forman and Pierson, 2002). These ages are broadly consistent with age estimates based on TL and 10Be accumulation in the lower Mississippi Valley and western Nebraska (Maat and Johnson, 1996; Markewich et al., 1998), and with the traditional assignment of Loveland Loess to the penultimate glaciation (Oxygen Isotope Stage 6 (OIS 6)). The Sangamon Geosol, the most prominent paleosol in the loess succession, developed into Loveland Loess across the Midcontinent. This paleosol commonly has a more strongly expressed Bt horizon, a thicker solum, and a redder hue than surface

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Fig. 1. (a) Location of study area (star) in relation to thick Peoria Loess distribution (Bettis et al., 2003; some small areas of thick loess not shown) and ice sheet margins during OIS 2. (b) Shaded relief map of part of eastern Nebraska, with locations of continuous cores (J) or mud-rotary testholes (K) used in this study. The study focused on Area A, but the larger Area B (whole map) was also used in evaluating regional thickness patterns at two scales. (c) Detailed view of Area A, indicating locations of cores (J) or testholes (K) mentioned in text and shown in Figs. 3–5.

soils that formed in loess during the Holocene (Ruhe, 1965; Follmer, 1983). Development of the Sangamon Geosol spanned OIS 5 and much of OIS 4 (Curry and Pavich, 1996; Markewich et al., 1998). The Gilman Canyon Formation is only 0.5–1 m thick in most outcrops in the study area, but the stratigraphically equivalent Pisgah Formation is 44 m thick at the Loveland Section. The Gilman Canyon Formation and Pisgah Formation have been radiocarbon dated at many localities in Nebraska and western Iowa. Most ages, from soil organic matter, fall between 41,000 and 20,000 14C yr BP (Souders and Kuzila, 1990; Johnson, 1993; Martin, 1993; May and Holen, 1993; Mandel and Bettis, 1995; Maat and Johnson, 1996; Muhs et al., 1999b); approximately 45,000–25,000 cal yr BP, based on the calibration proposed by Fairbanks et al. (2005). Ages from charcoal at the

Loveland Section, however, suggest that the Pisgah Formation was accumulating by about 45,000 14C yr BP (about 48,000–50,000 cal yr BP; Fairbanks et al. (2005)), consistent with IRSL dating (Forman and Pierson, 2002). The Peoria Loess can be traced almost continuously from northeastern Colorado to Ohio (Bettis et al., 2003). Peoria Loess thickness exceeds 30 m in a few outcrops in southwestern Iowa near the Missouri River (Bettis, 1990; Bettis et al., 2003). Numerous calibrated radiocarbon and luminescence ages indicate that most Peoria Loess was deposited between 25 and 11 ka (largely in OIS 2) (Bettis et al., 2003); however, the specific timing of deposition within that interval varied across the mid-continent, and in some areas most Peoria Loess accumulation may have occurred within a few thousand years (e.g. Roberts et al., 2003). In parts of Nebraska and Kansas, Bignell Loess was

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deposited during the Holocene, and is distinguished from Peoria Loess by the intervening Brady Soil (Johnson and Willey, 2000; Mason et al., 2003b). Miller (1964) described Bignell Loess in sections near Omaha, but the Brady Soil is not recognizable at those sites and the sediment interpreted as Bignell Loess is most likely upper Peoria Loess. It has generally been assumed that the Missouri River floodplain was the main source of loess in both the study area and adjacent southwestern Iowa, with a possible contribution from the valley shared by the Platte and Elkhorn rivers (Fig. 1b) (Miller, 1964). At the time of Peoria Loess deposition, the Missouri River drained the Laurentide Ice Sheet upstream of the study area during the last glaciation (Fig. 1a) and the Platte River system drained alpine glaciers in the Rocky Mountains, but the Elkhorn River had no glacial input. Aleinikoff et al. (1998) presented new data from geochemical tracers, indicating that Peoria Loess across much of Nebraska was derived from non-glacial sources on the western Great Plains. Mason (2001a) reached similar conclusions using Peoria Loess thickness trends, noting that the valleys of the Missouri River, the Platte River, and the Elkhorn River were all minor loess sources, detectable only over distances o60 km. Muhs and Bettis (2000) reported geochemical and grain size evidence for contributions from both the Missouri Valley and western sources to Peoria Loess at Loveland, Iowa. Both sources probably contributed to Peoria Loess in southwestern Iowa, which thins and fines eastward from the Missouri (Simonson and Hutton, 1954; Ruhe et al., 1967; Muhs and Bettis, 2000). 4. Methods We investigated loess stratigraphy, sedimentology, and paleopedology primarily on broad, nearly level upland summits, using cores and mud-rotary drilling. Observations were also made in outcrops and a few testholes on hillslopes. Nine continuous cores (5.1 cm diameter) were collected on upland summits between the Missouri and Platte/Elkhorn river valleys north of Omaha (Fig. 1b and c), using hollow-stem auger drilling or hydraulic directpush soil coring. A few older cores archived at the Conservation and Survey Division (CSD), University of Nebraska-Lincoln were also used in this study. Core recovery was 90–100% for most 1.5-m drives; greater loss occurred in a few cases, especially in low-cohesion lower Peoria and Loveland loess. We also used data from many mud-rotary test holes on upland summits in eastern Nebraska, drilled for this study or earlier geologic investigations by CSD (Fig. 1b and c). Mud-rotary drilling techniques used by CSD allow placement of lithologic boundaries within 70.5–1 m, with poor preservation of pedologic features (Mason, 2001a). Geophysical logs from the mud-rotary testholes provided additional data on lithological boundaries and grainsize variation. Single point resistance (SPR; one of several resistivity-based logging methods) logs were available from

all drilling sites, and natural gamma logs from sites drilled for this study. Higher SPR corresponds to coarser grain size, lower water content, and/or buried soil horizons with strong granular structure (Mason, 2001a). Natural gamma values increase with the abundance of radioisotopes of K, U, and Th, usually related to clay content. In exposures on hillslopes, we described both the vertical and horizontal variation of major lithologic and pedogenic horizons. We also made use of detailed hillslope crosssections prepared by USDA Soil Conservation Service geologists (unpublished data on file at CSD, University of Nebraska-Lincoln). Segments of three cores (10-B-98, 3-B-99, and 7-B-00, Fig. 1c) were sent to Spectrum Petrographics (Vancouver, WA) for preparation of thin sections (30 mm thick, 2 cm  3 cm), using impregnation with clear resin under vacuum. We attempted to collect samples from major horizons of the Sangamon Geosol, and at regular intervals elsewhere in the cores, but sample spacing was often controlled by the need to select intact, minimally disturbed core segments. Pedogenic features in the thin sections were described using the terminology of Bullock et al. (1985). Grain size analysis was carried out on samples from three cores (Stevens, 10-B-98, 7-B-00, Fig. 1c) using a Malvern Mastersizer 2000 laser diffraction particle size analyzer after sonication in deionized water for 3 min, with no other pretreatment. Experiments using 46 loess samples from across Nebraska and 12 samples collected near Omaha showed this methodology yielded clay and fine silt contents at least as high as those measured after a full set of chemical pretreatments (10% HCl, 30% H2O2, and Na-metaphosphate dispersion). Laser diffraction analysis after the 3-min sonication also produced results that were highly correlated with pipette/sieve analyses carried out on the same samples after full chemical pretreatment (HCl/H2O2/Na-metaphosphate) (Fig. 2). As in previous studies (Beuselinck et al., 1998; Mason et al., 2003a), laser diffraction measures much less clay than pipette analysis, although there is a strong linear relationship between results from these two methods (Fig. 2). 5. Results 5.1. Loess stratigraphy on upland summits Cores and test holes on broad, nearly level upland summits consistently reveal the same sequence of depositional units (Fig. 3 and see Table in the online version of this article). We assign the oldest sediments in this sequence to a new lithostratigraphic unit, the Kennard Formation, distinguishable from overlying Loveland Loess by greater clay and fine silt content, stronger pedogenic structure, and higher SPR. The three units above the Kennard Formation are correlated with existing lithostratigraphic units: Loveland Loess, Gilman Canyon Formation, and Peoria Loess. The four major loess units were differentiated in the field, on the basis of color, texture, and pedostratigraphy.

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Fig. 2. Relationship between %o2 mm as measured by laser diffraction (Malvern Mastersizer 2000, 3-min sonication) and pipette method (full chemical pretreatment). Open circles represent samples of loess and loess-derived soils from across Nebraska; gray squares are samples of loess from the study area.

Descriptions of three representative cores (online Table) illustrate typical lithological and pedological characteristics of the loess succession. 5.1.1. Kennard Formation (new) 5.1.1.1. Description. The type section of the Kennard Formation is Core 3-B-99 (411290 N, 961130 W, Figs. 1c and 3 and online Table), collected near the village of Kennard, Nebraska, and archived by the CSD, University of Nebraska-Lincoln. This formation has a total thickness of 6–11 m on wide ridgetops, is overlain by Loveland Loess, and usually rests directly on the uppermost preIllinoian glacial diamicton. At the Martin-Marietta quarry north of Springfield, Nebraska (41140 N, 96140 W), the Kennard Formation instead overlies fine-grained laminated sediment, probably lacustrine, containing an undated volcanic ash bed. No tephras have been observed within or above the Kennard Formation. The Kennard Formation is always fine-grained, ranging from clay or silty clay to silty clay loam. Rare to common granules and pebbles are often observed in the basal meter, but not in the rest of this unit. Geophysical logs indicate that the Kennard Formation has very low resistivity, consistent with high clay content (Fig. 4). Highly variable resistivity and natural gamma radiation logs within this unit suggest significant variations in clay content or mineralogy (Fig. 4). Moderate to strong pedogenic structure (especially fine to medium blocky) is present throughout much of the thickness of the Kennard Formation; root traces, clay coatings, and a wide variety of redoximorphic features are also common (online Table). Granular structure is clearly distinguishable in a few horizons. However, discrete A–B–C soil profiles are rarely identifiable, except at the base of the unit, where a clearly identifiable soil profile is developed through a gradational

contact with underlying diamicton, in 3-B-99 (online Table) and other cores. The contact of the Kennard Formation with underlying glacial diamicton is generally marked by a large downward increase in coarse sand and gravel content, and a downward increase in SPR. In cores, we identify the upper boundary, with Loveland Loess, on the basis of an upward decrease in clay content and weakening of pedogenic blocky structure. This boundary is gradational and difficult to place exactly in some cores, but it is clearly identifiable in SPR logs from mud-rotary testholes (Fig. 4). Natural gamma logs do not clearly differentiate the Kennard Formation from Loveland Loess, possibly because mineralogical differences offset differences in clay content. 5.1.1.2. Interpretation. We interpret the Kennard Formation as a stack of two to three thin depositional units, primarily fine-grained loess, although the basal 1–2 m of the unit probably contains some depression-fill, pond sediment, and local colluvium. The scattered coarse sand grains and pebbles in the basal part of this formation could have been mixed upward into loess by bio- or cryoturbation, or they could have been washed into depressions. The sediment was strongly overprinted by pedogenesis, resulting in such effective welding of overlapping soil profiles that they are no longer recognizable as discrete entities. 5.1.1.3. Correlation and age. The Kennard Formation has often been labeled ‘‘Sappa Formation’’ in older test hole logs but no clear correlation can be established to the type Sappa section in distant south-central Nebraska (Reed and Dreeszen, 1965). Possibly correlative units include the informally named Yarmouth Clay of southwestern Iowa (Ruhe and Cady, 1967) and the Ferrelview Formation of northern Missouri (Howe and Heim, 1968); both have been

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Fig. 3. Stratigraphic columns illustrating loess stratigraphy on wide, gently sloping upland summits of the study area, focusing on the intensively sampled region shown in Fig. 1c. Site locations shown in Fig. 1c. Gray shades approximately represent light-dark color variation in cores, in which more pedogenically altered zones are usually darker-colored.

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only the basal Kennard Formation (2 m), followed by slow or episodic accumulation of the rest of the unit, would also yield a reasonable fit with the data, however (Greg Balco, personal communication, 2005). The latter interpretation would be more consistent with the degree of pedologic evidence suggesting either slow accumulation or rapid accumulation in several increments.

Fig. 4. Geophysical logs from a mud-rotary testhole in the study area (location in Fig. 1c), illustrating typical log response to loess unit lithology. Lines separate stratigraphic units; gray band indicates Sangamon Geosol.

interpreted as primarily lacustrine sediment, but are similar to the Kennard Formation in lithology and stratigraphic position. The Kennard Formation probably correlates with some of the poorly understood pre-Loveland loess on the Great Plains (e.g. Feng et al., 1994), in southern Iowa (Woida and Thompson, 1993), and in the Mississippi Valley (e.g. Porter and Bishop, 1990; Jacobs et al., 1997; Markewich et al., 1998; Grimley et al., 2003). Much more work will be needed to make more definite correlations, however. Only a very approximate age estimate can be made for the Kennard Formation, but it most likely accumulated in the Middle Pleistocene, beginning after 780 ka and ending before OIS 6, when some or all of the overlying Loveland Loess was deposited. We did not find glacial sediment overlying the Kennard Formation at any locality. We cannot rule out removal of overlying glacial sediment by erosion, but it seems implausible that such widespread erosion would not also have removed the Kennard Formation or truncated the overlying loess sequence at some sites; thus, it is likely that the Kennard Formation postdates the most recent glaciation of the region. As discussed above, the most recent glaciation of eastern Nebraska probably occurred between 780 and 640 ka (Boellstorff, 1978a, b; Roy et al., 2004a), placing the onset of Kennard Formation deposition at some time after 780 ka. Consistent with this stratigraphic evidence, Balco et al. (2005) estimated the age of the basal Kennard Formation in Core 3-B-99 as 5807120 ka, using cosmogenic nuclide burial dating. This estimate was based on the assumption that the entire Kennard Formation was deposited rapidly (duration of deposition 103 or 104 yr, rather than 105 yr), which provides the best fit with the cosmogenic nuclide data (Balco et al., 2005). Rapid accumulation of

5.1.2. Loveland loess 5.1.2.1. Description. Loveland Loess is 6–16 m thick in continuous cores from upland summits (Fig. 3 and online Table), often exceeding the thickness of this unit at its paratype section (Daniels and Handy, 1959; Bettis, 1990) and any other known outcrops. High-quality geologic logs from mud-rotary testholes, supported by archived samples, indicate that Loveland Loess is up to 28 m thick on the bluffs bordering the Missouri Valley. In cores from summits in the study area, Sangamon Geosol is always present in the upper part of the Loveland Loess (Fig. 3 and online Table). The upper solum of this paleosol has well-preserved granular to very fine subangular blocky structure, which we interpret as A, AB, or BA horizons that have lost most organic carbon through postburial oxidation. The most strongly expressed Bt horizons of the Sangamon Geosol are usually reddish brown (7.5YR to 5YR hue) silty clay loam to clay, with strong subangular blocky structure and abundant clay coatings on ped faces (online Table). In geophysical logs, the Sangamon Bt is typically marked by relatively high natural gamma radiation, and low SPR (Fig. 4). Although the reddish brown color of the Sangamon Bt may in part reflect formation of iron oxide coatings on ped faces after burial (Thompson and Soukup, 1990), pedogenic rubification clearly played an important role, because ped interiors and clay coatings in the Bt both have redder hues than underlying less-altered Loveland Loess. In one core (2-B-98, Fig. 3), the Sangamon Geosol Bt, along with the rest of the Loveland Loess, has a gray color interpreted as the result of iron reduction in a poorly drained swale. Below the Sangamon Geosol profile, two zones can be distinguished within thick Loveland Loess on uplands of the study area (Fig. 3 and online Table). The upper zone of Loveland Loess is light-colored, generally yellowish brown silty clay loam or silt loam, typically with weak to moderate blocky structure. The upper zone of Loveland Loess can superficially resemble Peoria Loess, but often contains sparse clay coatings in larger pores or along fractures, unlike the Peoria (online Table). Within the upper zone, there are usually darker-colored bands with strong fine blocky structure that appear to be incipient A horizons (they are not pedogenic clay lamellae). One of these is often observed immediately below the Sangamon Geosol Bt or BC horizon. The lower zone of Loveland Loess is marked by a distinctly darker-colored horizon at its upper boundary, and the zone as a whole is generally slightly darker and redder than most of the upper zone. Within the lower zone, color value varies significantly, in

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alternating darker and lighter horizons. The two zones of Loveland Loess are difficult to distinguish, however, where the unit is less than about 5–7 m thick, and has a brown or reddish brown color and blocky structure throughout. 5.1.2.2. Interpretation. The Loveland Loess is interpreted as essentially a single depositional unit, because it contains no well-developed paleosols indicative of a long hiatus in loess deposition, below the Sangamon Geosol. The latter developed in the uppermost Loveland Loess after loess accumulation had ceased or slowed to a very low rate. The dark bands (incipient buried A horizons) in the lower zone of Loveland Loess suggest at least short intervals of slower accumulation and/or more effective pedogenesis, however. The lower zone of Loveland Loess as a whole may have experienced more pedogenic alteration as it accumulated, but the color difference between zones (especially the redder hue in the lower zone) may also reflect changing provenance. 5.1.2.3. Correlation and age. Loveland Loess north of Omaha is clearly correlative with the same unit at its nearby paratype section , where luminescence ages indicate deposition in OIS 6 (Forman et al., 1992; Forman and Pierson, 2002). The lower zone of Loveland Loess is not apparent in the paratype section, and could conceivably be older than the ages obtained there. The absence of a welldeveloped soil developed into the lower zone, however, suggests that it was not deposited long before the upper zone, and probably also accumulated in OIS 6. The two zones identified here have not been explicitly described in the study area , but the ‘‘early Sangamon (?) or Illinoian soil’’ described by Thorp et al. (1951) at the Yankee Hill Brickyard near Lincoln, Nebraska, may correspond to the lower Loveland Loess in the Omaha area. Grimley (1996, pp. 84–85) noted a darker-colored lower zone in thick Loveland Silt sections along the Mississippi River valley. Here, we designate Core 3-B-99 (Figs. 1c and 3 and online Table) as a reference section for Loveland Loess, representing both the unusual thickness and the well-defined zonation of this unit in the study area. 5.1.3. Gilman Canyon Formation 5.1.3.1. Description. In ridgetop cores from the study area, this unit is usually thin (o1 m), although it is almost 6 m thick in one core on the west side of the Missouri Valley (3-B-98, Fig. 3). The Gilman Canyon Formation is brown or dark brown silt loam to silty clay loam. The uppermost part of this unit is often macroscopically massive, but granular or fine blocky structure is increasingly well-expressed with depth. The lower boundary of the Gilman Canyon Formation is indefinite at most sites, because clay content and structural grade increase gradually with depth into the upper solum of the Sangamon Geosol (online Table). In many geophysical logs from eastern Nebraska, the Gilman Canyon Formation is marked by a distinct SPR peak (Fig. 4), probably because

of strong aggregation that creates a response similar to that of sand (Mason, 2001a). 5.1.3.2. Interpretation. We interpret the pedogenic alteration of the lower Gilman Canyon Formation as upward growth of the Sangamon Geosol under conditions of slow loess accumulation. As the Gilman Canyon Formation accumulated, the accumulation rate increased and/or the rates of pedogenic processes decreased, leading to the upward trend of decreasing alteration. 5.1.3.3. Correlation and age. The Gilman Canyon Formation can be confidently correlated in loess sections across much of Nebraska. In core 2-B-98 (Figs. 1c and 3), the Gilman Canyon Formation was deposited in a poorly drained swale on a broad ridgetop. At this site, the unit has very high organic matter content, and wood fragments at 0.6 m below the top of the unit yielded an age of 24,010790 14C yr BP (Beta-119107), consistent with ages obtained elsewhere. 5.1.4. Peoria Loess 5.1.4.1. Description. Peoria Loess caps all upland ridgetops in the study area, and is at least 30 m thick in core 4-A-76 (Fig. 1c; Boellstorff, 1978a) on the Missouri River bluffline (the base of the Peoria Loess is unusually difficult to identify in that core). Elsewhere, Peoria Loess ranges from 21 to 8 m thick on broad upland summits (Fig. 3). As elsewhere across the Great Plains and Midwest, Peoria Loess in eastern Nebraska is light yellowish brown silt loam or silty clay loam with limited macroscopic evidence of pedogenic alteration below the surface soil profile (online Table). Portions of the Peoria Loess do display weak blocky structure, rare insect and rodent burrows, and common root traces. ‘‘Dark bands’’ (weak buried A horizons), which appear in a few very thick Peoria Loess sections in southwestern Iowa (Ruhe et al., 1971) were not observed in the study area. Millimeter- to centimeter-scale subhorizontal color banding or lamination is visible in the lower Peoria Loess at virtually all sites, generally faint, but occasionally accentuated by iron stains (online Table). Macroscopic observations do not provide clear evidence that this banding has a sedimentary origin, such as distinct grain size variation within individual bands or laminae, but alternative explanations are difficult to identify. Clearly identifiable wind-ripple lamination was observed in one outcrop of sandy Peoria Loess immediately west of the Missouri valley, and in the Loveland section in Iowa (Bettis, 1990). 5.1.4.2. Interpretation. The minor degree of pedogenic alteration observed in Peoria Loess below the surface soil is interpreted primarily as evidence of rapid deposition, consistent with geochronological data (e.g. Roberts et al., 2003). Assuming that the weak lamination in this unit is

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sedimentary, its preservation could also be explained by rapid deposition, relative to the rate of bioturbation. 5.1.4.3. Correlation and age. No new radiocarbon ages were obtained from Peoria Loess in the study area, but it can clearly be correlated to well-dated loess sections elsewhere in the Great Plains and Midwestern US. Based on that correlation, it was probably deposited some time within the interval of 25–12 ka (OIS 2). 5.2. Loess stratigraphy on hillslopes Outcrops and cores on hillslopes provide important additional insight on loess stratigraphy. All four stratigraphic units observed on summits also crop out on hillslopes or narrow spur ridge summits, but all are truncated to varying degrees in that setting (Mason, 2001b). On the steepest backslopes, all loess has been removed by erosion and glacial till is at the surface. Thin Peoria Loess rests directly on glacial till on some steep slopes. Throughout the study area, the reddish-brown Sangamon Geosol is exposed in bands across hillslopes. In a few large pits, it is clear that the Sangamon Geosol marks sloping paleolandsurfaces with gradients parallel to, but not as steep as, the overlying modern hillslopes. The Loveland Loess is much thinner on lower slopes than on broad ridgetops nearby, and the Sangamon Geosol takes

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up much of its thickness (Fig. 5). The Kennard Formation also occurs on at least a few low spur ridges and upper footslopes (Fig. 5), indicating development of at least some of the present valley network by some time during Kennard Formation accumulation. These observations, similar to those of Ruhe and Cady (1967) in southwestern Iowa, lead to three important interpretations. First, erosional truncation of the loess succession has been pervasive everywhere below wide, nearly level summits. Second, a subdued version of the modern topography may have developed by the time of Kennard Formation deposition (prior to OIS 6), and certainly was present during Sangamon Geosol formation (OIS 5 and 4). Finally, erosional truncation on hillslopes is not a recent phenomenon, but has been recurrent throughout the accumulation of the loess-paleosol sequence. 5.3. Regional thickness trends Assuming modest rates of erosion on upland summits, regional trends of loess unit thickness on those summits should approximately represent corresponding trends in loess accumulation rates, allowing reconstruction of loess transport direction and sources. We investigated regional thickness trends at two scales, following Mason (2001a), who noted that the Missouri valley loess source influenced Peoria Loess thickness only at sites o60 km west of the

Fig. 5. Contrast between loess stratigraphy in a core on a wide, gently sloping ridgetop (10-B-98) and on a lower slope nearby along edge of same ridge (2-B-99). Top of cores offset by difference in elevation. Top of 10-B-98 is natural ground surface, while Core 2-B-99 was sampled within a shallow roadcut. About 1.5–2 m of sediment originally covered Sangamon Geosol in 2-B-99, based on adjacent roadcut face. Sangamon Geosol in 2-B-99 is similar in thickness and degree of development to same soil in 10-B-98. Gray shades approximately represent light-dark color variation in cores.

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Missouri. Across the rest of Nebraska, Peoria Loess thins southeastward from nonglacial sources on the western Great Plains. A portion of the study area north of Omaha (Area A, Fig. 1b and c) was selected to capture local trends produced by transport from the Missouri valley source, while a larger region (Area B, Fig. 1b) was used to characterize trends related to western Great Plains sources. All thickness measurements used in this analysis are from upland summits. In Area A, we used continuous cores in which stratigraphic boundaries are clearly visible, as well as data from older mud-rotary testholes. Nearby cores and testholes were compared to develop criteria for consistently differentiating the units in the mud-rotary drilling logs, using texture, color, geophysical logs, and drilling speed changes. These criteria were then applied across Area B, where only mud-rotary testhole data were available. We could not consistently identify the Kennard Formation across Area B, and therefore only report data for that unit in Area A. Thickness trends of the Gilman Canyon Formation were not analyzed in either area, since the thickness of this unit at most sites (o1 m) is only slightly greater than the uncertainty in placing its boundaries. In Area A, Peoria Loess and Loveland Loess both systematically thin westward from the Missouri River valley, except for the Loveland Loess in core 3-B-98 (Fig. 6). In that core, the anomalously thin Loveland Loess contains coarse sand grains and a few pebbles, suggesting reworking by slope processes, and possible truncation by erosion. The trend of Loveland Loess thickness is steeper than that for Peoria Loess; near the Missouri, Loveland Loess is thicker, but beyond a distance of about 13 km, Peoria Loess is thicker (Fig. 6). Peoria Loess thickens slightly at the westernmost sites, just east of the Platte/ Elkhorn valley, but a similar trend reversal is not apparent for Loveland Loess. Kennard Formation thickness also tends to decrease slightly with distance from the Missouri,

Fig. 6. Thickness of individual loess units in relation to distance west from the western edge of the Missouri River valley.

but this trend is due almost entirely to the two westernmost measurements. Fig. 7a plots Peoria Loess thickness in Area B against distance downwind from the northwestern edge of thick Peoria Loess, which represents the minimum distance downwind from the regional nonglacial sources of Peoria Loess (Mason, 2001a). Peoria Loess thins gradually southeastward from those nonglacial sources, across most of Area B, corresponding to the distal part of the regional trend across Nebraska described by Mason (2001a). A cluster of points above that trend are all from sites within 10 km of the Missouri River Valley. Loveland Loess thickness across Area B follows a quite similar southeastward thinning trend (Fig. 7b), but offset to lower values and with more scatter. A cluster of points with thicker Loveland Loess lie within 25 km of the Missouri

Fig. 7. Thickness of Peoria Loess and Loveland Loess in relation to distance southeast from the northwestern edge of thick loess deposits in central Nebraska, as defined by Mason (2001a); this distance variable represents distance from nonglaciogenic loess source areas on Great Plains: (a) Peoria Loess, and (b) both Loveland and Peoria loess.

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River valley, and a few other high outliers occur farther west (Fig. 7b). 5.3.1. Interpretation Except near the Missouri River valley, thickness trends indicate that both Peoria and Loveland loesses were transported southeastward from similar northwestern source areas. Those sources were clearly much more productive during Peoria Loess deposition (OIS 2) than during Loveland Loess deposition (OIS 6), producing an offset in thickness trends. Higher trapping efficiency of OIS 2 vegetation versus OIS 6 vegetation could also account for some of this offset , but the importance of this effect cannot be evaluated given the lack of information on the region’s OIS 6 vegetation. Both loess units were also transported from the Missouri River Valley, a major local source. The influence of this Missouri Valley source extended farther west for the Loveland Loess, partly because it was superimposed on a smaller influx from northwestern sources. Even near the Missouri River, 435% of the total thickness of Peoria Loess is attributable to northwestern sources, based on extrapolation of the southeastward thinning trend (Fig. 7a). In contrast, almost all of the thick Loveland Loess near the Missouri Valley is probably derived from that valley (Fig. 7b). Data from Area A hint at a minor influx of Peoria Loess from the Platte/Elkhorn valley. A few thick outliers of Loveland Loess in Area B cannot be explained by proximity to the Missouri or other major rivers, and may represent other local sources, or simply errors in interpreting the drilling logs. The transport direction of the loess that makes up most of the Kennard Formation cannot be inferred with confidence from the thickness data. The slight apparent westward thinning may be an artifact of incomplete Kennard Formation preservation at the two westernmost sites (Fig. 6), since even very slow erosion on summits could have had a significant effect over the long time period represented by this unit.

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some cases a distinct fine silt ‘‘shoulder’’ on the primary peak (e.g. Loveland Loess at 1080 cm, Fig. 8b). Particle size profiles are strikingly similar, across all three cores (Fig. 9). Major shifts in the mode of the PSD, reflecting the position of the primary peak, occur at unit boundaries or at consistent stratigraphic positions within individual units (e.g. abrupt coarsening above the dark horizon marking the top of the lower Loveland Loess) (Fig. 9a). Particle size shifts at the same depths are also evident in plots of the clay (o2 mm) and fine silt (2–16 mm) (Fig. 9b). The Kennard Formation is consistently the finestgrained unit at a given site; clay contents measured by laser diffraction are 8–14% (Fig. 9b); equivalent to about 34–50% clay as measured by pipette (Fig. 2). The mode of Kennard Formation samples ranges from 10 to 20 mm.

5.4. Particle size analysis The three cores selected for particle size analysis are representative of loess stratigraphy on upland summits, and span a wide range of thickness for individual loess units. Stevens, 10-B-98, and 7-B-00, are from sites 6.3, 13.3, and 20.7 km west of the Missouri River valley, respectively (Fig. 1c). Core 7-B-00 is also immediately east of the Platte/ Elkhorn valley. Two cores contain the entire sequence from the contact between glacial till and the Kennard Formation up through Peoria Loess. One (Stevens) only penetrates the uppermost part of the Kennard Formation, where the depth limit of the coring equipment was reached. Particle size distributions (PSDs) of samples from all four loesses are similar to those reported for loess in other regions (Mason et al., 2003a; Sun, 2004), displaying a prominent primary peak between 10 and 40 mm, with a smaller secondary peak in the clay fraction (Fig. 8), and in

Fig. 8. Examples of particle size distributions determined by laser diffraction.

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Fig. 9. Vertical profiles of particle size parameters in three representative cores: (a) coarse silt content and mode and (b) clay and fine silt content (including both actual values obtained by laser diffraction, and estimated pipette clay content, from Omaha-area equation in Fig. 2). Gray bands identify the Sangamon Geosol and dark horizons in the lower Loveland Loess, to illustrate particle size variation associated with these zones (discussed in text).

Particle size profiles within the Kennard Formation are similar at sites 10-B-98 and 7-B-00, with peaks in the middle of the unit. The field-identified boundary between the Kennard Formation and the overlying Loveland Loess is marked by a distinct upward shift to a coarser mode and less fine silt (Fig. 9). Loveland Loess contains less coarse silt and more fine silt than the overlying Peoria Loess at the two western sites (10-B-98 and 7-B-00), but not at the eastern end of the transect (Stevens). The upper part of the Sangamon Geosol, with a peak clay content of 9–11% (laser diffraction) or about 38–40% (pipette), is finegrained compared to the rest of the upper Loveland Loess

(Figs. 8b and 9). The lower Loveland Loess consistently contains peaks of clay and fine silt content as high as those in the overlying Sangamon Geosol, however. These peaks are not associated with evidence of advanced pedogenesis, and their occurrence on high ridgetops appears to preclude a non-eolian origin. A distinct upward coarsening marks the transition from upper Sangamon Geosol to Gilman Canyon Formation at all three sites (Fig. 9). The Peoria Loess uniformly contains the least fine silt and clay of any of the units, and always has a relatively coarse mode. Particle-size trends along the three-core transect are summarized in Fig. 10. The Kennard Formation is omitted,

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Fig. 10. Variation of particle size mode with distance west from the Missouri River valley, using values averaged over portions of each stratigraphic unit. Upper and lower Loveland Loess distinguished as described in text, Peoria loess arbitrarily averaged over upper 4 m and lower 3 m, to capture any major changes in grain size over time.

since it was adequately sampled in only two of the cores, and Peoria and Loveland loess are each split into upper and lower parts. Overall, the loess units diverge with distance westward: upper and lower Loveland Loess are finer at the westernmost site (7-B-00) than at the eastern site (Stevens), while both upper and lower Peoria Loess and the Gilman Canyon Formation are coarser to the west. For all but the lower Loveland Loess, however, there is at least a slight grain size minimum at the middle site (10-B-98), with an increase westward to 7-B-00. 5.4.1. Interpretation Weathering and clay illuviation have certainly affected parts of the loess sequence, such as the Sangamon Geosol. Nonetheless, we believe that much of the observed particlesize variation is sedimentary in origin, because all of the up-section changes in particle size are marked by major shifts in the mode of the PSD (Fig. 9a). Studies of unweathered loess indicate that such large variation of the mode can clearly be produced by sorting during eolian transport, as a function of transport distance and/or wind speed (e.g. Mason et al., 2003a; Sun, 2004). It is much less plausible to attribute the up-section variation of the mode to weathering. For example, if the fine mode (o20 mm) of the Kennard Formation in Core 7-B-00 resulted from weathering of loess originally as coarse as the overlying Peoria Loess (mode 430 mm), a large percentage of the coarse and medium silt grains originally present must have been destroyed or greatly reduced in size by weathering. This would require intense or very long-term weathering, however, since it is evident from thin sections that quartz and feldspars dominate the coarse and medium silt fraction

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of all the loess units. In fact, much of the variation in fine silt and clay content tracks along with changes in the mode, suggesting the predominance of sedimentary control even for those size fractions. An alternative explanation for sedimentary control of PSD variation is that some units were transported in a more aggregated condition than others. For example, the abundant fine silt and clay in the Kennard Formation could have been deposited as coarse silt-size aggregates, with a much coarser effective mode than indicated by the disaggregated PSD. Assuming that the variation in the mode of the PSD, and perhaps other grain-size parameters, is largely sedimentary in origin, the particle size profiles can be used to test stratigraphic correlations between the three cores. The similarities between the three profiles, extending even to minor peaks and troughs within units, strengthens the argument that the same near-continuous sequence of loess units is preserved at all three sites. Furthermore, the especially fine mode of the Kennard Formation probably reflects transport over longer distances, by weaker winds, and/or in a more aggregated condition than the overlying units. The same may be true of fine-grained zones in lower Loveland Loess. Particle size trends are broadly consistent with thickness trends across the same area. Peoria Loess thickness reaches a minimum between 10 and 15 km west (Fig. 6) of the Missouri River, where the finest particle size was observed in the three-core transect (Fig. 10). The coarsening of Peoria Loess from 10-B-98 to 7-B-00 corresponds to an increase in thickness, indicating increasing predominance of western loess sources and/or input from a local Elkhorn/ Platte Valley source. Loveland Loess is much coarser at the site nearest the Missouri River (Stevens) than at the westernmost site (7-B-00), consistent with thickness data indicating predominance of the Missouri Valley source for that unit in the study area. 5.5. Micromorphology Thin-section analysis provides information on the pedogenic processes that were active during or after loess deposition, and provides additional paleoenvironmental insights. We systematically examined micromorphology in three cores (Fig. 11), similar to the approach used by Kemp et al. (1996, 1997). The Peoria Loess was thoroughly sampled only in Core 3-B-99 (Fig. 11). In that core, the upper Peoria Loess below the surface soil has weak blocky, granular, or channel structure. Granular or spongy structure is produced mainly near the ground surface by soil faunal activity (FitzPatrick, 1993, p. 124; Bullock et al., 1985, p. 46), and where it occurs well below the modern surface it is likely to have developed during upper Peoria Loess accumulation (i.e. the structure formed in recently deposited near-surface loess before it was deeply buried by ongoing accumulation). Blocky structure is commonly observed in subsoil horizons and is produced mainly by wetting and drying (FitzPatrick,

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Fig. 11. Micromorphological observations on thin sections from three representative cores.

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1993, pp. 118–131). Channel structure is similar to a massive condition except for the presence of linear biogenic pores such as root channels (Bullock et al., 1985, p. 46). Blocky and channel structure could have developed both during and after loess accumulation. The lower several meters of Peoria Loess typically has a variety of lenticular or ‘‘banded’’ microstructures (Figs. 11 and 12a, b), in 3-B-99 and the more limited intervals sampled in the other cores. In some cases these microstructures are visible in hand specimens, and occur at roughly the same depths as macroscopic banding observed in the field. They are not artifacts of coring, because hydrous iron oxides coat some of the planar voids defining the structure. In some cases (Fig. 12a), the lenticular or banded structure resembles that produced by freeze–thaw cycles in experiments by Coutard and Mucher (1985), and observed in full-glacial loess-derived paleosols in the Yukon by Sanborn et al. (2006). In other samples (Fig. 12b), we also observed thin (o0.1 mm) bands of laminated material, resembling surface seals or crusts formed by rainsplash or local slopewash (Coutard and Mucher, 1985). Similar lenticular or banded structure was

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not observed below the Peoria Loess (Fig. 11). Where it does not have this structure, lower Peoria Loess is usually massive. Granular structure is strongly expressed in the lower Gilman Canyon Formation and uppermost Sangamon Geosol (Figs. 11 and 12d). This interval is also the only part of the loess sequence in which we observed clearly recognizable excrements of soil fauna (Bullock et al., 1985, pp. 133–137) and vermiforms (FitzPatrick, 1993, pp. 131–133). The Sangamon Geosol has strong blocky structure (Figs. 11 and 12g, h). Loveland Loess below the Sangamon Geosol typically has a mixture of weak blocky, channel, or granular structure, similar to that observed in upper Peoria Loess (Fig. 11). Two of the dark colored bands in the Loveland Loess that resemble incipient soils in core 7-B-00 have well-preserved granular or spongy structure (Fig. 12e), while similar bands in the other cores are massive or have channel structure. Most thin sections from the Kennard Formation display weak to strong blocky and granular structure (Figs. 11 and 12f). Illuvial clay coatings in pores are absent in thin sections of Peoria Loess except within B horizons of the surface soil

Fig. 12. Photomicrographs. Vertical sections, top of frame toward groundsurface unless noted; PPL ¼ plane polarized light, XPL ¼ cross-polarized light: (a) Lenticular structure, lower Peoria Loess (3-B-99, 717 cm, PPL); (b) feature interpreted as soil crust, lower Peoria Loess (3-B-99, 673 cm, PPL); (c) massive lower Peoria Loess (3-B-99, 717 cm, PPL); (d) strong granular structure with dark iron oxide coatings (arrows) on some aggregates, Gilman Canyon Formation (10-B-98, 1027 cm, PPL); (e) spongy (compressed granular) structure in a dark band (incipient soil?) of the lower Loveland Loess (7-B00, 1373 cm, PPL); (f) granular to spongy structure in the Kennard Formation (7-B-00, 1857 cm, PPL); (g) blocky structure with illuvial clay coatings in pores (examples indicated by arrows), lower B horizon of Sangamon Geosol; also note rounded aggregate (ag) embedded in a matrix of contrasting color, suggesting physical mixing at some time in this horizon’s development (3-B-99, 1312 cm, PPL); (h) same as (g), but XPL; (i) illuvial clay coatings (arrows) lining pores, Sangamon Geosol B horizon (7-B-00, 1206 cm, XPL, up to left); (j) illuvial clay filling a large pore in the Kennard Formation (7-B-00, 1807 cm, XPL); (k) infilling of clean silt in a pore (si), transition from Gilman Canyon Formation to Sangamon Geosol (7-B-00, 1120 cm, PPL); (l) illuvial clay coatings (cc) and striated b-fabric (str) in lower Kennard Formation (7-B-00, 2192 cm, XPL).

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profile, and are rare in the Gilman Canyon Formation , but are nearly ubiquitous in thin sections from the upper Sangamon Geosol downward through Loveland Loess and the Kennard Formation (Fig. 11). Coatings are thickest and most abundant in Bt horizons of the Sangamon Geosol and especially in some of the welded Bt horizons of the Kennard Formation (Figs. 11 and 12g, h, but also occur in BC or C horizons in Loveland Loess. Features indicating disruption and soil mixing appear near the Gilman Canyon Formation to Sangamon Geosol transition in all three cores, and near the contact of the Kennard Formation with the underlying till in core 3-B-99. These include infillings of relatively clean silt (Fig. 12k) and rounded fragments of soil material (‘‘pedorelicts’’) embedded in material with distinctive color and texture (Fig. 12g and h). Disruption and silt translocation may result from frequent, deep freeze–thaw cycles (FitzPatrick, 1993, p. 185), intensified tree throw or burrowing by soil fauna. Clay illuviation clearly post-dates disruption in some cases (e.g. Fig. 12g and h), while in other cases, clay coatings embedded within aggregates suggest disruption after illuviation. Redoximorphic features, mainly pedofeatures produced by local iron (hydr)oxide accumulation, are common throughout all three cores (Fig. 11). Given the partial saturation of the loess sequence observed today, these features could have formed recently and do not necessarily record conditions during loess accumulation. Striated b-fabric in parts of the Kennard Formation records stresses generated during shrink–swell (wetting–drying) cycles in this clay-rich material, probably when it was much closer to the ground surface (Figs. 11 and 12l). 5.5.1. Interpretation Weak granular structure in the upper Peoria Loess records moderate soil faunal activity during loess accumulation. The preservation of surface seals and lenticular structure indicate that the lower Peoria Loess has been minimally affected by bioturbation or other pedogenic processes. Given the low clay content of the lower Peoria Loess and its frequent near-saturation, however, it is possible that hydroconsolidation (Rogers et al., 1994) has compressed some original granular structure into a massive condition. In contrast, the strong granular structure in the Gilman Canyon Formation almost certainly reflects longer residence time within the zone of pedogenesis, consistent with numerical ages indicating slow accumulation. The abundant illuvial clay and strong blocky structure of the Sangamon Geosol clearly indicate long-term exposure to pedogenesis. Clay translocation was probably minimal during Peoria Loess and Gilman Canyon Formation deposition because loess accumulation was more rapid than depletion of exchangeable or readily soluble Ca2+ and Mg2+, which inhibit clay mobility (Schaetzl and Anderson, 2005, pp. 363–366). Extensive clay illuviation in the Sangamon Geosol probably occurred during OIS 5 and

OIS 4, after Loveland Loess deposition had ceased, and deep leaching of base cations occurred. The sparse illuvial clay at depths up to 7 m below the top of Loveland Loess may reflect either some base cation leaching and clay translocation during loess accumulation, or unusually deep illuviation during Sangamon Geosol formation. If any part of the Loveland Loess originally had lenticular structure, relict surface seals, or similar features, they have been destroyed by bioturbation. We interpret this observation as indicating that Loveland Loess accumulation was on average slower than Peoria Loess accumulation, and/or soil fauna were more active during Loveland accumulation. Some post-burial loss of granular structure through hydroconsolidation is even more likely in lower Loveland Loess than in Peoria Loess, however, because of greater burial depth and overburden pressure. Where granular structure is preserved in dark bands of the lower Loveland Loess (Core 7-B-00), it is finer-grained, more cohesive, and probably more resistant to compaction. The extensive illuvial clay in the Kennard Formation is consistent with its macroscopic resemblance to an overthickened B horizon. The variable microstructure of this unit is consistent with multiple minor pulses of loess accumulation, with intervening pedogenesis and clay translocation. 6. Discussion Our results highlight both the pervasive erosion of loess units across much of the study area landscape, and the critical importance of the more complete stratigraphic record preserved beneath upland summits. Despite the modest relief of the study area, and its distance from late Pleistocene ice margins (Fig. 1a) and associated periglacial conditions, all loess units have clearly been truncated on hillslopes and narrower ridgetops, where most loess outcrops occur. Cores and other drill holes on upland summits reveal much thicker Loveland Loess than is exposed in any known outcrop, moreover, we find that pre-Loveland loess is much more extensive than previously reported. We suggest that much outcrop-based work on loess stratigraphy elsewhere in the mid-continent, where such highdensity (and high-quality) subsurface data are rarely available, may be biased by the effects of erosion. Despite the strong evidence for post-depositional erosion, our results also provide compelling new evidence that a relatively complete sequence of Middle to Late Pleistocene loess is preserved on wide, gently sloping summits in parts of the central US. The same sequence of loess units appears under almost all broad, gently sloping summits across the study area, and particle size profiles are similar at localities several kilometers apart. The regional thickness trends followed by individual loess units, particularly within Area A, are most readily explained as products of decreasing deposition rate with distance from loess sources. It is unlikely that depositional trends would be preserved after significant erosion on the ridgetops, and even more

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unlikely that the thickness trends could be produced by erosion. In fact, the thickest Peoria and Loveland Loess is in the much more dissected areas with greater relief near the Missouri Valley. Finally, loess stratigraphy on hillslopes suggests the present valley network was partially developed when the oldest loess unit (the Kennard Formation) was accumulating, and was clearly present by OIS 6. This observation implies long-term stability of major interfluves, favoring long-term loess preservation on wide and level ridgetops. The relatively complete stratigraphic record of loess under broad summits prompts reconsideration of the concept of a Middle to Late Pleistocene acceleration in long-term loess accumulation rates. Even considering the large uncertainty in the estimated age of the Kennard Formation, its 6–11 m thickness represents all loess accumulated over several hundred thousand years, while the much thicker overlying loess probably accumulated since the beginning of OIS 6 (o200 ka). Thus, the loess accumulation rate, when averaged over such long time intervals, was greater from OIS 6 onward. The long-term rate of Kennard Formation accumulation may have been limited because glaciation of the Missouri Basin above the study area was infrequent and/or limited in extent during the part of the Middle Pleistocene represented by this unit. In fact, the fine-grained nature of the Kennard Formation suggests that it could be dominated by distal loess from western Great Plains sources. The slight westward thinning of the Kennard Formation conflicts with this interpretation, but could be an artifact of incomplete preservation. Based on thickness trends in Fig. 6, if the Loveland Loess was derived entirely from western Great Plains sources, it would be only 1–3 m thick in easternmost Nebraska. The Kennard Formation as a whole could contain several distal loess increments of this magnitude. The shift from slow Kennard Formation accumulation to rapid deposition of the exceptionally thick Loveland Loess probably records the initiation of the Missouri River valley as a major drainage way for glacial melt water, and as a major loess source. Our results, combined with existing evidence that Loveland Loess thins and fines eastward from the Missouri River in Iowa (Ruhe and Cady, 1967), confirm the glaciogenic origin of thick Loveland Loess in eastern Nebraska and southwestern Iowa. This probably indicates extensive glaciation of the Missouri Basin during OIS 6. This study does not provide clear evidence of a major increase in loess accumulation rate from the Loveland Loess to Peoria Loess near the Missouri River. In fact, the Loveland Loess is actually thicker than Peoria Loess and the Gilman Canyon Formation combined at sites near the Missouri River, suggesting that the Missouri Valley glaciogenic loess source was less productive in OIS 3 and 2 than in OIS 6. Our new subsurface data reveal Loveland Loess and Peoria Loess to be broadly similar units near the Missouri: both are thick loesses with only modest

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pedogenic overprinting, and much of the pedogenic alteration is actually ‘‘top-down’’ soil development after loess accumulation had largely ceased. Thus, given another 60,000–70,000 years of surface exposure, Peoria Loess might look much like the thick Loveland Loess near the Missouri, having a very well-developed surface soil with deep illuviation of clay. Paleopedologic evidence does suggest the possibility of a modest difference in accumulation rate between Loveland and lower Peoria loess. Loveland Loess may have accumulated at a slower rate because it lacks the early post-depositional features common in the lower Peoria, such as lenticular structure and the putative remnants of surface seals. Numerous dark bands in the Loveland Loess, most likely incipient soils, suggest multiple episodes of slower accumulation and/or more effective pedogenesis. These contrasts might also reflect differing climatic conditions during loess accumulation, however. Despite the similarities of Peoria and Loveland loesses near the Missouri Valley, a broader regional perspective demonstrates a key contrast, which is the most important result of this study. Thickness trends reported here (Figs. 6 and 7) indicate that nonglaciogenic loess sources became much more important during OIS 2 than during OIS 6, and probably at any time since the onset of Kennard Formation deposition in the Middle Pleistocene. Most of the Kennard Formation and a small fraction of the Loveland Loess in the study area may be derived from nonglaciogenic Great Plains sources, but both represent low rates of dust influx to the study area. The Missouri Valley glacial loess source was apparently less productive during OIS 2 than during OIS 6, but at the same time there was a major increase in the influx of dust from nonglacial sources on the Great Plains. This interpretation is consistent with evidence that Peoria Loess west of the study area in central Nebraska, closer to those nonglacial sources, experienced some of the highest mass accumulation rates known worldwide (Roberts et al., 2003). Furthermore, extensive nonglacial dust production occurred during OIS 2 on the Iowan Erosion Surface of northeastern Iowa and southeastern Minnesota, producing some of the thickest Peoria Loess in the upper Midwest (Mason et al., 1994, 1999; Bettis et al., 2003). The high rate of dust production from nonglacial sources in both the Great Plains and the Upper Midwest during OIS 2, relative to OIS 6 and earlier glaciations, is difficult to explain through local, non-climatic mechanisms. Instead, this regional pattern strongly suggests a unique set of climatic conditions during OIS 2, most likely some combination of lower temperatures and lower effective moisture south of the ice margin. Possible mechanisms of enhanced dust production under these conditions include sparse vegetation cover on winderodible sediments, frequent exposure of bare surface soil through frost action, and accelerated hillslope erosion and sediment delivery to sparsely vegetated floodplains that were dust sources.

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From a broader perspective, especially cold and/or dry conditions during OIS 2 could help explain the minimal degree of pedogenesis or even bioturbation evident in lower Peoria Loess in the study area and other parts of the midcontinent. To the extent that the unusual thickness of glaciogenic Peoria Loess along the lower Mississippi or Ohio rivers, relative to older loesses, is not an artifact of post-depositional erosion, unique climatic conditions during OIS 2 may help explain that contrast as well. Low temperature or effective moisture would limit the colonization of floodplain surfaces by vegetation, possibly increasing dust production from valley sources. These same climatic conditions may also have fostered instability of unglaciated landscapes southwest of the ice sheet, activating extensive non-glacial dust sources. Simulations of the last glacial maximum climate with general circulation models of widely varying resolution and complexity have consistently indicated much lower temperatures and precipitation than at present in areas where abundant nonglaciogenic loess production occurred in OIS 2 (e.g. Manabe and Broccoli, 1985; Bartlein et al., 1998; Shin et al., 2003). These climatic conditions are in part a direct effect of the ice sheet itself, and the height of the ice sheet is a key factor in their occurrence (Felzer et al., 1996). Consistent with these results, Williams (2002) recently inferred low tree-cover density across much of North America during OIS 2 from a large pollen dataset. A recent modeling study suggested occasional intense rainstorms along the ice margin, superimposed on generally dry conditions (Bromwich et al., 2005), which is consistent with some of the mechanisms for nonglaciogenic dust production described above (e.g. accelerated slope erosion). It is not clear, however, why these ice sheet effects would be greater in OIS 2 than in OIS 6. Illinoian (OIS 6) glaciation was more extensive in the Midwest than the Wisconsin (OIS 2) glaciation, and marine records suggest similar global ice volume in both stages. Clark and Pollard (1998) proposed that the Laurentide Ice Sheet was thicker but less extensive at glacial maxima after about 1 Ma than earlier in the Pleistocene or Pliocene, a transition attributed to a decreasing area covered by soft regolith beneath the ice sheet (Clark and Pollard, 1998; Roy et al., 2004b). Perhaps the OIS 2 ice sheet represented an extreme in ice sheet thickness, covering less area but standing higher than even the OIS 6 ice sheet, and generating conditions unusually favorable for dust production in adjacent landscapes.

7. Conclusions In eastern Nebraska, broad upland summits are underlain by a relatively complete Middle to Late Pleistocene loess-paleosol sequence. Post-depositional erosion has nonetheless been a strong influence on the preservation of the loess succession across the landscape as a whole, particularly on steeper slopes.

The fine-grained Middle Pleistocene Kennard Formation, newly recognized in this paper, may represent multiple thin increments of distal loess from northwestern sources on the Great Plains. The thick Loveland Loess in the study area probably marks the first time the middle Missouri River valley functioned as a major loess source. Peoria Loess (deposited in OIS 2) and Loveland Loess (OIS 6) share many characteristics, but Peoria Loess incorporated much more nonglaciogenic dust from Great Plains sources. This contrast may reflect colder and/or drier climatic conditions uniquely conducive to nonglaciogenic dust production during OIS 2. Acknowledgements This research was supported by the US Geological Survey (STATEMAP Program), the Conservation and Survey Division, School of Natural Resources, U NL-Lincoln, and the National Science Foundation (EAR-0087916 and EAR-0087572). The CSD testhole network is the legacy of the late Vince Dreeszen, who also encouraged this study. Greg Grosch and Jim Roberts carried out the extensive drilling required, with help from Bryan Penas, Mike Ponte, Mike Beshore, Joni Koskovic, Cynda Timperley, and many others. We thank Rob Kemp and Dave Grimley for helpful reviews and Paul Hanson for core descriptions. Appendix A. Electronic Supplementary Material The online version of this article contains additional supplementary data. Please search for doi:10.1016/j.quascirev. 2006.10.007.

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