CHAPTER SOILS AND PALAEOSOLS IN GLACIAL ENVIRONMENTS 17 J.A. Mason1 and P.M. Jacobs2 1University of Wisconsin-Madison, Madison, WI, United States, 2University of Wisconsin-Whitewater, Whitewater, WI, United States
17.1 INTRODUCTION In 1873, the geologist Amos Henry Worthen recognized two buried soils within the Quaternary stratigraphy of Sangamon County, Illinois, USA (Worthen, 1873; Follmer, 1978). One of these sep- arated the surficial loess mantle from the underlying ‘boulder clay’ or till and the other occurred below the same till. The upper soil was later named the Sangamon Soil by Frank Leverett (1898), separating sediments of the last glacial cycle from those of the penultimate glaciation, Marine Isotope Stage (MIS) 6 (Fig. 17.1), and still recognized as one of the most important Quaternary stratigraphic markers of the central United States (Follmer, 1983). At roughly the same time in the late 19th century, Albrecht Penck and Eduard Bru¨ckner noted the stratigraphic and geomorphic evi- dence for several glaciations in the foreland of the Alps, including—in modern terms—the contrast- ing degree of soil development on the surface of terraces correlated with each glaciation. In particular, the reddish clay-rich weathering profiles on the terraces attributed to aggradation during the Mindel glaciation were seen as evidence of a ‘great interglacial’ before the subsequent Riss gla- ciation (Penck and Bru¨ckner, 1909; Kukla, 2005). Thus, by the late 19th century, two of the most important ways in which soils can help interpret gla- cial stratigraphy were widely recognized. Buried palaeosols form stratigraphic markers representing a significant interval of nondeposition and surface exposure of the sediments they are formed in. Surface soils, through their degree of morphological development and weathering, can be used as evidence of the relative age of glacial landforms. Two other types of evidence available from palaeosols are now also well-established. First, the preservation of a palaeosol at a particular stratigraphic boundary demon- strates the lack of substantial erosion when the overlying sediment was deposited. If the overlying sedi- ment is till, that observation may be relevant to interpretations of glacial processes. Conversely, local or widespread truncation of palaeosols may be important evidence in interpreting postglacial landscape evolution. Second, palaeosols in glacial deposits can also potentially be sources of evidence on past cli- mate and vegetation. All of these interpretations of palaeosols depend heavily on a large body of research on surface soils and practical soil surveys, which provide essential information on how soils vary with environmental factors at the global and local scale, and how they develop over time. This chapter first reviews the chronosequence approach that has often been applied in studies of sur- face soils in glacial sediments, and the many insights it has provided on processes and rates of soil
Past Glacial Environments. DOI: http://dx.doi.org/10.1016/B978-0-08-100524-8.00018-X © 2018 Elsevier Ltd. All rights reserved. 587 588 CHAPTER 17 SOILS AND PALAEOSOLS IN GLACIAL ENVIRONMENTS
FIGURE 17.1 Sangamon soil in Illinois. (A) Stratigraphic section from Thomas Quarry, Illinois (39 310 N, 90 400 W), illustrating Sangamon soil as key stratigraphic marker between sediments of last glaciation and penultimate glaciation. Section also includes an older palaeosol (Yarmouth soil). (B) Sangamon soil profile at Thomas Quarry. (C) Pedocomplex consisting of Sangamon soil and overlying Farmdale soil at site in northern Illinois (41 510 N, 88 280 W) described by Jacobs et al. (2009). Sangamon soil formed in Illinoian (MIS 6) glacial diamicton and was buried by aeolian and colluvial Robein silt during MIS 3; Farmdale soil developed in Robein silt. See Fig. 17.4 for site stratigraphy. formation over time. We then discuss the continued value of this information in glacial geomorphic map- ping and stratigraphy, and the importance of recognizing how soils vary across glaciated landscapes and how those patterns change with ongoing postglacial landscape evolution. Finally, we turn to key issues that have arisen in the recognition and interpretation of palaeosols in glacial settings, as stratigraphic markers and as indicators of past environments (and in some cases, past geomorphic processes).
17.2 SURFACE SOILS IN GLACIATED LANDSCAPES 17.2.1 CHRONOSEQUENCE STUDIES OF SOIL DEVELOPMENT IN GLACIAL SETTINGS In Hans Jenny’s well-known model of pedogenesis, the present state of a soil is a function of fac- tors including climate, organisms, relief (more accurately described as geomorphic setting), parent material, and time since the start of pedogenesis (Jenny, 1941). While alternatives to Jenny’s model 17.2 SURFACE SOILS IN GLACIATED LANDSCAPES 589
have been proposed (e.g., Simonson, 1959; Runge, 1973; Johnson et al., 1990), a large proportion of all subsequent pedological research has followed research designs compatible with Jenny’s model. Essentially, a sequence of soils in which one factor varies much more substantially than the others is used to assess the effects of that factor. Chronosequences of soils formed in glacigenic sediments and loess of similar lithology but vary- ing age are especially common and relevant to this chapter (Fig. 17.2). In most cases, a quantitative model, or chronofunction, is fit to the observed data from a chronosequence study. While chronose- quences clearly provide insight on rates and processes of soil formation, chronosequence research has also been motivated by the goal of inferring the age of sediments and landforms from the soils formed on them. Many early, groundbreaking chronosequence studies emphasized relative ages or broadly estimated numerical ages of the soils that were compared (e.g., Birkeland, 1964). More recent chronosequence studies have taken advantage of major advances in glacial chronology, and are based on numerical ages from cosmogenic radionuclide (CRN) analysis or other methods (e.g., Douglass and Bockheim, 2006; Dahms et al., 2012). Few if any chronosequences do not involve at least minor violations of the assumption that only one factor—soil age—varies across the soils studied. Minor variations in parent material or vegetation history are difficult to avoid; more importantly, older soils
FIGURE 17.2 Soils in glacial sediment of three ages in the Wind River Mountains, Wyoming. (A) Soil formed in Alice Lake till (middle to early Holocene) above modern treeline at B3380 m elevation (soil BFL-7 in Dahms, 2002). (B) Soil formed in Late Pinedale till (15 16 ka) at modern treeline, B3260 m (soil WL-1 in Dahms, 2002). (C) Soil formed in Middle Pleistocene till at B2030 m (HR-1 in Dahms, 2004). Soil age is the primary factor in the increasing soil depth and reddish-brown B horizon development from (A) (C), though the soils also differ in climate and vegetation history because of elevation. Scales in centimeters. Photos by Dennis Dahms, used with permission. 590 CHAPTER 17 SOILS AND PALAEOSOLS IN GLACIAL ENVIRONMENTS
developed over one or more full glacial cycles will have experienced a different range of climatic conditions than those formed since the last glaciation. The degree to which the assumptions of a chronosequence can be relaxed is essentially a philosophical issue, and it is clear that chronose- quences provide important information on the broad patterns of soil development in glacial settings, in spite of obvious variation in other factors. Early soil development in glacial sediments is especially well understood from chronosequences of soils developed on surfaces deglaciated within the past few centuries, often with excellent age control from historical observations. Soil pH declines quickly in some cases, and organic matter (and organic nitrogen) accumulation is initially rapid but quickly decreases as an approximately steady state is approached (Crocker and Major, 1955; Jacobson and Birks, 1980; James, 1988; Burt and Alexander, 1996; Egli et al., 2006b). Silicate mineral weathering is detectable in the first 150 years of pedogenesis in granitic tills in the Alps, consistent with high initial rates of that process as well (Mavris et al., 2010). Studies of older soils on alpine glacial moraines and glaciofluvial terraces have revealed trends in soil morphology that are common to long-term chronosequences in a wide range of glacial and nonglacial settings (Bockheim, 1980; Harden and Taylor, 1983; Birkeland, 1990; Vidic, 1998; Sauer et al., 2015). In subhumid to semiarid climates (e.g., those characterizing glaciated parts of many mountain ranges of the western United States), the most important field-observed morpholog- ical changes over 10,000 years or longer timescales involve B (subsoil) horizons that thicken, develop stronger structure, become redder, and accumulate clay and sometimes silt (e.g., Birkeland, 1964; Mahaney, 1978; Burke and Birkeland, 1979; Hall and Shroba, 1993). Pedogenic carbonate accumulation, a process well-known from chronosequences in non-glacial settings (Machette, 1985), becomes progressively more important in drier climates (e.g., Hall and Shroba, 1993; Douglass and Bockheim, 2006). Podzolization (translocation of iron, aluminium, and organic matter to the B horizon) predominates over clay translocation in some humid settings (Stevens and Walker, 1970; Birkeland, 1984). In the cold climate of Yukon, Canada, soils in a long-term chrono- sequence in glacial deposits do not display reddening or evidence of clay translocation, but depth of cryoturbation is greater in older soils (Dampier et al., 2009). In cold, hyperarid Antarctica, long- term salt accumulation occurs in Pliocene to Pleistocene soils (Bockheim, 2013). Laboratory analyses confirm long-term accumulation of pedogenic iron oxides and hydroxides, consistent with B horizon reddening that is evident in the field (Birkeland et al., 1989; Mahaney et al., 2009; Dahms et al., 2012). Clay and silt that accumulate over time may originate through weathering of the original parent material, but clay, silt, and carbonate can also be added to the soil surface over time as dust (Mokma et al., 1972; Dahms, 1993; Applegarth and Dahms, 2004; Bockheim and Douglass, 2006), and clay can also be produced by rapid weathering of volcanic ash added to the soil in regions near volcanic centres (Burkins et al., 1999). Weathering can be assessed directly through mass balance analysis of major element losses, based on an index element assumed to be relatively stable during long-term pedogenesis, usually Ti or Zr (Brimhall and Dietrich, 1987; Chadwick et al., 1990). Mass balance analysis of soils on cosmogenic nuclide-dated moraines up to 1 million years old demonstrates progressive loss of base cations through silicate mineral weather- ing, and in some cases significant loss of Si and Al, with clear evidence for a decreasing rate of loss over time (Taylor and Blum, 1995; Dahms et al., 2012). A decline in weathering rates over time is expected on theoretical grounds, reflecting depletion of the more weatherable minerals (Colman, 1981). Similarly, net gain of organic matter should 17.2 SURFACE SOILS IN GLACIATED LANDSCAPES 591
decline to near zero as inputs are balanced by decomposition, an expectation supported by chrono- sequences of young soils on newly deglaciated terrain. Jenny (1941) suggests that rates of pedo- genic processes in general should decline as the soil approaches equilibrium with factors other than time. If so, then there is an upper age limit beyond which a soil property cannot be used to estimate the age of the landform or the time represented by a palaeosol within a stratigraphic succession. It has been questioned whether long-term chronosequences support the idea that many soil morpho- logical properties reach a steady state (Bockheim, 1980; Birkeland, 1990), though B horizon clay content has been noted to approach near steady-state values after 100 200 kilo years of soil forma- tion in Pleistocene interglacial climates of the Midwestern United States (Grimley et al., 2003). Some functions often fit to chronosequence data have this assumption of declining rate of change built-in, emphasizing the importance of chronofunction model selection (Schaetzl et al., 1994). The influence of climate on rates of pedogenic processes should, in theory, be evident when chronosequences from different climatic regimes are compared. In a comparison of chronose- quences from the European Alps and the Wind River Range of Wyoming, Dahms et al. (2012) found evidence for more rapid weathering in A horizons in the wetter Alps, and also inferred a cli- matic effect on accumulation of poorly crystalline forms of Al, Si, Fe, and Mn. Other studies have found contrasts in pedogenic process rates related to climatic variation within the same region, including microclimatic effects of slope aspect (Birkeland, 1994; Egli et al., 2006a, 2008). Bockheim’s (1980) analysis of a large number of chronosequences identified climatic effects on some but not all soil properties considered. Many chronosequence studies involve small numbers of soils, but where larger numbers of soils have been sampled, it is clear that variability among soils of the same age can be quite high in gla- ciated landscapes (e.g., Schaller et al., 2010; Dahms et al., 2012). This kind of scatter does not negate the value of the information on soil development trends available from many chronose- quences, but it does imply large uncertainties in inverting a chronofunction to estimate the numeri- cal age of a soil. In fact, it appears that chronosequence data have most often been used to inform more qualitative distinctions between landforms of substantially different ages, such as moraines formed in Holocene, Late Pleistocene, or earlier glaciations. In the United States, the Sierra Nevada of California and Wyoming’s Wind River Range provide especially good examples of the value of soils (e.g., Fig. 17.2)—together with other evidence such as moraine morphology—for regional cor- relation and mapping of sediments and landforms representing major glacial advances (e.g., Birkeland, 1964; Richmond, 1965; Burke and Birkeland, 1979; Berry, 1994; Dahms, 2002). In those areas, early work on CRN dating was largely focused on testing and refining key aspects of the resulting age framework, such as the distinction between Pinedale (last glaciation) and older Bull Lake and pre-Bull Lake moraines (Phillips et al., 1990, 1997; Gosse et al., 1995).
17.2.2 POSTGLACIAL SOIL LANDSCAPE EVOLUTION Glaciated landscapes undergo substantial postglacial modification over timescales ranging from 102 to 106 years. Moraines and till plains dating to MIS 2 display subtle to very pronounced hummocky topography with many closed depressions. Extension and filling-in of the drainage network results in more profound topographic modification (Ruhe, 1952) and eventual disappearance of most gla- cial landforms (Fig. 17.3). In alpine settings, moraines become lower and smoother over time (Meierding, 1984; Colman and Pierce, 1986). The erosion and deposition involved in this landscape 592 CHAPTER 17 SOILS AND PALAEOSOLS IN GLACIAL ENVIRONMENTS
FIGURE 17.3 Example of long-term postglacial landscape evolution in Iowa, USA, associated with major changes in soil-landscape patterns. (A) Topographic transition from area glaciated in MIS 2 by Des Moines Lobe of Laurentide Ice Sheet, to landscape last glaciated in the Middle Pleistocene and reshaped since then by hillslope and fluvial erosion. (B) Detailed view of Des Moines Lobe landscape, dominated by minor moraines and other glacial landforms, with many closed depressions and minimal drainage network development. (C) Detail of older landscape with fully integrated drainage; broad ridgetops are capped by loess over glacial sediment. Hillshades produced from LiDAR data by Iowa DNR. evolution inherently alter the landscape-scale pattern of soils, affecting their morphology, mineral- ogy, and other properties. Pedologists have extensively investigated soil patterns on glaciated land- scapes, often in the context of the catena, a predictable downslope sequence of soils, usually linked by transfers of water and sediment (Milne, 1936; Sommer and Schlichting, 1997). Substituting space for time, catenas in glaciated landscapes of varying age indicate how postgla- cial landscape evolution alters the soil landscape. Chronosequences of catenas on alpine moraines 17.2 SURFACE SOILS IN GLACIATED LANDSCAPES 593
indicate that B horizon thickness, clay accumulation, and mineral weathering are greater in lower slope positions, and this contrast is more pronounced on older moraines (Swanson, 1985; Berry, 1987). While these topographic effects may in part be related to greater water flux through the soil on lower slopes, they also reflect modification of the initially rugged topography by slope pro- cesses. Schaller et al. (2009) found that failure to account for erosion of moraine crests led to sub- stantial underestimation of weathering rates in that landscape position using mass balance analysis. In relatively young glacial landscapes, sediment eroded from slopes and ridge crests accumulates in the many closed depressions (Walker and Ruhe, 1968; Burras and Scholtes, 1987; Dahms, 1994), and soils may build upward as they develop (Hall and Anderson, 2000). Depressions are also often locations of at least seasonal upward transport of solutes into the soil profile. As a consequence, soils in the depressions are usually poorly drained and in subhumid to semiarid climates can accu- mulate pedogenic carbonates at a shallow depth or accumulate exchangeable sodium, which enhances formation of clay-rich B horizons (Seelig and Richardson, 1994). Where more time has passed since glaciation and stream dissection has made more progress, proximity to valley margins and the loess and till stratigraphy on broad drainage divides strongly influences soil drainage (Coleman and Fenton, 1982). A final important set of processes involves recurrent aeolian deposition and/or erosion on older glaciated landscapes. Where thin, episodically deposited loess is incorporated into an upward-building soil, but where it is thicker it buries exist- ing soils and resets pedogenesis (McDonald and Busacca, 1990; Woida and Thompson, 1993; Almond and Tonkin, 1999). Hall (1999) describes a chronosequence of soils in glacial sediments in a semiarid environment with strong winds today, near the foot of the Wind River Range in Wyoming, USA, in which morphological characteristics such as B horizon thickness show no long- term trends, although the oldest soil has accumulated more pedogenic carbonate than the others. This pattern was attributed to recurrent wind erosion followed by deposition of loess, which effec- tively resets pedogenesis in the upper part of these soils. Palaeosols should display similar variation across landscapes that were buried at various stages of postglacial evolution. The limited existing data on that variation suggest it may be essential information for accurate reconstruction of palaeoenvironments from palaeosols in glacial settings. In east-central Illinois, Follmer (1982) found substantial morphological variation of the Sangamon soil across a landscape underlain by Illinoian (MIS 6) till and now buried by loess, mainly related to palaeodrainage conditions and redistribution of soil into depressions. Jacobs et al. (2009) described fairly uniform texture-contrast morphology along a catena of Sangamon soils formed in highly dolomitic glacial sediments in northern Illinois (Fig. 17.4), suggesting long-term stability of the interglacial landscape at that study site, with little downslope change in soil morphology related to erosion or deposition. Jacobs (1998b) found that soil morphological trends across a palaeolands- cape marked by the Sangamon soil in Indiana mimicked those on the modern land surface, but clay mineral weathering patterns differed between the modern and buried landscape, apparently reflect- ing greater seasonal extremes in hydrology in the modern environment. In both of the latter two studies, clay mineral weathering indicated a greater degree of mineral alteration in the palaeosols compared to the modern landscape. The influence of glaciogenic sediment particle size on soil mor- phology and weathering characteristics was still evident and important after approximately 100 kilo years of soil formation, however (Jacobs, 1998a). Finally, research since the 1950s in southern Iowa (Fig. 17.3) has identified a characteristic spatial pattern of palaeosols buried by loess in a landscape last glaciated in the Early Middle Pleistocene, and now dissected by a well-developed 594 CHAPTER 17 SOILS AND PALAEOSOLS IN GLACIAL ENVIRONMENTS
FIGURE 17.4 Stratigraphic section illustrating catena of Farmdale Sangamon soil complex at a site in northern Illinois described by Jacobs et al. (2009), modified from Fig. 3 in that paper. Equality Formation is proglacial lake sediment deposited in MIS 2, Robein silt is aeolian and colluvial sediment (MIS 3), and Pearl Formation is glacigenic sediment from MIS 6. stream network (Ruhe, 1967). Palaeosols toward the margins of the divides are often reddish brown and relatively thin, probably the effect of episodic erosion and better drainage in that setting, and palaeosols are not preserved on the steeper slopes descending to stream valleys. In contrast, palaeo- sols in the middle of broad tabular divides are thick, clay-rich, and grey-coloured and display evi- dence of multiple phases of upward growth through additions of loess to the developing soil profile (Woida and Thompson, 1993).
17.3 PALAEOSOLS IN GLACIAL SETTINGS 17.3.1 PALAEOSOL PRESERVATION In rare cases it is possible to trace a soil across a former ice margin, from where it is the surface soil to where it is a palaeosol buried under till (Matthews and Caseldine, 1987). More generally, however, specific conditions are probably required for preservation of palaeosols overridden by ice or exposed to periglacial conditions just outside the ice margin. Multiple palaeosols can be pre- served where there is long-term net accumulation of glacigenic sediment. For example, stacks of multiple Late Pliocene and Pleistocene tills separated by palaeosols are exposed near the lower lim- its reached by montane glaciation in northwestern Canada and in Montana, USA (Karlstrom, 1988; Duk-Rodkin and Barendregt, 2011). Palaeosol preservation may also have implications for glacier bed conditions. Rovey and McLouth (2015) argue that the common preservation of palaeosols within a thick sequence of pre-Illinoian (Early and Middle Pleistocene) tills in the central United States, and the scarcity of deformation structures within those palaeosols, indicate general absence of deforming bed conditions. They also argue that the clay-rich nature of the palaeosols and the tills 17.3 PALAEOSOLS IN GLACIAL SETTINGS 595
they formed in prevented drainage of subglacial water and enhanced basal sliding. Piotrowski et al. (2001) argued against extensive deforming bed conditions under ice of the last glaciation in north- ern Germany, based on widespread preservation of the underlying Eemian soil. Karlstrom and Osborn (1992) interpreted features of Holocene palaeosols in western Canada as recording episodic colluvial deposition and disturbance as an alpine glacier advanced over them. Burial of a palaeosol by proglacial sediment before the ice advances over it is another mecha- nism that can explain preservation in some cases. For example, Jacobs et al. (2009) found that the entire pedogenetic profile of the Sangamon soil was preserved in northern Illinois, probably because the landscape was buried under proglacial lake sediment ahead of the advancing MIS 2 glacial ice (Fig. 17.4). The greatest preservation potential is associated with loess sedimentation (Fig. 17.1A), which can bury a former landscape and the soils on it intact (e.g., Follmer, 1982). In the Midwestern United States outside ice limits of MIS 2, both poor preservation of older soils and anomalously thin loess cover have been used as evidence of widespread erosion in the periglacial environment (Ruhe et al., 1968; Mason, 2015). Recognition that the erosion indicated by a missing soil or palaeosol where one would be expected can also be important in interpreting glacial stratigraphy (Ruhe et al., 1968; Meer et al., 2011). Evidence of palaeosols being eroded and incorporated into till deposits includes increased organic carbon content of tills (Dredge and McMartin, 2011) or intact blocks of soil and logs in basal tills (Johnson and Hansel, 1999). Clay-enriched B horizons are often the most readily recog- nized part of a palaeosol profile, and may also be more resistant to subglacial or hillslope erosion (Catt, 1990). The apparent absence of easily recognizable A or E horizons often leads to the inter- pretation of partial erosion before burial. It should be emphasized, however, that palaeosol A hori- zons may be difficult to recognize because of postburial compaction and oxidation of humus, obscuring pedogenic structure and lightening originally dark colours (Fig. 17.1B and C).
17.3.2 PALAEOSOL RECOGNITION Recognition of palaeosols in glacial settings is primarily based on soil morphology and mineral weathering features. Catt (1990) provides a thorough review of the issues involved. Palaeosols should display a sequence of horizons and other pedogenic features at least approximately similar to soils of about the same age, in similar environments of soil formation. Micromorphological fea- tures observed in thin sections can often be diagnostic of specific pedogenic processes, such as translocation of clay through soil pores, structure formation, and soil faunal activity (Stoops, 2003) (Fig. 17.5). While similar features such as coatings of translocated clay can sometimes occur in sediments unaltered by pedogenesis (Meer and Menzies, 2011), a combination of micro- and macromorphological evidence for multiple paedogenic processes is diagnostic of a soil. This use of modern analogues has historically been the dominant method in recognizing palaeosols and interpreting palaeoenvironments from them, and remains an essential first step. In fact, palaeosols representing substantial pedogenesis through one or more interglacials are often easily recognized in the field. The distinctive reddish-brown buried soils that formed in well-drained settings during and just after the last interglacial in North America and Eurasia are a good example (Fig. 17.1). Most disputes over palaeosol recognition involve much more weakly developed or highly truncated or deformed profiles. FIGURE 17.5 Photomicrographs of vertical thin sections of soil horizon features in Farmdale Sangamon soil complex in Northern Illinois. (A) A horizon of MIS 3 Farmdale soil with moderately separated granular microstructure with some tendency toward platy structure, probably as a result of compression; note common humus staining throughout; opaque areas are Fe/Mn oxides that may have precipitated around organic materials (plane polarized light, PPL). (B) Bg horizon of Farmdale soil that in outcrop appears massive, but in thin section has microstructure that appears to consist of compressed humus-stained granules (gr) separated by unstained silt (si) (PPL). (C) Bt horizon of Sangamon soil showing well-oriented illuvial clay lining channels (yellow-coloured areas with striated or laminated appearance, upper part of photo) that was translocated from overlying horizons and a fine matrix containing yellow stippled areas marking clay domains that likely originated from in situ weathering of mica and other minerals (cross-polarized light). Jacobs, P.M., Konen, M.E., Curry, B.B., 2009. Pedogenesis of a catena of the Farmdale–Sangamon Geosol complex in the north central United States. Palaeogeogr. Palaeoclimatol. Palaeoecol. 282, 119–132. 17.3 PALAEOSOLS IN GLACIAL SETTINGS 597
Palaeoenvironmental interpretation of palaeosols is another matter. A major difficulty in the use of modern analogues for that purpose is that both palaeosols and potential analogues at the modern land surface may often be polygenetic (Johnson et al., 1990; Catt, 1991), in the sense that they have experienced multiple changes in climate and vegetation, and may also have been affected by slow erosion even on relatively stable landscape positions or alternatively, experienced slow upbuilding by loess accumulation (e.g., Woida and Thompson, 1993). Evidence of mineral weathering is an important complement to morphological evidence in inter- preting palaeosols. In humid regions where a positive water balance leads to long-term leaching of solutes, the effects include dissolution of carbonate minerals, hydrolysis of feldspar and ferromag- nesian minerals to produce phyllosilicates and oxides, and the progressive alteration and loss of layer charge in some phyllosilicate minerals inherited from parent materials (Curtis, 1990). As dis- cussed above, progressive mineral weathering over time is evident in many chronosequences of soils formed in glacial sediment. Importantly, studies of the Sangamon soil in the central United States reveal large variations in the depth of carbonate dissolution and the degree of silicate mineral weathering—implying quite different weathering rates—related to the lithology of the glacigenic parent material; grain-size-related differences in permeability are likely to be especially important (Brophy, 1959; Jacobs, 1998a). Wilson (1999) reviews extensive research on alteration of the detrital clay mineral assemblage over time in soils. In an oxidizing and leaching environment, the sequence generally consists of rapid alteration of chlorite, followed by the loss of interlayer K from micas to form vermiculite, while further loss of interlayer cations and layer charge produce smectite under certain conditions. In poorly drained soils where base cations are abundant, the degree of alteration is less and neofor- mation of smectite may occur. The formation of kaolinite from detrital 2:1 clay minerals appears to require significant lengths of time ( . 10 kilo years) and often proceeds through a mixed-layer kaolinite-smectite phase (Hughes et al., 1993).
17.3.3 PALAEOSOLS IN GLACIAL SETTINGS AS EVIDENCE OF PALAEOENVIRONMENTS Early pedologists strongly emphasized the importance of climate as the key factor determining global soil geography (e.g., Glinka, 1914) and even today many introductory courses take this approach. However, some key features of soil morphology that are well-preserved after burial have a fairly complex relationship to climate. For example, a soil profile in which clay-poor A and/or E horizons overlie a clay-enriched B horizon is commonly associated with forests of humid midlati- tude regions, but can also be observed in semiarid grasslands or shrublands. On the other hand, where forests bordered on subhumid grasslands during the Holocene, as in central North America and Russia, a sharp pedologic boundary formed, separating soils with much greater A to B horizon textural contrast on the forest side from those with much less on the grassland side. Thick A hori- zons with strong granular structure formed under grassland also contrast with thin A horizons on the forest side (Fig. 17.6)(White and Riecken, 1955; Bailey et al., 1964; Severson and Arneman, 1973; Bronger, 1991; Miedema et al., 1999). Pedogenic carbonate accumulation in the B horizon is especially common in semiarid to arid regions, but can occur locally in more humid settings (Schaetzl and Anderson, 2005, p. 402). 598 CHAPTER 17 SOILS AND PALAEOSOLS IN GLACIAL ENVIRONMENTS
FIGURE 17.6 Contrast between soils formed in till (MIS 2) under forest and grassland near the transition between those vegetation types in northwestern Minnesota (47 degrees North, 96 degrees West). (A) Typical morphology, with thin A horizon and distinct E horizon over clay-rich Bt horizon under forest and thicker A over weak Bw horizon under grassland. (B) Clay content profiles of four soils including those in (A), showing large clay increase from A and E to B in forest soils, but not in prairie soils. All soils are within a 40-km radius in essentially identical climates.
A limited number of soil features may be more diagnostic of specific climatic conditions. These certainly include ice-wedge casts and relict sand wedges indicative of past permafrost, and involu- tions or other evidence of cryoturbation in a cold climate (Kemp, 1987; Catt, 1990; Kemp et al., 1993; Mason et al., 1994; Dampier et al., 2009). These features have been identified in palaeosols and sediments as old as the Precambrian, and have been used as evidence for palaeoclimatic condi- tions associated with pre-Quaternary glaciation (Williams, 1986; Retallack, 1999). Podzolization (accumulation of iron, aluminium, and humus in the B horizon) is largely associated with forest or heath vegetation in humid climates (Schaetzl and Anderson, 2005, pp. 440 441). Many other important pedogenic processes are not inherently diagnostic of particular climatic conditions; however, they would often be expected to occur at different rates depending on the cli- mate, especially through precipitation effects on the flux of water through the soil profile, and tem- perature effects on silicate mineral weathering kinetics. As noted above, chronosequence studies do not always support this assumption, but it underlies longstanding, recurrent efforts to interpret palaeosols as representing interglacial climates that were warmer and/or wetter than the Holocene, because they have thicker, more clay-rich, and/or redder profiles than comparable Holocene soils 17.3 PALAEOSOLS IN GLACIAL SETTINGS 599
(e.g., Karlstrom, 1991). Chronosequence studies demonstrate that those soil properties are also a function of the duration of pedogenesis, so disentangling effects of climate and soil age is a crucial problem for interpreting these well-developed interglacial palaeosols (Boardman, 1985). The duration of pedogenesis was clearly longer for many interglacial palaeosols than for surface soils formed since the last deglaciation in the same environments. For example, the Sangamon soil in the Midwestern United States was at the land surface undergoing pedogenesis not only during the last interglacial in a strict sense (MIS 5), but also during the buildup of the last glaciation (MIS 4). Burial of this palaeosol began with loess deposition around 55 ka, but it was still near the surface until as late as 25 30 ka in many localities; thus, pedogenesis spans at least 75,000 years (Curry and Follmer, 1992; Curry and Pavich, 1996). Similar constraints from stratigraphy and numerical dating are available for major palaeosols in Eurasia (Fischer et al., 2012; Kreutzer et al., 2012). From these data, it might in theory be possible to use chronosequence studies to evaluate whether the degree of morphological development and weathering in interglacial palaeosols required a warmer or wetter environment than the Holocene. In practice, the scatter observed in chronosequence data and the substantial effects of parent material lithology and local landscape set- ting on pedogenic processes limit this approach. Jacobs (1998a) noted that the depth of carbonate leaching, solum thickness, and B horizon reddening in the Sangamon soil varied substantially over a small area of Indiana, USA, as a function of the sand content of the various glacigenic sediment facies that the palaeosol formed in. If various individual profiles sampled in this study were inter- preted in isolation, the inferred palaeoclimate might vary widely. To minimize parent material effects, Catt (1988) recommended comparison of palaeosols within relatively uniform loess successions to gauge climate evolution through the Quaternary. In his anal- ysis from 20 locations in Europe and Asia, Catt used modern analogues to rank interstadial and interglacial palaeosols according to inferred environment of formation, from cold and dry to warm and wet. In a till sequence in northern Missouri, USA, Rovey and Balco (2015) describe Early to Late Pleistocene palaeosols formed in six tills and overlying loess. They used CRN to determine the duration of exposure and relate that to soil morphological features such as soil colour, B hori- zon thickness, depth of carbonate mineral weathering, and secondary carbonate accumulation. A change from secondary carbonate accumulation in Early Pleistocene palaeosols to carbonate-free profiles by the Middle to Late Pleistocene was interpreted as indicating an increase in mean annual precipitation. Redder colours, despite shorter times of soil formation for the Yarmouth and Sangamon soils, formed after about 0.4 Ma, were interpreted to indicate wetter and/or warmer inter- glacial conditions than the older soils formed in. The study by Rovey and Balco (2015) points to the value of using CRN analysis to quantify the duration of surface exposure, as an important step toward distinguishing effects of climate and time on palaeosol properties. We would also emphasize the need to characterize local variation of each palaeosol across the palaeolandscape it formed on, and across a range of parent materials. Studying geographic variation in palaeosols that are preserved over large areas, such as the Sangamon soil, is also a potentially fruitful line of research that deserves more attention. The large morphological changes observed in modern surface soils along transects from humid forest-covered regions to semiarid grasslands—and the particularly abrupt change at the forest grassland transition (Fig. 17.6)—should be reproduced in a palaeosol that spans the same geographic extent, but it may be shifted in a direction indicating warmer or wetter than modern climatic conditions. Ruhe (1965) 600 CHAPTER 17 SOILS AND PALAEOSOLS IN GLACIAL ENVIRONMENTS
suggested that comparison of Sangamon and modern soils in the Midwestern USA may indicate only minor shifts in climate and vegetation zones. In Ukraine and southern Russia, however, Velichko et al. (1984) used the morphology of the Mikulino (last interglacial) soil to infer forest expansion and a warmer and wetter climate than the present one. Micromorphological analysis of palaeosols in the Loess Plateau of China also found evidence for different forest steppe boundary locations during Middle and Late Pleistocene interglacials (Bronger and Heinkele, 1989).
17.4 CONCLUSIONS With adequate attention to the complexities of soil formation over time and across glaciated land- scapes, both surface soils and buried palaeosols can provide valuable information on the relative age of glacial sediments and landforms, patterns of postglacial landscape evolution, and the palaeoenvironments in which they formed. Chronosequences of soils formed in glacigenic sediment not only provide vital information on processes and rates of soil formation, but are also the basis for using soil morphology as a tool in mapping and correlating glacial sediments and landforms, and in guiding the application of numerical dating methods. Studies of surface soils—not only chronosequences but also investigation of variation with landscape position and parent material lithology—are essential background for palaeoenvironmental and geomorphic interpretation of palaeosols in glacial settings.
REFERENCES
Almond, P., Tonkin, P., 1999. Pedogenesis by upbuilding in an extreme leaching and weathering environment, and slow loess accretion, south Westland, New Zealand. Geoderma 92, 1 36. Applegarth, M.T., Dahms, D.E., 2004. Aeolian modification of moraine soils, Whiskey Basin, Wyoming, USA. Earth Surf. Process. Landforms 29, 579 585. Bailey, L., Odell, R., Boggess, W., 1964. Properties of selected soils developed near the forest-prairie border in east-central Illinois. Soil Sci. Soc. Am. Proc. 28, 257 263. Berry, M.E., 1987. Morphological and chemical characteristics of soil catenas on Pinedale and Bull Lake moraine slopes in the Salmon River Mountains, Idaho. Quat. Res. 28, 210 225. Berry, M.E., 1994. Soil-geomorphic analysis of late-Pleistocene glacial sequences in the McGee, Pine, and Bishop Creek drainages, east-central Sierra Nevada, California. Quat. Res. 41, 160 175. Birkeland, P.W., 1964. Pleistocene glaciation of the northern Sierra Nevada, north of Lake Tahoe, California. J. Geol. 810 825. Birkeland, P.W., 1984. Holocene soil chronofunctions, Southern Alps, New Zealand. Geoderma 34, 115 134. Birkeland, P.W., 1990. Soil-geomorphic research—a selective overview. Geomorphology 3, 207 224. Birkeland, P.W., 1994. Variation in soil-catena characteristics of moraines with time and climate, South Island, New Zealand. Quat. Res. 42, 49 59. Birkeland, P.W., Burke, R.M., Benedict, J.B., 1989. Pedogenic gradients for iron and aluminum accumulation and phosphorus depletion in arctic and alpine soils as a function of time and climate. Quat. Res. 32, 193 204. Boardman, J., 1985. Comparison of soils in midwestern United States and western Europe with the interglacial record. Quat. Res. 23, 62 75. REFERENCES 601
Bockheim, J., 1980. Solution and use of chronofunctions in studying soil development. Geoderma 24, 71 85. Bockheim, J., 2013. Soil formation in the Transantarctic Mountains from the middle Paleozoic to the Anthropocene. Palaeogeogr. Palaeoclimatol. Palaeoecol. 381, 98 109. Bockheim, J.G., Douglass, D.C., 2006. Origin and significance of calcium carbonate in soils of southwestern Patagonia. Geoderma 136, 751 762. Brimhall, G.H., Dietrich, W.E., 1987. Constitutive mass balance relations between chemical composition, vol- ume, density, porosity, and strain in metasomatic hydrochemical systems: results on weathering and pedo- genesis. Geochim. Cosmochim. Acta 51, 567 587. Bronger, A., 1991. Argillic horizons in modern loess soils in an ustic soil moisture regime: comparative studies in forest-steppe and steppe areas from Eastern Europe and the United States. Adv. Soil Sci. 15, 41 90. Bronger, A., Heinkele, T., 1989. Micromorphology and genesis of paleosols in the Luochuan loess section, China: pedostratigraphic and environmental implications. Geoderma 45, 123 143. Brophy, J.A., 1959. Heavy mineral ratios of Sangamon weathering profiles in Illinois, Illinois State Geological Survey Circular, vol. 273. pp. 1 22. Burke, R., Birkeland, P.W., 1979. Reevaluation of multiparameter relative dating techniques and their application to the glacial sequence along the eastern escarpment of the Sierra Nevada, California. Quat. Res. 11, 21 51. Burkins, D.L., Blum, J.D., Brown, K., Reynolds, R.C., Erel, Y., 1999. Chemistry and mineralogy of a granitic, glacial soil chronosequence, Sierra Nevada Mountains, California. Chem. Geol. 162, 1 14. Burras, C.L., Scholtes, W.H., 1987. Basin properties and postglacial erosion rates of minor moraines in Iowa. Soil Sci. Soc. Am. J. 51, 1541 1547. Available from: http://dx.doi.org/10.2136/sssaj1987.03615995005100060025x. Burt, R., Alexander, E., 1996. Soil development on moraines of Mendenhall Glacier, southeast Alaska. 2. Chemical transformations and soil micromorphology. Geoderma 72, 19 36. Catt, J., 1988. Soils of the Plio-Pleistocene: do they distinguish types of interglacial? Philos. Trans. R. Soc. B Biol. Sci. 318, 539 557. Catt, J.A., 1990. Paleopedology manual. Quat. Intl. 6, 1 95. Catt, J.A., 1991. Soils as indicators of Quaternary climatic change in mid-latitude regions. Geoderma 51, 167 187. Chadwick, O.A., Brimhall, G.H., Hendricks, D.M., 1990. From a black to a gray box—a mass balance inter- pretation of pedogenesis. Geomorphology 3, 369 390. Coleman, T., Fenton, T., 1982. Stratigraphy and Water Table Relationships of Upland Loess-Derived Soils in South-Central Iowa. Soil Sci. Soc. Am. J. 46, 82 87. Colman, S.M., 1981. Rock-weathering rates as functions of time. Quat. Res. 15, 250 264. Colman, S.M., Pierce, K.L., 1986. Glacial sequence near McCall, Idaho: weathering rinds, soil development, morphology, and other relative-age criteria. Quat. Res. 25, 25 42. Crocker, R.L., Major, J., 1955. Soil development in relation to vegetation and surface age at Glacier Bay, Alaska. J. Ecol. 43, 427 448. Curry, B.B., Follmer, L.R., 1992. The last interglacial-glacial transition in Illinois: 123 25 ka. In: Clark, P.U., Lea, P.D. (Eds.), The Last Interglacial-Glacial Transition in North America. Special. Geological Society of America, Boulder, CO, pp. 71 88. , Paper 270. Curry, B.B., Pavich, M.J., 1996. Absence of glaciation in Illinois during marine isotope stages 3 through 5. Quat. Res. 46, 19 26. Curtis, C., 1990. Aspects of climatic influence on the clay mineralogy and geochemistry of soils, palaeosols and clastic sedimentary rocks. J. Geol. Soc. 147, 351 357. Dahms, D., Favilli, F., Krebs, R., Egli, M., 2012. Soil weathering and accumulation rates of oxalate-extractable phases derived from alpine chronosequences of up to 1Ma in age. Geomorphology 151, 99 113. Dahms, D.E., 1993. Mineralogical evidence for eolian contribution to soils of late Quaternary moraines, Wind River Mountains, Wyoming, USA. Geoderma 59, 175 196. 602 CHAPTER 17 SOILS AND PALAEOSOLS IN GLACIAL ENVIRONMENTS
Dahms, D.E., 1994. Mid-Holocene erosion of soil catenas on moraines near the type Pinedale Till, Wind River Range, Wyoming. Quat. Res. 42, 41 48. Dahms, D.E., 2002. Glacial stratigraphy of Stough Creek Basin, Wind River Range, Wyoming. Geomorphology 42, 59 83. Dahms, D.E., 2004. Relative and numeric age data for glacial deposits and diamictons in and near Sinks Canyon, Wind River Range, Wyoming, U.S.A. Arct. Antarct. Alp. Res. 36, 59 77. Dampier, L., Sanborn, P., Bond, J., Clague, J.J., Smith, S., 2009. Soil genesis in relation to glacial history in central Yukon. In: Weston, L.H., Blackburn, L.R., Lewis, L.L. (Eds.), Yukon Exploration and Geology 2008. Yukon Geological Survey, pp. 113 123. Douglass, D.C., Bockheim, J.G., 2006. Soil-forming rates and processes on Quaternary moraines near Lago Buenos Aires, Argentina. Quat. Res. 65, 293 307. Dredge, L.A., McMartin, I., 2011. Glacial stratigraphy of northern and central Manitoba. Geological Survey of Canada Bulletin 600. Natural Resources Canada. Duk-Rodkin, A., Barendregt, R., 2011. Stratigraphical record of glacials/interglacials in northwest Canada. In: Ehlers, J., Gibbard, P.L., Hughes, P.D. (Eds.), Quaternary Glaciations—Extent and Chronology. A Closer Look. Developments in Quaternary Science, vol. 15. Elsevier, Amsterdam, p. 49. Egli, M., Mirabella, A., Sartori, G., Zanelli, R., Bischof, S., 2006a. Effect of north and south exposure on weathering rates and clay mineral formation in Alpine soils. Catena 67, 155 174. Egli, M., Wernli, M., Kneisel, C., Haeberli, W., 2006b. Melting glaciers and soil development in the pro- glacial area Morteratsch (Swiss Alps): I. Soil type chronosequence. Arct. Antarct. Alp. Res. 38, 499 509. Egli, M., Mirabella, A., Sartori, G., 2008. The role of climate and vegetation in weathering and clay mineral formation in late Quaternary soils of the Swiss and Italian Alps. Geomorphology 102, 307 324. Fischer, P., Hilgers, A., Protze, J., Kels, H., Lehmkuhl, F., Gerlach, R., 2012. Formation and geochronology of Last Interglacial to Lower Weichselian loess/palaeosol sequences case studies from the Lower Rhine Embayment, Germany. Eiszeitalter Ggw. Quat. Sci. J. 61, 48 63. Follmer, L., 1978. The Sangamon soil in its type area—a review. In: Mahaney, W. (Ed.), Quaternary Soils. GeoAbstracts, Norwich, p. 165. Follmer, L.R., 1982. The geomorphology of the Sangamon surface: its spatial and temporal attributes. In: Thorn, C.E. (Ed.), Space and Time in Geomorphology. Allen and Unwin, London, pp. 117 146. Follmer, L.R., 1983. Sangamon and Wisconsinan pedogenesis in the Midwestern United States. In: Porter, S.C., Wright Jr, H. (Eds.), The Late Pleistocene. University of Minnesota Press, Minneapolis, MN, pp. 138 144. Glinka, K.D., 1914. Typen der Bodenbildung, ihre Klassifikation und geographische Verbreitung. Gebruder Borntraeger, Berlin. Gosse, J.C., Klein, J., Lawn, B., Middleton, R., Evenson, E., 1995. Beryllium-10 dating of the duration and retreat of the last Pinedale glacial sequence. Science 268, 1329 1333. Grimley, D.A., Follmer, L.R., Hughes, R.E., Solheid, P.A., 2003. Modern, Sangamon and Yarmouth soil devel- opment in loess of unglaciated southwestern Illinois. Quat. Sci. Rev. 22, 225 244. Hall, R.D., 1999. Effects of climate change on soils in glacial deposits, Wind River Basin, Wyoming. Quat. Res. 51, 248 261. Hall, R.D., Anderson, A.K., 2000. Comparative soil development of Quaternary paleosols of the central United States. Palaeogeogr. Palaeoclimatol. Palaeoecol. 158, 109 145. Hall, R.D., Shroba, R.R., 1993. Soils developed in the glacial deposits of the type areas of the Pinedale and Bull Lake glaciations, Wind River Range, Wyoming, USA. Arct. Alp. Res. 25, 368 373. Harden, J.W., Taylor, E.M., 1983. A quantitative comparison of soil development in four climatic regimes. Quat. Res. 20, 342 359. REFERENCES 603
Hughes, R., Moore, D., Reynolds, R., 1993. The nature, occurrence, and origin of kaolinite/smectite. In: Murray, H.H., Bundy, W.M., Harvey, C.C. (Eds.), Kaolin Genesis and Utilization. Clay Mineral Society Special Publication 1, pp. 291 323. Jacobs, P.M., 1998a. Influence of parent material grain size on genesis of the Sangamon Geosol in south- central Indiana. Quat. Intl. 51, 127 132. Jacobs, P.M., 1998b. Morphology and weathering trends of the Sangamon soil complex in south-central Indiana in relation to paleodrainage. Quat. Res. 50, 221 229. Jacobs, P.M., Konen, M.E., Curry, B.B., 2009. Pedogenesis of a catena of the Farmdale Sangamon Geosol complex in the north central United States. Palaeogeogr. Palaeoclimatol. Palaeoecol. 282, 119 132. Jacobson, G., Birks, H., 1980. Soil development on recent end moraines of the Klutlan Glacier, Yukon Territory, Canada. Quat. Res. 14, 87 100. James, L.A., 1988. Rates of organic carbon accumulation in young mineral soils near Burroughs Glacier, Glacier Bay, Alaska. Phys. Geogr. 9, 50 70. Jenny, H., 1941. Factors of Soil Formation: A System of Quantitative Pedology. McGraw-Hill, New York, NY, p. 281. Johnson, D.L., Keller, E.A., Rockwell, T.K., 1990. Dynamic pedogenesis: new views on some key soil con- cepts, and a model for interpreting Quaternary soils. Quat. Res. 33, 306 319. Johnson, W.H., Hansel, A.K., 1999. Wisconsin Episode glacial landscape of central Illinois: a product of sub- glacial deformation processes? In: Mickelson, D.M., Attig, J.W. (Eds.), Glacial Processes Past and Present. Geological Society of America Special Paper 337. Boulder, Colorado, pp. 121 135. Karlstrom, E.T., 1988. Multiple paleosols in pre-Wisconsinan drift, northwestern Montana and southwestern Alberta. Catena 15, 147 178. Karlstrom, E.T., 1991. Paleoclimatic significance of Late Cenozoic paleosols east of Waterton—Glacier Parks, Alberta and Montana. Palaeogeogr. Palaeoclimatol. Palaeoecol. 85, 71 100. Karlstrom, E.T., Osborn, G., 1992. Genesis of buried paleosols and soils in Holocene and late Pleistocene tills, Bugaboo Glacier area, British Columbia, Canada. Arct. Alp. Res. 24, 108 123. Kemp, R.A., 1987. Genesis and environmental significance of a buried Middle Pleistocene soil in eastern England. Geoderma 41, 49 77. Kemp, R.A., Whiteman, C., Rose, J., 1993. Palaeoenvironmental and stratigraphic significance of the Valley Farm and Barham Soils in eastern England. Quat. Sci. Rev. 12, 833 848. Kreutzer, S., Fuchs, M., Meszner, S., Faust, D., 2012. OSL chronostratigraphy of a loess-palaeosol sequence in Saxony, Germany using quartz of different grain sizes. Quat. Geochronol. 10, 102 109. Kukla, G., 2005. Saalian supercycle, Mindel/Riss interglacial and Milankovitch’s dating. Quat. Sci. Rev. 24, 1573 1583. Leverett, F., 1898. The weathered zone (Sangamon) between the Iowan loess and Illinoian till sheet. J. Geol. 6, 171 181. Machette, M.N., 1985. Calcic soils of the southwestern United States. In: Weide, D.L. (Ed.), Soils and Quaternary Geology of the Southwestern United States, Special Paper 203. Geological Society of America, Boulder, CO, pp. 1 22. Mahaney, W., 1978. Late Quaternary Stratigraphy and Soils in the Wind River Mountains, Western Wyoming, Quaternary Soils. Geo Abstracts, Norwich, pp. 223 264. Mahaney, W., Kalm, V., Kapran, B., Milner, M., Hancock, R., 2009. A soil chronosequence in Late Glacial and Neoglacial moraines, Humboldt Glacier, northwestern Venezuelan Andes. Geomorphology 109, 236 245. Mason, J.A., 2015. Up in the refrigerator: geomorphic response to periglacial environments in the Upper Mississippi River Basin, USA. Geomorphology 248, 363 381. Mason, J.A., Milfred, C.J., Nater, E.A., 1994. Distinguishing soil age and parent material effects on an Ultisol of north-central Wisconsin, USA. Geoderma 61, 165 189. 604 CHAPTER 17 SOILS AND PALAEOSOLS IN GLACIAL ENVIRONMENTS
Matthews, J.A., Caseldine, C.J., 1987. Arctic-alpine brown soils as a source of palaeoenvironmental informa- tion: further 14C dating and palynological evidence from Vestre Memurubreen, Jotunheimen, Norway. J. Quat. Sci. 2, 59 71. Mavris, C., Egli, M., Plo¨tze, M., Blum, J.D., Mirabella, A., Giaccai, D., et al., 2010. Initial stages of weathering and soil formation in the Morteratsch proglacial area (Upper Engadine, Switzerland). Geoderma 155, 359 371. McDonald, E.V., Busacca, A.J., 1990. Interaction between aggrading geomorphic surfaces and the formation of a Late Pleistocene paleosol in the Palouse loess of eastern Washington state. Geomorphology 3, 449 469. Meer, J.J.M. van der., Menzies, J., 2011. The micromorphology of unconsolidated sediments. Sedimentary Geology 238, 213 232. Meer, J.J.M. van der., Thomas, G.S.P., Chiverall, R.C., Lee, J., Phillips, E., Rijsdijk, K., et al., 2011. Macrostructural analysis: unraveling polyphase glacitectonic histories. In: Phillips, E., Lee, J., Evans, H.M. (Eds.), Glacitectonism: A Guide to the Description and Analysis of Glacitectonic Deformation Structures with Case Studies from Norfolk. Quaternary Research Association, England, pp. 54 100. Meierding, T., 1984. Correlation of Rocky Mountain Pleistocene deposits by relative dating methods: a perspective. In: Mahaney, W. (Ed.), Correlation of Quaternary Chronologies. Geo Books, Norwich, pp. 455 477. Miedema, R., Koulechova, I.N., Gerasimova, M.I., 1999. Soil formation in Greyzems in Moscow district: micromorphology, chemistry, clay mineralogy, and particle size distribution. Catena 34, 315 347. Milne, G., 1936. Normal erosion as a factor in soil profile development. Nature 138, 548 549. Mokma, D., Syers, J., Jackson, M., Clayton, R., Rex, R., 1972. Aeolian additions to soils and sediments in the South Pacific area. J. Soil Sci. 23, 147 162. Penck, A., Bru¨ckner, E., 1909. Die alpen im Eiszeitalter, vol. 3. Tauchnitz, Leipzig. Phillips, F.M., Zreda, M.G., Smith, S.S., Elmore, D., Kubik, P.W., Sharma, P., 1990. Cosmogenic chlorine-36 chronology for glacial deposits at Bloody Canyon, eastern Sierra Nevada. Science 248, 1529 1532. Phillips, F.M., Zreda, M.G., Gosse, J.C., Klein, J., Evenson, E.B., Hall, R.D., et al., 1997. Cosmogenic 36Cl and 10Be ages of Quaternary glacial and fluvial deposits of the Wind River Range, Wyoming. Geol. Soc. Am. Bull. 109, 1453 1463. Piotrowski, J.A., Mickelson, D.M., Tulaczyk, S., Krzyszkowski, D., Junge, F.W., 2001. Were deforming sub- glacial beds beneath past ice sheets really widespread? Quat. Intl. 86, 139 150. Retallack, G.J., 1999. Carboniferous fossil plants and soils of an early tundra ecosystem. Palaios 14, 324 336. Richmond, G.M., 1965. Glaciation of the Rocky Mountains. In: Wright Jr, H.E., Frey, D.G. (Eds.), The Quaternary of the United States. Princeton University Press, Princeton, NJ, pp. 217 230. Rovey, C.W., Balco, G., 2015. Paleoclimatic interpretations of buried paleosols within the pre-Illinoian till sequence in northern Missouri, USA. Palaeogeogr. Palaeoclimatol. Palaeoecol. 417, 44 56. Ruhe, R.V., 1965. Quaternary paleopedology. In: Wright Jr., H.E., Frey, D.G. (Eds.), The Quaternary of the United States. Princeton University Press, Princeton, New Jersey, pp. 755 764. Rovey, C.W., McLouth, T., 2015. A near synthesis of pre-Illinoian till stratigraphy in the central United States: Iowa, Nebraska and Missouri. Quat. Sci. Rev. 126, 96 111. Ruhe, R.V., 1952. Topographic discontinuities of the Des Moines lobe. Am. J. Sci. 250, 46 56. Ruhe, R.V., 1967. Geomorphology of parts of the Greenfield Quadrangle, Adair County, Iowa; chapter 2. In: Ruhe, R.V., Daniels, R.B., Cady, J.G. (Eds.), Landscape Evolution and Soil Formation in Southwestern Iowa. Technical Bulletin 1349, U. S. Department of Agriculture, Washington, DC, pp. 93 161. Ruhe, R.V., Dietz, W.P., Fenton, T.E., Hall, G.F., 1968. Iowan Drift Problem, Northeastern Iowa. Report of Investigations 7. Iowa Geological Survey, Iowa City. Runge, E., 1973. Soil development sequences and energy development models. Soil Sci. 115, 183 193. Sauer, D., Schu¨lli-Maurer, I., Wagner, S., Scarciglia, F., Sperstad, R., Svendga˚rd-Stokke, S., et al., 2015. Soil development over millennial timescales-a comparison of soil chronosequences of different climates and REFERENCES 605
lithologies. IOP Conference Series: Earth and Environmental Science. IOP Publishing, Bristol, England, p. 012009. Schaetzl, R., Anderson, S., 2005. Soils. Genesis and Geomorphology. Cambridge University Press, Cambridge. Schaetzl, R., Barrett, L., Winkler, J., 1994. Choosing models for soil chronofunctions and fitting them to data. Eur. J. Soil Sci. 45, 219 232. Schaller, M., Blum, J., Ehlers, T., 2009. Combining cosmogenic nuclides and major elements from moraine soil profiles to improve weathering rate estimates. Geomorphology 106, 198 205. Schaller, M., Blum, J.D., Hamburg, S.P., Vadeboncoeur, M.A., 2010. Spatial variability of long-term chemical weathering rates in the White Mountains, New Hampshire, USA. Geoderma 154, 294 301. Seelig, B., Richardson, J., 1994. Sodic soil toposequence related to focused water flow. Soil Sci. Soc. Am. J. 58, 156 163. Severson, R., Arneman, H., 1973. Soil characteristics of the forest-prairie ecotone in northwestern Minnesota. Soil Sci. Soc. Am. Proc. 37, 593 599. Simonson, R.W., 1959. Outline of a generalized theory of soil genesis. Soil Sci. Soc. Am. Proc. 23, 152 156. Sommer, M., Schlichting, E., 1997. Archetypes of catenas in respect to matter—a concept for structuring and grouping catenas. Geoderma 76, 1 33. Stevens, P., Walker, T., 1970. The chronosequence concept and soil formation. Q. Rev. Biol. 45, 333 350. Stoops, G., 2003. Guidelines for analysis and description of soil and regolith thin sections. Soil Science Society of America, Madison, Wisconsin, 184 p. Swanson, D.K., 1985. Soil catenas on Pinedale and Bull Lake moraines, Willow Lake, Wind River Mountains, Wyoming. Catena 12, 329 342. Taylor, A., Blum, J.D., 1995. Relation between soil age and silicate weathering rates determined from the chemical evolution of a glacial chronosequence. Geology 23, 979 982. Velichko, A.A., Bogucki, A.B., Morozova, T.D., Udartsev, V.P., Khalcheva, T.A., Tsatskin, A.I., 1984. Periglacial landscapes of the East European Plain. In: Velichko, A.A., Wright Jr, H.E., Barnosky, C. (Eds.), Late Quaternary Environments of the Soviet Union. University of Minnesota, Minneapolis, MN. Vidic, N.J., 1998. Soil-age relationships and correlations: comparison of chronosequences in the Ljubljana Basin, Slovenia and USA. Catena 34, 113 129. Walker, P., Ruhe, R., 1968. Hillslope models and soil formation. II. Closed system, Transactions 9th International Congress of Soil Science, Adelaide, Australia, vol. 4. pp. 561 568. White, E.M., Riecken, F.F., 1955. Brunizem-Gray Brown Podzolic soil biosequences. Soil Sci. Soc. Am. Proc. 19, 504 509. Williams, G., 1986. Precambrian permafrost horizons as indicators of palaeoclimate. Precambrian Res. 32, 233 242. Wilson, M., 1999. The origin and formation of clay minerals in soils: past, present and future perspectives. Clay Miner. 34, 7 25. Woida, K., Thompson, M., 1993. Polygenesis of a Pleistocene paleosol in southern Iowa. Geol. Soc. Am. Bull. 105, 1445 1461. Worthen, A.H., 1873. Geology and Paleontology, vol. V. Geological Survey of Illinois.