Abstract: 

Paleosols formed from weathering of alluvial mudstones in the Late Cretaceous (Maastrichtian) Prince Creek Formation, North Slope Alaska, are dominated by detrital smectite, discrete illite, kaolinite, chlorite, quartz, and pedogenic illite–smectite (I/S) mixed-layer clays. In the fine clay fraction (< 0.2 µm) illite–smectite mixed-layer clay is the main clay mineral and is interpreted as pedogenic in origin, whereby the I/S is a product of illitization of inherited smectite during weathering and pedogenesis. We consider the detrital clay minerals to be derived from pre-existing sediments eroded from the Brooks Range, mixed with reworked volcanic ash-fall-derived bentonites. In the Prince Creek Formation, smectitic parent materials were deposited by epiclastic volcanic ash-rich alluvium that accumulated on imperfectly drained floodplains. Diagenetic transformation of smectite to illite is unlikely in the Prince Creek Formation, in as much as maximum burial temperatures never exceeded ∼ 48° C. The predominance of bentonite-derived smectite (> 80%), low bulk density, phosphorus accumulation, Fe and Al mass-balance trends, and the presence of Fe–Al–humus complexes in one paleosol profile is interpreted as evidence of andic soil properties, and these paleosols are interpreted, therefore, as Andept-like alluvial soils. These results demonstrate that clay mineralogical studies, in conjunction with geochemical data of paleosols, can be used to identify paleo-andic soil properties which have not been widely recognized in the ancient rock record. Alternating wetting and drying conditions, required to form pedogenic I/S in these alluvial paleosols, resulted from a highly seasonal moisture regime in the Late Cretaceous Arctic.

Introduction

Clay mineral studies of pre-Quaternary paleosols are increasingly common and can provide significant qualitative and, in some cases, quantitative information on paleoclimate and soil environmental properties, surface weathering characteristics, and diagenesis (Środoń 1999; Vitali et al. 2002; Huggett and Cuadros 2005; Sheldon and Tabor 2009). Clay minerals originate from a variety of different processes, but only those clay minerals that are the product of surface weathering may provide information about pedogenesis and paleoclimate (Singer 1980; Singer 1984).

Studies of clay minerals from high-latitude paleosols remain scarce (Kodama et al. 1976; Foscolos et al. 1977). The Prince Creek Formation, North Slope, Alaska, U.S.A., deposited at a paleolatitude of 83–85° N during the Late Cretaceous (Ziegler et al. 1983; Lawver et al. 2002), provides an exquisite opportunity to examine clay-mineral development in paleosols formed under paleo-arctic greenhouse climate conditions (Fig. 1). Importantly, a more refined understanding of surface processes operating under ancient, warm-Earth climate states has potentially significant societal implications as Earth undergoes a current phase of global warming (Montañez et al. 2007; Parrish et al. 2012).

In this paper, we identify the origin of the clay minerals in three paleosols and their parent materials, interpret the pedogenic processes (illitization and andosolization) acting on these ancient floodplains, and discuss the paleosol types present. The paleosols of the Prince Creek Formation, previously classified as Aquic Inceptisols (Flaig et al. 2013), are further refined herein as non-allophanic Andept-like paleosols on the basis of clay-mineral signatures and geochemical trends interpreted to represent evidence of andic soil-forming processes. Our results further refine our understanding of alluvial soil formation in a much warmer arctic, but one with a polar light regime similar to that experienced at high latitudes today (Spicer 2003; Spicer and Herman 2010; Tomsich et al. 2010).

Geologic Setting

Stratigraphy

The Prince Creek Formation, exposed along the Colville River, Alaska, is a Late Cretaceous (Maastrichtian) alluvial–deltaic coastal-plain succession that was deposited in the Colville Basin (Fig. 1). The Prince Creek Formation interfingers with, and is overlain by, marine and marginal-marine sediments of the Schrader Bluff Formation (Fig. 1; Mull et al. 2003; Phillips 2003).

The age of the entire Prince Creek Formation extends from Campanian to Paleocene (Conrad et al. 1990; Bice et al. 1996); however, recent biostratigraphic and geochronologic analyses from exposures along the Colville River (Fig. 1) indicate an early Maastrichtian age for all of the samples included in our data set (Brandlen 2008; Fiorillo et al. 2010a, 2010b; Flaig 2010; Flaig et al. 2011, 2013). The Prince Creek Formation was buried to depths of ∼ 600 to 2000 m (Burns et al. 2005) and vitrinite reflectance values (n  =  4, avg. 0.3825) suggest a maximum burial temperature of ∼ 48° C, which is too low to have caused diagenetic changes in paleosol mineralogy (Barker and Pawlewicz 1986; Robinson 1989; Johnson and Howell 1996).

Depositional Environments

The Prince Creek Formation consists of interbedded, very fine- to fine-grained sandstone, conglomerate, organic-rich siltstone, and coal with thin bentonites and tuffs, and abundant remains of both vertebrates and fossil plants (Mull et al. 2003). Detailed descriptions of sedimentary facies and alluvial architecture of the Prince Creek Fm. are presented elsewhere (Phillips 2003; Brandlen 2008; Fiorillo et al. 2010a, 2010b; Flaig et al. 2011). The Prince Creek Formation is interpreted as a large, tidally influenced, alluvial–deltaic depositional system consisting of large meandering trunk channels, smaller meandering and anastomosing distributary channels, and mud-rich floodplain deposits with paleosols, small lakes or ponds, and swamps (Phillips 2003; Fiorillo et al. 2010a, 2010b; Flaig et al. 2011, 2013). The depositional system grades northward from the Brooks Range, an arctic mountain range at least 1500 m high even in the Cretaceous (Spicer 2003), into deltaic distributary channels and marginal marine interdistributary bays and tidal flats (Phillips 2003; Flaig et al. 2011; Fiorillo et al. 2010a).

Floodplains and Paleosols.—

Alluvial plains in the Prince Creek Formation are dominated by fine-grained sediments reflecting a variety of floodplain sub-environments. Small sheet-like sandstones and siltstones are interpreted as crevasse splays and levees (Flaig et al. 2011). Thin packages of interbedded, ripple cross-laminated, rarely rooted or burrowed siltstone and mudstone are interpreted as small floodplain lakes and ponds (Flaig et al. 2011). Organic-rich mudstone, carbonaceous shale, and coal (typically < 0.5 m thick) record organic deposition in swamps or on poorly drained floodplains. Thin (< 1 m thick) tuffs and bentonites are locally present (Flaig et al. 2011). Drab-colored, rooted and mottled siltstones and mudstones are interpreted as paleosols (Brandlen 2008; Fiorillo et al. 2010a, 2010b; Flaig et al. 2013). Field characteristics, micromorphology, geochemistry and palynology have been described in detail for nine paleosol types (pedotypes) representing the range of pedogenic development in the Prince Creek Formation (Flaig et al. 2013). The paleosols are all relatively immature and are characterized by blocky structure, brown or gray color, redoximorphic features, abundant carbonaceous material, Fe-depletion coatings, and siderite, suggesting that the soils were periodically water saturated and anoxic. The presence of Fe-rich mottles, ferruginous and manganiferous nodules and void and grain coatings, burrows, and weakly developed illuvial clay coatings suggest oxidizing conditions and periodic drying of some of the soils (Brandlen 2008; Fiorillo et al. 2010a, 2010b; Flaig et al. 2013).

Paleosols for Clay Mineralogical Study.—

In this paper we focus on clay minerals from three paleosol profiles at North Kikak–Tegoseak (NKT), Kikiakrorak River Mouth (KRM), and Liscomb Bonebed (LBB), exposed in a generally south-to-north transect along the Colville River, Alaska (Fig. 1). Flaig et al. (2013) interpreted these three paleosols as similar to Aquic Inceptisols (Soil Survey Staff 2014). Detailed characteristics of these paleosols and their associated biota have been published previously (Flaig et al. 2013), and we briefly summarize their field and micromorphological characteristics in Figures 2 and 3.

Methods

Clay-Mineral Identification

Nineteen paleosol samples from NKT, KRM, and LBB (Figs. 2, 3) and two bentonite samples (06KKT-20.5, PFDV 17; Fig. 1) were selected for detailed clay mineralogical analyses. The < 2 µm clay fraction was separated by sedimentation using the pipette method (Burt 2004). Preferred-orientation clay-mineral slides were prepared with the filter transfer method (Moore and Reynolds 1997). The < 0.2 µm fraction was isolated from the total clay by centrifugation using the time and speed calculated with Centriset (Poppe and Eliason 2009). The clay mineralogy of coarse (< 2 µm) and selected fine (< 0.2 µm) clay-size fractions was determined by X-ray diffraction (XRD) using a Panalytical X'PERT PRO Materials Research Diffractometer (MRD), at the University of Alaska Fairbanks Advanced Instrumentation Laboratory (AIL). The instrument utilized a Cu k-α source (45 kV and 40 mA), a Ni filter, 0.02 radian Soller slits, a one-half-degree divergence slit, and an Xcellerator strip detector. Three scans, 2° to 60° 2θ, were performed for each sample using a 0.0125° step size and a count time of 20 s/step. Each sample was interpolated to a mean step value of 0.02°.

Seven XRD runs were performed for each < 2 µm sample; these include an air dried sub-sample, Ca2+-, Mg2+-, and K+-saturated subsamples, an ethylene-glycol (EG) solvated subsample (in vapor for 24 hours at room temperature), and subsamples heated at 300 °C and 550 °C for 4 hours. Random powder mounts of the air-dried < 2 µm clay fractions were prepared in 0.7 mm glass capillary tubes to determine illite polytypes and distinguish between dioctahedral and trioctahedral clay minerals. Two XRD runs were performed on selected < 0.2 µm samples following Mg2+-saturation and EG solvation.

Clay minerals were identified according to Moore and Reynolds (1997). Illite–smectite (I/S) mixed-layer clays were identified following Środoń (1984, 1999). The I/S interstratification analysis (random, partially ordered, and ordered) was determined qualitatively by comparison with calculated diffraction profiles and tables that show the relationship between clay composition and peak position (Reynolds 1980; Moore and Reynolds 1997). R  =  0 (random ordered) was taken from the “Reichweite” (R) notation (Reynolds 1980). Polytypes were identified by diagnostic peaks (Moore and Reynolds 1997) and quantified using the intensity ratio of peaks at 2.80 Å /2.58 Å (Maxwell and Hower 1967). Semiquantitative analyses of the diffraction data were carried out using mineral intensity factors (MIF) taken from the literature (Appendix 1; Biscaye 1965; Laves and Jähn 1972; Tributh 1991; Kahle et al. 2002) and also calculated using NEWMOD (Reynolds 1985). The amount of each clay mineral was calculated based on the MIF using the 100% approach (Moore and Reynolds 1997; Kahle et al. 2002). We assumed that smectite (montmorillonite), illite, kaolinite, chlorite, quartz and I/S mixed-layer clays constitute 100% of the mineralogy in the total clay fraction. The following basal (001) peaks in the Mg2+-EG samples were used for clay-mineral identification: 17 Å peak for smectite (montmorillonite), 10 Å peak for illite, 7.2 Å peak for kaolinite, 7.1 Å peak for chlorite, and 4.26 Å peak for quartz. The illite/smectite (I/S) 002/003 reflection was used to estimate the amount of illite in the mixed-layer clay.

Geochemistry

The weight percent of clay minerals, determined from semiquantitative analyses of the XRD data, were verified with geochemical data obtained using a Cameca SX-50 electron microprobe at the University of Alaska Fairbanks AIL. Abundances (in wt. %) of the light major oxides were measured from bulk samples using a PANalytical Axios wavelength-dispersive X-ray fluorescence spectrometer (WD-XRF) at the University of Alaska Fairbanks AIL. Major-oxide concentrations were used for geochemical mass-balance calculations (Brimhall and Dietrich 1987; Brimhall 1991a, 1991b) after bulk-density measurements were determined by the clod method (Blake and Hartge 1986). Geochemical mass-balance calculations are used to identify residual enrichment (Zr, Ti), volume changes, and element translocation that may reflect weathering trends or additions of material in soils and paleosols (Brimhall et al. 1991a, 1991b; Driese et al. 1992; McCarthy and Plint 2003).

Extractable Fe, Al, and Si

Unconsolidated paleosol samples were processed at the University of Alaska Matanuska Experimental Farm Soils Laboratory. Extraction procedures of acid-ammonium-oxalate and sodium-pyrophosphate (Parfitt and Childs 1988; McKeague 1978) were used to identify amorphous forms of Fe and Al. Acid ammonium oxalate extracts Al and Fe from allophane, imogolite, and Al–Fe–humus complexes (Alo and Feo) (McKeague 1978; Wada 1989), as well as Si from allophane and imogolite (Sio) and Fe from ferrihydrite (Parfitt and Childs 1988). The acid ammonium oxalate method is used to determine amorphous inorganic Fe and Al and organic-complexed Fe and Al in soils. Sodium pyrophosphate slightly extracts amorphous inorganic forms of Al and Fe (Alp and Fep) but effectively extracts Fe from ferrihydrite and goethite. Additionally, sodium pyrophosphate–extractable Al corresponds to Al in Al–humus complexes (Parfitt and Childs 1988; van Breeman and Buurman 2002). Alp/Alo and Fep/Feo ratios in Andisols indicate the relative proportions of active Al and Fe in organic complexes (Pinheiro et al. 2004). Alp/Alo ratios can be used to separate allophanic Andisols (ratio < 0.5) from nonallophanic Andisols (ratio > 0.5) (Saigusa et al. 1991, in Shoji et al. 1993). The acid ammonium oxalate extraction is also used to identify andic soil properties, which are diagnostic soil characteristics for Andisol classification (Soil Survey Staff 2014).

Results

Paleosol profiles NKT, KRM, and LBB all have similar clay-fraction X-ray diffraction patterns, indicating similarity in the clay mineralogy regardless of provenance. Variability between the profiles is restricted to relative abundances in the mineral assemblages. In all three paleosols, NKT (Fig. 4A), KRM (Fig. 4B), and LBB (Fig. 4C), the total clay fraction (< 2 µm) contains discrete smectite (S), discrete illite (I), kaolinite (Kaol), chlorite (Chl), and mixed-layer illite–smectite (I/S). Quartz is identified in the clay-size fraction by peaks at 4.26 Å and 3.34 Å (Figs. 4A, B, C).

Smectite Characterization

Discrete smectite (S) is indicated by a shift of the 13.74 Å and 12.08 Å peaks in the Ca2+- and Mg2+- saturated samples to ∼ 17 Å in the Ca2+- and Mg2+-EG-solvated samples, respectively (Fig. 4C). Montmorillonite is the only variety of smectite identified in all three paleosols. It was identified by a sharp peak at ∼ 17 Å (001) of Mg2+-EG-solvated samples (Fig. 5A). The 002 and 003 basal reflections of the smectite Mg2+- and Ca2+- EG-solvated samples are weak and occur in the region of the 001/002 and 002/003 basal reflections for I/S. Higher-order reflections (e.g., ∼ 2.81Å in Mg2+- and Ca2+-EG-solvated samples) are detectable, indicating the presence of discrete smectite (Fig. 5A, B). Heating K+-saturated samples to 300°C and 550°C results in collapse of the ∼ 17 Å peak to 10 Å, further suggesting the presence of smectite (Fig. 5C). The 006 reflection in randomly oriented < 2 µm powder mounts shows a montmorillonite peak at 1.49 Å in the 62.22° to 61.67° 2θ region (inset at top right of Fig. 6). Smectite variability between the three profiles is between 60 and 80 wt %, and KRM has ≥ 70% smectite in all horizons (Fig. 7A).

Microprobe analysis of the clay samples (Table 1) further suggests that montmorillonite is the type of smectite present in the clay-mineral assemblages when comparing the geochemical data with the Na-montmorillonite (SWy-2) standard (Table 2). The excess Fe detected in the microprobe analysis is probably due to amorphous Fe precipitates in pedogenic features such as iron oxide mottles, nodules, and coatings (Figs. 2, 3) (Flaig et al. 2013).

Illite Characterization

Discrete illite (I) is found in all samples and identified by peaks at 10 Å, 5 Å, and 3.3 Å. Discrete illite basal reflections do not change with any treatment, showing a periodicity in the 00L diffraction patterns. The discrete illite rational series of d(001) compared to the non-regular reflections of the mixed-layer I/S is useful to differentiate discrete illite from I/S mixed-layer clays (Środoń 1984) (Fig. 4). The 00L (001, 002, and 004) illite peaks do not shift with glycolation and the d ≈ 5 Å (Fig. 4) and d ≈ 1.99 Å (Fig. 5B) peaks are present in most samples and do not change position regardless of the treatments. Illite has a higher wt % in all horizons of the NKT profile (Fig. 7A).

In randomly oriented samples, the 060 reflections are used to identify illite polytypes (Fig. 6; Moore and Reynolds 1997). All samples show at least three of the 2M1 polytype diagnostic reflections at 3.88 Å, 3.49 Å, 3.20 Å, 2.98 Å, 2.86 Å, and 2.79 Å (Fig. 6). The 2M1 polytype reflections at 3.49 Å, 3.20 Å, 2.98 Å, 2.86 Å, and 2.79 Å are well defined (Fig. 6). The 1M polytype reflections are weak or absent except for a weak reflection at 3.07 Å that is present in samples LBB 20 and LBB 23, suggesting small amounts of the 1M polytype in that profile (Fig. 6). For example, we estimated the percentage of the 2M1 polytype in LBB 23 by applying the intensity ratio at 2.80 Å /2.58 Å (Maxwell and Hower 1967) where the peak area of the 2.80 Å is 0.3 counts, the peak area of the 2.58 Å is 1.62 counts, and the 2.80 Å/2.58 Å ratio is 0.19, which gives a percentage of the 2M1 illite polytype of 75%.

Illite–Smectite (I/S) Characterization

Interstratified illite–smectite consists of various combinations of illite layers and smectite layers stacked parallel to the c axis (Reynolds 1980). In the NKT, KRM, and LBB profiles, mixed-layer illite–smectite (I/S) is identified from EG-solvated samples by peaks near 8.5–8.9 Å and 5.5–5.6 Å (Figs. 4A, B, C). The wt % in the three profiles is very similar and < 20% (Fig. 7A). Smectite is the main component of the I/S mixed-layer clays (> 70%) (Fig. 7B). In the total clay diffraction patterns, two broad peaks (∼ 8.7Å and ∼ 5.6 Å) in the region near 10° 2θ and 16 to 17.7° 2θ show reflections of the I/S 001/002 and 002/003 peaks, respectively (Fig. 4). The significant shift in the peak of the EG-solvated samples that occurs as a shoulder on the 001 illite reflection is diagnostic of I/S (Fig. 5C). In the EG-solvated samples the ordering type (Reichweite) of the I/S is determined by the position of the reflection between 5° and 8.5° 2θ (Moore and Reynolds 1997). The presence of a peak near 8.5–8.7 Å in EG-solvated samples suggests that the I/S interstratification is random (R  =  0) and smectite-rich (Moore and Reynolds 1997; Parry et al. 2002). R signifies the most distant layer in an interstratified sequence that affects the probability of occurrence of the final layer (Reynolds 1980). R  =  0 means random or no neighboring dependence and, therefore, no preferred sequence in stacking of layers (Środoń 1999).

The percent of illite layers in the I/S can be estimated by the difference between the 001/002 and 002/003 reflections (Δ2θ) (Moore and Reynolds 1997). In all the samples the illite content of the I/S mixed-layer clay varies between 10 and 20% (Fig. 7B). In the < 0.2 µm fraction, I/S is the main mixed-layer clay (Fig. 8A).

Kaolinite Characterization

Kaolinite (Kaol) is identified by peaks at 7.2 Å and 3.58 Å that collapse upon heating to 550°C (Fig. 5C). Heating at 550°C collapses the kaolinite peaks and increases the intensity of the illite peak (10 Å) of the K+-saturated samples (Fig. 5C). The 001 and 002 kaolinite reflections are nearly superimposed upon the chlorite 00L series. Nevertheless, in the NKT and KRM profiles it is possible to differentiate kaolinite (3.58 Å) and chlorite (3.54 Å) as two sharp peaks next to each other at 25° 2θ (Fig. 5C). The presence of the 003 peaks of both chlorite and kaolinite allow the identification of each mineral since neither of these reflections interferes with the reflection of the other (Figs. 4, 5). Kaolinite is mainly present in the NKT profile (Fig. 7A).

Chlorite Characterization

Chlorite (Chl) is identified by peaks at 14.15 Å, 7.1 Å, 4.72 Å, and 3.54 Å. The 001 and 003 reflections of chlorite, d ≈ 14.15 Å and d ≈ 4.72 Å respectively, allow for differentiation of chlorite from kaolinite (Fig. 4). The 14.15 Å chlorite peak persists even after glycolation of Mg2+-saturated samples (Fig. 5C). After heating at 550°C, the 001 chlorite reflection increases in intensity and the 002, 003, and 004 reflections are weakened, indicating the presence of discrete chlorite (Fig. 5C) (Moore and Reynolds 1997). Chlorite is present mainly in the LBB profile (Figs. 4C, 7A).

Semiquantitative Clay-Mineral Analysis

Figure 7A illustrates the weight percent distribution for clay minerals in the NKT, LBB, and KRM paleosol profiles determined using the MIF 100% approach (Kahle et al. 2002) (Appendix 1). Smectite is the dominant clay mineral in all three profiles (60–81%), and almost the only constituent of the bentonites (> 95%) (Fig. 8B). Discrete illite (4–21%) and I/S mixed-layer clays (5–15%) are present in all paleosol samples, while kaolinite (0–8%) and chlorite (0–9%) are a minor component of the clay assemblages. I/S mixed-layer clays are between R0 illite (0.10)/smectite to R0 illite (0.25)/smectite (Fig. 7B). The discontinuities marked by the Ti/Zr ratio (Figs. 2, 3) coincide with changes in the weight-percent distribution for clay minerals within the NKT, LBB, and KRM paleosol profiles (Fig. 7A).

Clay-Mineral Geochemistry

Using the standard composition of kaolinite (KGa-1), Na-montmorillonite (SWy-2), illite (IMt-1), and chlorite (CCa-2) (Van Olphen and Fripiat 1980; Table 2), and weight-percent values obtained from XRD data (Fig. 7A), we calculated the chemical composition of the paleosol clay-mineral samples (Table 3). The calculated chemical compositions are compared with microprobe analyses of the clays as a cross-check on the overall analysis (Fig. 9A; Tables 1, 3). There is a good correlation between the two methods, particularly for the more abundant components (SiO2, Al2O3, and Fe2O3; R2  =  0.9884). The correlation is not as good for the less abundant components (MgO, K2O, CaO, Na2O, and TiO2; R2  =  0.36), which might be due to analytical errors or variability in the composition of soil minerals (particularly smectite, which can be highly variable in chemical composition).

Geochemical Composition of Bentonites

The monomineralic character (pure montmorillonite) of the 06KKT-20.5 and PFDV bentonite samples is shown in Figure 8B. The series of basal reflections, with no splitting associated with glycolation, indicates well-ordered stacking sequences with no interstratification (e.g., Olsson and Karnland 2011) (Fig. 8B). Furthermore, a plot of standard montmorillonite against the geochemical data from the bentonites (clay fraction) and bulk-sample geochemistry shows a very good 1∶1 correlation (Fig. 9B, Table 4). These results indicate that the chemical composition of these bentonites is similar to the chemical composition of a standard montmorillonite. The two bentonite samples (06KKT-20.5 and PFDV-17, Fig. 1) plot in the rhyolitic to rhyodacitic compositional fields on a Zr/TiO2-Nb/Y classification diagram (Fig. 10; Winchester and Floyd 1977).

In order to test the hypothesis that clay minerals in the Prince Creek paleosols are derived, at least partly, from bentonite parent materials with a chemical composition similar to 06KKT-20.5 and PFDV 17, pure smectite and pure illite end members were plotted using the Al/K–Si/K diagram (Fig. 11A; De Caritat et al. 1994). We used the chemical composition from a measured bentonite at the Kikak–Tegoseak dinosaur quarry (near the 06KKT-20.5 profile) for the smectite end member (Fiorillo et al. 2010a, 2010b; Flaig et al. 2011) (Table 1). We used an ideal illite stoichiometry (IMt-1) for the illite end member (Van Olphen and Fripiat 1980; Table 2). When clay-mineral chemical compositions of NKT, KRM, and LBB profiles obtained from the microprobe are plotted on the diagram, there is a strong linear compositional relationship (R2  =  0.98) between illite, bentonite, and the clay samples (Fig. 11A).

Geochemical Mass-Balance Calculations

Figures 2, 3A, and 3B show lithologic discontinuities in the three paleosol profiles, based on depth distributions of Ti/Zr ratios. In uniform parent materials, ratios of Ti/Zr should change gradually and uniformly with depth, without sharp inflections or reversals in trends (Birkeland 1999). In the three paleosols, depth plots of calculated Ti/Zr ratios indicate large inflections that coincide with parent-material discontinuities identified on the basis of grain size and micromorphological features. For example, the NKT profile shows a change in the trend of the Ti/Zr ratio at NKT 42 (Fig. 2). In the LBB profile there is a change in this ratio at LBB 20 (Fig. 3B) and in KRM there is a gradual decrease in the Ti/Zr ratio with depth and a small inflection at KRM 23 (Fig. 3A).

To interpret weathering trends in the paleosols, the C horizons were used as an approximation of the parent material of the soil profiles. The mass-transport calculations show gains (positive values) or losses (negative values) of each major element in the paleosols compared to the parent materials (Figs. 2, 3A, B). The NKT profile has Fe depletion, and minor depletion of Al and Si, in the NKT 46, NKT 44, and NKT 42 samples (Bg and Bw horizons), and significant accumulation in the NKT 40 sample (Bt horizon). Phosphorus shows a net loss, except in the NKT 40 sample, where it shows a gain of 30% (Fig. 2). In the KRM profile, all samples (except the uppermost one, KRM 25) show Si and Al depletions (maximum loss of 36%), and phosphorus mass-balance values between 0.2 and 2 (20–200%), indicating pronounced phosphorous accumulation.

Geochemical values in the LBB profile are highly variable (Fig. 3B). Given that this is a composite paleosol, the variability is probably due mainly to lithological discontinuities derived from different depositional episodes rather than from pedogenic accumulation or loss of elements (Flaig et al. 2013).

Extractable Iron and Aluminum

The acid ammonium oxalate and sodium pyrophosphate extractable Fe and Al values (Table 5) show that active Fe and Al are significant components of the NKT and KRM paleosol profiles. The NKT 42 and KRM 25 samples have 3.8% and 3.4%, respectively, of Alo+1/2 Feo (Table 5), which meets the requirement for andic soil properties (Soil Survey Staff 2014). The data in Table 5 show that NKT 42 and KRM 25 also have Alp/Alo ratios > 0.5.

Discussion

The Late Cretaceous Arctic contrasts markedly with modern polar regions in having both higher temperatures and an intensified hydrological cycle (Suarez et al. 2013), although the polar light regime was similar to the present (Spicer and Herman 2010). Flaig et al. (2013) suggested that, in the absence of ground ice, paleosols formed at high latitudes under greenhouse conditions would be unlikely to have unique pedogenic signatures but that increased temperatures and moisture availability might result in increased biological activity and mineral weathering leading to more rapid and mature development of soils. Evidence for pedogenic illite–smectite formation in the Prince Creek Formation stands in stark contrast to much more modest clay mineral transformations typical of modern arctic soil-forming environments (Ping et al. 1999; Borden et al. 2010). Instead, the Late Cretaceous Arctic of northern Alaska was one in which rapid weathering of volcanic ash and alluvium in a seasonally wet environment resulted in a rich ecosystem with soils that supported a diverse flora and a year-round dinosaur fauna (Fiorillo et al. 2010a, 2010b; Flaig et al. 2014). The Late Cretaceous Prince Creek Formation, therefore, provides an exquisite example of a flourishing high-latitude ecosystem in a greenhouse world.

Detrital Clays

The smectite (montmorillonite), discrete illite, chlorite, and kaolinite are interpreted to be a detrital clay suite inherited from the parent material (floodplain alluvium). The low concentrations of soluble silica and base cations (Ca2+, Mg2+, Na+) (Figs. 2, 3), high concentrations of K+ from the micas present in bulk samples (Figs. 2, 3), and the organic-rich nature of the paleosols are suggestive of an acidic groundwater regime, likely with a pH < 4.5 (Flaig et al. 2013). Such conditions are incompatible with the formation of pedogenic montmorillonite (e.g., Essington 2004; Borchardt 1989). Soil-formed smectites are typically Fe-rich beidellites, which form at the expense of montmorillonite and are the most stable smectites in modern pedogenic environments (Šucha et al. 2001; Środoń 1999). Since smectite in the Prince Creek paleosols is montmorillonite (inset at top right of Fig. 6) we interpret it as bentonitic in origin rather than pedogenic. Furthermore, the Al/K–Si/K diagram (Fig. 11A; De Caritat et al. 1994) shows a compositional relationship between illite, bentonite, and clay samples from Prince Creek paleosols. All of the clay samples plot between the two endmembers (R2  =  0.98), suggesting that their composition is the result of a mixture of detrital illite with bentonite-derived smectite. Smectite-rich bentonites are considered the potential source for the epiclastic smectite (reworked and redeposited product of weathered volcanic ash) in the NKT, KRM, and LBB profiles. This is consistent with the fact that smectites are typical in sediments and sedimentary rocks after periods of above-average volcanism (Borchardt 1989), such as that which occurred in northern Alaska during the Late Cretaceous. It is likely that most, if not all, of these Late Cretaceous bentonites were sourced to the west in northeastern Russia (Bergman et al. 2006; Kelley et al. 1999).

Burial diagenesis is unlikely to have had a major impact on the clays in the Prince Creek Formation because maximum burial temperatures, based on vitrinite-reflectance values, never exceeded 48° C (Robinson 1989; Johnson and Howell 1996). This is consistent with the observation that illite, illite polytypes, and I/S mixed-layer clays identified in this study do not show evidence of progressive illitization associated with burial diagenesis. During diagenesis, smectite can alter to illite through a transformation sequence of smectite to I/S of different compositions and finally to illite (Cuadros and Altaner 1998; Meunier 2005). However, the most abundant illite polytype in our samples, the 2M1 illite polytype (Fig. 6), corresponds with the beginning of the lowest-temperature metamorphic zone during burial diagenesis (Moore and Reynolds 1997). Consequently, we interpret discrete illite (2M1 polytype) as detrital in origin and likely sourced from very low- to low-grade metamorphic rocks from the Brooks Range to the south. We also consider detrital illite as the potential source of K+ for the pedogenic smectite illitization process (discussed below). In bulk samples, the K2O plot of XRF data vs. microprobe data of the clay fraction shows a good correlation (Fig. 11B), suggesting that illite is the major source of K+.

The occurrence of chlorite and kaolinite in the three profiles is minimal, and their abundances do not exhibit any trends with depth in the profiles (Fig. 7A), suggesting that they have a detrital origin. Primary chlorite in soils is typically of detrital origin (Barnhisel and Bertsch 1989). The chlorites found in our samples exhibit the characteristic 00L reflections (7 Å) of trioctahedral chlorites, which are considered inherited from the parent material, as opposed to pedogenic hydroxy-interlayered vermiculite (HIV) or pedogenic chlorite (Allen and Hajek 1989). The lack of intermediate species such as chlorite–vermiculite, vermiculite, or chlorite–smectite in our samples, and the overall weak paleosol development, suggests that a detrital origin for the chlorites is the most reasonable interpretation.

Kaolinite formation in situ is also not common in weakly developed soils (Wilson 1999). In the three Prince Creek paleosols kaolinite is interpreted as detrital rather than authigenic since kaolinite formation under surface weathering conditions is more typical of highly developed soils in the modern tropics (e.g., Aristizabal et al. 2005). Generally, it occurs in freely drained, acid- and base-depleted tropical environments, where an abundant supply of water ensures the availability of the required silica and alumina cations (Wilson 1999).

Chlorite and kaolinite are common clay minerals in recent sediments from the Beaufort Sea and Arctic Ocean, where they are derived from weathering of sedimentary and metasedimentary rocks from the Brooks Range and the North Slope of Alaska (Naidu et al. 1971). We suggest that the detrital chlorite and kaolinite in our paleosols likely had a similar provenance.

Pedogenic I/S Mixed-Layer Clays

I/S is the only mixed-layer clay identified in our paleosol samples (Fig. 4). Illite–smectite comprises 5–15% of the total clay fraction in all samples (Fig. 7A). The random stacking order of I/S found in the NKT, KRM, and LBB samples is consistent with a supergene, rather than diagenetic, illitization (e.g., Huggett and Cuadros 2005). Furthermore I/S is the main clay mineral in the fine clay fraction (< 0.2 µm) (Fig. 8A) where it is interpreted to represent pedogenic clay, since no quartz (a common mineral in the detrital fraction) is present (Vitali et al. 2002). Although illitization of smectite commonly occurs during burial diagenesis, it is also known to occur in soil environments (Huggett and Cuadros 2005). Huggett and Cuadros (2005) suggest that the transition from smectite to illite at low temperature occurs through Fe reduction in octahedral sites leading to increased layer charge, coupled with K+ fixation by smectite. The Fe reduction is interpreted to be the result of wetting (reducing) and drying (oxidizing) cycles in gley soils, in which reoxidation of reduced Fe is never complete (Huggett and Cuadros 2005). Similar drainage conditions are present in gleyed horizons (Bg) of the NKT and KRM paleosols (Figs. 2, 3), which suggests that the illitization of smectite could result from multiple wetting (reducing) and drying (oxidizing) cycles. Wetting and drying conditions are interpreted from alluvial-coastal-plain subenvironments of the Prince Creek Formation based on paleosol characteristics and vegetation assemblages (Flaig et al. 2013). The wetting and drying is probably the result of both fluctuations in local water-table conditions and more regional flooding due to periods of intense seasonal runoff probably related to melting snow and ice in the ancestral Brooks Range (Flaig et al. 2013; Fiorillo et al. 2010a; Spicer 2003). In the NKT (NKT42, NKT 44) and KRM (KRM 21, KRM 22, KRM 23), profiles the illite percentage in the I/S increases in the Bg horizons suggesting a relationship between illite formation and gley horizons (Figs. 2, 3, 7). Therefore, we conclude that for generating I/S mixed-layer clays in the Prince Creek Formation, the probable mechanism is low-temperature pedogenic illitization of primary smectite. Given the low burial temperatures for the Prince Creek Formation, a pedogenic origin for the I/S mixed-layer clays seems much more reasonable than a diagenetic origin (Huggett and Cuadros 2005).

Other Pedogenic Processes

The geochemical trends in the NKT and KRM profiles show characteristics that can be interpreted as the result of redox processes and andic soil properties (Figs. 2, 3). For example, the depletion of Fe and Al in NKT 46 and NKT 44 and the gain of Fe and Al in NKT 40 can be attributed to a possible lithologic discontinuity and/or accumulation in the Bw horizon of Fe and Al complexes translocated from the upper horizons. Some trends in the KRM profile are interpreted as phosphorus accumulation and Al and Si depletion (Figs. 2, 3). In addition, values from acid-ammonium-oxalate extractable Fe and Al (Feo, Alo) and pyrophosphate-extractable Fe and Al (Fep, Alp) (Table 5) indicate that hydrous oxides of Fe and Al are significant components of the NKT and KRM paleosols and form complexes with humus.

In order to be recognized as having andic properties when no volcanic glass is present, as is the case for the Prince Creek paleosols, a soil is required to contain less than 25 weight percent organic carbon and meet the following additional criteria: Al + 1/2 Fe percentages (by ammonium oxalate) totaling 2% or more; a bulk density of 0.9 g/cm3 or less; and phosphate retention of 85% or more (Soil Survey Staff 2014). Diagenesis can be problematic in ancient soils, and caution is required when attempting to make interpretations based on modern soil criteria since some properties may be modified or incompletely retained in paleosols due to burial compaction and/or increased temperatures (Retallack 2001). However, given the relatively low burial temperatures of Prince Creek paleosols (48 °C max.), their well-preserved micromorphological features, and the fact that the paleosols remain unlithified, we are fairly confident that diagenetic modification of values of ammonium oxalate and sodium pyrophosphate Fe and Al is minimal. Nevertheless, we cannot rule out the possibility that some of the original extractable Fe and Al has been lost from these paleosols following burial.

The NKT 42 horizon has 3.8% of Alo + 1/2 Feo; however, the bulk density and phosphate retention (based on wt. % of P2O5) (Fig. 2), based on mass-balance calculations, do not meet the requirements for andic soil properties. This may be, at least partially, because the mass-balance geochemical calculations are based on the assumption that the C horizon approximates the original alluvial parent material. Ammonium-oxalate-extractable Feo and Alo is > 2%, indicating that even though phosphate retention and bulk density do not meet criteria for andic soil properties, Fe- and Al-organic complexes were still translocated in the profile but that the soil was insufficiently well developed for Fe and Al complexation to reach the levels necessary for identification of andic properties.

The KRM profile exhibits trends of depletion of total Al and Si and significant phosphorus gain (60–200%) (Fig. 3A). Values of phosphorus accumulation, Al and Si depletion, and low bulk density are interpreted as evidence of andic trends (Bestland 2002). The KRM 25 horizon has 3.4% of Alo + 1/2 Feo, low bulk density, and phosphorus gain > 100%, all of which we interpret as evidence for andic soil properties. The organic-rich nature of the paleosol, and high Alp/Alo, suggests a low pH and formation of Fe–Al–humus complexes rather than amorphous silicates, consistent with a non-allophanic trend in modern Andisols (Shoji et al. 1993). Mn accumulation (Fig. 3) is consistent with the presence of manganese nodules (Flaig et al. 2013). The high amount of Fe3+ interpreted from the acid ammonium oxalate extraction is consistent with the high Fe2O3 values obtained by microprobe analysis (Table 1) and the presence of abundant ferruginous features (Table 4). The presence of amorphous Fe3+ precipitates suggests a balance in the reducing and oxidizing tendencies that is diagnostic for redoximorphic soils (Essington 2004).

Andic Properties and Paleo-Andisols

Andisols are typically volcanic-ash soils with properties inherited from or associated with properties of tephra, but they can also be formed from non-tephric materials such as sedimentary rocks and mixed materials of tephra and loess (Shoji et al. 1993). Andisols include amorphous materials commonly formed during weathering of tephra and andic materials, and they typically have a significant content of volcanic glass, low bulk density, and high organic matter content (Birkeland 1999; Soil Survey Staff 2014).

Key features, such as volcanic glass and andic soil properties, typically do not persist in paleosols (Bestland 2002). However, we demonstrate that the KRM profile shows evidence of andic soil properties within 20 cm of a Bg horizon, which partially satisfies the criteria to be classified as a paleo-Andisol. According to Nettleton et al. (2000), paleosols can be classified as paleo-Andisols if they have andic soil properties in subhorizons that total 30 cm or more within 50 cm of the paleosol surface. However, the NKT, KRM, and LBB paleosols are truncated soils where the diagnostic A horizon is absent, presenting another difficulty to classification of the soil type. Alternatively, Andept-like paleosols can be identified based on geochemical trends of phosphorus accumulation, potassium retention, and alumina and iron depletion, assumed to represent traces of andic soil-forming processes (e.g., Bestland 2002). The KRM profile also shows trends of phosphorus accumulation and potassium retention, suggesting an Andept-like alluvial paleosol. Amorphous forms of silica and aluminum, like volcanic glass and allophane, change with time into minerals like clays (Mack et al. 1993). The presence of allophane cannot be demonstrated in our paleosols. However, modern non-allophanic Andisols are dominated by 2∶1 clays (Shoji et al. 1993), and smectite (montmorillonite) is the main clay mineral (60–81%) in all three of our paleosol profiles. One profile (KRM) displays clear evidence of andic soil properties, but the other two (NKT, LBB) do not meet modern soil criteria for andic properties (Soil Survey Staff 2014). Nevertheless, the NKT and LBB profiles still display evidence of andic trends, which suggests either that they were too immature to fully develop andic properties or that some of the necessary geochemical properties for classification as Andiosols were lost during burial despite the low-temperature diagenetic conditions. Consequently, it is most appropriate to refer to these Prince Creek paleosols as Andept-like paleosols.

Paleosol Formation in the Late Cretaceous Arctic

The alluvial parent material of the NKT, KRM, and LBB paleosols is dominated by smectite (> 60%) derived from bentonites. On the Prince Creek coastal plain, the bentonite-derived smectite was incorporated as detrital material (epiclasts) that was subsequently deposited as alluvium on floodplains.

The observation that the parent material is mostly of volcanic epiclastic origin, bulk densities are low, phosphorus and organic-matter contents are high, and Fe–Al–humus complexes are present suggests that the genesis of these paleosols was strongly influenced by andic soil-forming processes (andosolization). The properties of the bentonite-derived smectite parent material, incorporated as detrital material, and the pedogenic environment (organic-rich paleosols in an alternating wet and dry environment) probably favored an anti-allophanic process that inhibited the formation of allophane and imogolite (Shoji et al. 1993). Non-allophanic Andisols are characterized by acidic pH values (Shoji et al. 1993), which were probably prevalent in the Prince Creek Formation during pedogenesis. Chromas of 2 and morphologic features in the NKT, KRM, and LBB profiles (Figs. 2, 3) indicate poorly drained paleosols that were reduced for significant periods of time. Under poor drainage, organic matter accumulates and its microbial decomposition is retarded by seasonal reducing conditions, formation of Al–humus complexes, and humification (Shoji et al. 1993). The weathering of volcanic glass under organic-matter-rich conditions with a pH < 4.9 results in organic acids being available to form metal–humus complexes. Humus competes for Al, reducing the Al available for co-precipitation with silica to form aluminosilicate minerals (Shoji et al. 1993). However, periodic desiccation of allophane may cause crystallization to phyllosilicate minerals, including smectite, which are also known to occur in modern Andisols either due to this process or as an inherited component (detrital) or as the result of additions from eolian dust (Jongmans et al. 1993). In summary, these trends lead us to interpret these ancient soils as non-allophanic Andept-like paleosols (Soil Survey Staff 2014).

Studies of volcanic ash soils in various parts of the world illustrate the effect of climate on the soil colloidal fraction (Zehetner et al. 2003; Ugolini and Dahlgren 2002). Soils characterized by non-allophanic clay mineralogy, an abundance of 2∶1 layer phyllosilicates, abundant formation of Al– and Fe–humus complexes, and a marked accumulation of humus are known to be developed from non-colored volcanic glass under mean-annual-precipitation values > 1000 mm (Shoji et al. 1993). The final transformation of smectite to illite in the mixed-layer I/S clays occurred in situ in the paleosol profiles. Redox processes and illitization of smectite are also known to occur in other soils and paleosols subjected to alternating phases of wetting and drying (Huggett and Cuadros 2005; Rosenau et al. 2013). The paleosols formed on floodplains of the Prince Creek Formation reveal features of wet–dry cycles (e.g., Fiorillo et al. 2009; Fiorillo et al. 2010b; Flaig et al. 2010; Flaig et al. 2011), perhaps as a result of seasonal flooding due to variations in temperature and precipitation related to the high paleo-latitude of Alaska (82–85° N) in the Late Cretaceous (Flaig et al. 2011), or as a result of more frequent fluctuations in water-table levels. Mn nodules and the presence of amorphous Fe3+ precipitates are indicators of redoximorphic features in hydric soils (Vepraskas and Lindbo 2012), and are present in all of the paleosols (Figs. 2, 3). In waterlogged soils, the reduced iron and gray colors in gley horizons are due in part to bacterial decomposition of organic matter at temperatures above freezing, and pH values close to 5. These pedogenic processes are consistent with current cool temperate paleoclimatic reconstructions of Late Cretaceous North Slope ecosystems (Spicer and Herman 2010; Flaig et al. 2011, 2013, 2014).

The presence of pedogenic I/S in paleo-Arctic paleosols of the Prince Creek Formation contrasts strongly with clay-mineral formation regimes in modern arctic soils. In these regions, permafrost and cryoturbation processes dominate in soils and chemical weathering is a relatively slow process such that clay mineralogy is quite similar to that of the parent material. For example, clay minerals in arctic tundra Gelisols from northern Alaska are dominated by detrital illite and kaolinite with low amounts of pedogenic vermiculite derived from detrital illite (Borden et al. 2010). This regime of weathering and clay-mineral formation is strikingly different from that observed in the Late Cretaceous Prince Creek Formation, where higher temperatures, abundant moisture, and the presence of volcanic-ash-derived smectite resulted in formation of pedogenic I/S and at least some paleosols with andic soil properties which may at least be partly attributable to spring flooding in the ancestral Brooks Range and North Slope coastal plain. Furthermore, these pedogenic processes are consistent with evidence of seasonality (spring runoff in response to warmer temperatures) recognized from other aspects of the Prince Creek Formation. For example, root structures are often present in channel deposits, which suggest some form of wet–dry seasonality (Fiorillo et al. 2010b; Flaig et al. 2011, 2014). Similarly, the taphonomic pattern of the distribution of several bone beds in the Prince Creek Formation, as mortality events, has been attributed to a seasonal pattern of Late Cretaceous spring runoff from the ancestral Brooks Range (Fiorillo et al. 2010b). Lastly, details of the histologic structure of dinosaur bone shows that dinosaurs grew to adulthood in a seasonal environment that affected growth rates, with slower growth during times of reduced food resources (winter) and faster growth during times of resource abundance (summer; Chinsamy et al. 2012). Though the degree of seasonality recorded across these lines of evidence may differ, it is striking nonetheless that such disparate data sets repeatedly suggest seasonality in the high-latitude Latest Cretaceous Prince Creek ecosystem, an ecosystem that likely experienced seasonal changes in moisture availability on a scale similar to that observed in the modern Arctic.

Conclusions

We interpret the origin of clay minerals and the pedogenic processes that occurred in paleosols from the Prince Creek Formation during the Late Cretaceous to result mainly from a seasonal moisture regime related to annual changing conditions of light and temperature in the Late Cretaceous paleo-Arctic. In the NKT, KRM, and LBB profiles the clay minerals are mainly detrital and only I/S is the product of surface weathering. Pedogenesis is heavily influenced by the volcanic-ash-charged parent material, and alternating wetting and drying resulting from a highly seasonal moisture regime.

Arguments that favor a pedogenic origin for I/S in the Prince Creek Formation rather than a diagenetic one include the following:

  • 1).

    Vitrinite-reflectance data from the Prince Creek Formation indicates a maximum burial temperature of 48°C, which is not sufficient to trigger the mineral reactions required to transform smectite into illite during diagenesis (Pollastro 1993; Meunier 2005). For example, the rate of reaction in smectite is a function of geothermal gradient, with most changes occurring at 100°C and a depth of about 1.5 km (Velde et al. 1986, in Borchardt et al. 1989). These values are higher than the burial temperature estimated for the Prince Creek Formation (Robinson 1989; Johnson and Howell 1996).

  • 2).

    In the first stages of diagenetic transformation, illite forms randomly ordered mixed-layer minerals with smectite. As the reaction progresses, the stacking sequence becomes ordered and the illite content increases (Meunier 2005). Illite in Prince Creek I/S mixed-layer clays is randomly ordered (R  =  0) and its content varies between 10 and 25% in all of the NKT, KRM, and LBB samples. There is no progression in the stacking order or in the amount of the illite.

  • 3).

    Bentonite layers do not contain any I/S mixed-layer clays. As a consequence, the process of illitization in the Prince Creek Formation appears to be exclusively of a low-temperature pedogenic origin.

In summary, smectite is the most abundant clay mineral in the three profiles (60–81%) followed by illite (4–21%), I/S (5–15%), chlorite (0–9%), and kaolinite (0–5%) (Fig. 7A). Discrete illite, smectite, chlorite, and kaolinite are interpreted as detrital in origin. In contrast, I/S is interpreted as pedogenic in origin. Detrital illite is the most likely source of K+ for the transformation of smectite to I/S mixed-layer clay in the pedogenic environment. For the NKT, KRM, and LBB paleosol profiles, the parent material is interpreted as a mixture of detrital smectite (from a bentonitic source) and detrital illite from older Cretaceous rock units in the Brooks Range (Naidu et al. 1971), probably also mixed with reworked alluvial sediments. Reworked material derived from sedimentary and metasedimentary rocks from the Brooks Range and the North Slope of Alaska is a likely explanation for the presence of kaolinite and chlorite (Naidu et al. 1971). The bentonites (06KKT-20.5 and PFDV 17) are monomineralic rocks composed of smectite, specifically montmorillonite. The fine clay fraction (< 0.2 µm) is considered to be purely pedogenic because I/S is the main component and no detrital material, such as quartz, is present (Fig. 8A).

Geochemical mass-balance, bulk density, and acid-ammonium-oxalate and pyrophosphate-extractable Fe and Al indicate that: 1) hydrous oxides of Fe and Al are significant components of the NKT and KRM paleosols and are forming complexes with humus, 2) KRM has a low bulk density and phosphorus gain > 100%, which we interpret as evidence for andic soil properties, and 3) the NKT paleosol meets only one of the three criteria for andic soil properties (extractable Fe and Al > 2%), indicating that pedogenic processes were similar to KRM but NKT was probably more weakly developed or, alternatively, geochemical indicators were not fully preserved during burial.

Low pH, phosphorus accumulation, Fe– and Al–humus complexes, high organic matter, and an abundance of 2∶1 clay minerals indicate that the genesis of the paleosols was strongly influenced by the properties of the parent material and that the clay minerals from epiclastic sources contributed to the development of non-allophanic andic properties.

Taken together, our results indicate formation of pedogenic I/S under alternating wet and dry conditions in the presence of high amounts of organic matter that also resulted in migration of Fe–Al–organic complexes that typically form in andic soils. Wetting and drying may have resulted from seasonal flooding due to snow melt in the ancestral Brooks Range or from more frequent water-table fluctuations on the low-lying coastal plain. Consequently, we conclude that the presence of volcanic ash-derived bentonites, along with warmer and wetter paleo-Arctic conditions, led to the formation of andic soil properties and pedogenic I/S on this Late Cretaceous high-latitude coastal plain.

Financial support for this research was provided by the National Science Foundation (OPP-425636 to McCarthy and OPP-424594 to Fiorillo), the Alaska Geological Society (AGS), the International Association of Sedimentologists (IAS), the University of Alaska Museum of the North, and the Advanced Instrumentation Laboratory (AIL) at the University of Alaska Fairbanks. The authors thank Colciencias Colombia for economic support and Melissa Dick at University of Alaska Matanuska Experiment farm soils laboratory for the soil analyses. The authors also thank Peter Flaig for providing all the paleosol background pedogenic information, and Rainer Newberry, Andrés Ochoa, and Ken Severin, for their assistance with data analyses and interpretations. We are sincerely grateful for the detailed and insightful reviews provided by Erik Gulbranson and Nick Rosenau, and Associate Editor Greg Ludvigson, whose comments helped us to improve clarity and focus.