Oxygen and hydrogen isotope compositions of paleosol phyllosilicates: Differential burial histories and determination of Middle–Late Pennsylvanian low-latitude terrestrial paleotemperatures

Palaeogeography, Palaeoclimatology, Palaeoecology 392 (2013) 382–397

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Oxygen and hydrogen isotope compositions of paleosol phyllosilicates: Differential burial histories and determination of Middle–Late Pennsylvanian low-latitude terrestrial paleotemperatures Nicholas A. Rosenau ⁎, Neil J. Tabor 1 Roy M. Huffington Department of Earth Sciences, Southern Methodist University, 3225 Daniel Avenue, Dallas, TX 75275-0395, USA

a r t i c l e

i n f o

Article history: Received 20 April 2013 Received in revised form 22 August 2013 Accepted 19 September 2013 Available online 26 September 2013 Keywords: Illinois basin Carboniferous Paleosol Stable isotope Clay mineral

a b s t r a c t The clay mineralogy, chemistry, and stable hydrogen and oxygen-isotope compositions were measured from 20 phyllosilicate samples representing 11 Pennsylvanian-age paleosol profiles taken from three cores in the Illinois basin in order to assess their utility as proxies of low-latitude terrestrial paleotemperatures. The majority of the samples are mineralogical mixtures of illite–smectite (I/S), kaolinite, and rarely discrete illite. Samples from a shallowly-buried locality in the northern part of the basin are dominantly composed of smectite-rich I/S, with variable amounts of kaolinite, and no discrete illite. Samples from deeply-buried, interior parts of the basin are composed of illite-rich I/S, variable amounts of kaolinite, and discrete illite. These phyllosilicate mixtures have δ18OV-SMOW and δDV-SMOW values that range from 17.2‰ to 23.0‰ and −56‰ to −27‰, respectively. Assuming that the phyllosilicates preserve a record of isotopic equilibrium with Pennsylvanian meteoric waters, these oxygen and hydrogen isotope values correspond to crystallization temperatures ranging from 22 ± 3 °C to 55 ± 3 °C. The clay mineralogy, phyllosilicate δ18O and δD values and calculated crystallization temperatures of 44 °C to 55 °C from deeply buried localities in the interior of the basin are not consistent with a pedogenic origin. Instead, these trends are considered to be the result of diagenetic recrystallization of pedogenic minerals in response to greater depths of burial (by ~1.5 to 3 km) in the southerly, basin-center localities, as well as an interval of middle Permian elevated heat flow associated with magmatic intrusions in the southern part of the basin. Phyllosilicate mineralogy, δ18O and δD values, and calculated phyllosilicate crystallization temperatures from a shallowly buried, northern locality in the Illinois basin are consistent with a pedogenic origin, and reveal a longterm warming trend from an average temperature of 23 ± 3 °C in the lower Desmoinesian to an average temperature of 32 ± 3 °C in the Missourian. This temperature change is coincident with a significant change in the composition of wetland vegetation in Euramerica, which has been attributed to a shift in low-latitude Pennsylvanian climate towards warmer and drier conditions in the Late Pennsylvanian. This study reveals the presence of a dynamic Late Paleozoic paleoequatorial icehouse climate characterized by significant low-latitude temperature variability unprecedented on modern Earth. © 2013 Elsevier B.V. All rights reserved.

1. Introduction During the Late Mississippian to Late Permian the earth was entrained in the Late Paleozoic Ice Age (LPIA; Isbell et al., 2012). Recent advances in our understanding of the LPIA reveal that it consisted of multiple glaciations that waxed and waned across Gondwanaland (Frakes et al., 1992; Isbell et al., 2003; Rygel et al., 2008; Isbell et al., 2012). Global compilations of climate-sensitive sediments, mineralogic assemblages, fossil floras and geochemical proxy records have provided insight into the paleoclimate and paleoenvironments during the LPIA (Nairn and Smithwick, 1976; Phillips and Peppers, 1984; Phillips et al., 1985a,b; ⁎ Corresponding author. Tel.: +1 214 768 3313. E-mail addresses: [email protected] (N.A. Rosenau), [email protected] (N.J. Tabor). 1 Tel.: +1 214 768 3313. 0031-0182/$ – see front matter © 2013 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.palaeo.2013.09.020

Rowley et al., 1985; Ziegler et al., 1987; Parrish and Peterson, 1988; Scotese et al., 1999; Tabor and Montañez, 2005; Came et al., 2007; Tabor and Poulsen, 2008; Tabor et al., 2008; Falcon-Lang and DiMichele, 2010). However, the far-field (tropical to sub-tropical) paleoclimatic impacts of the glaciations remain controversial (Ziegler et al., 1987; Cecil, 1990; Heckel, 1994; West et al., 1997; Isbell et al., 2008; Soreghan et al., 2008, 2009; Bishop et al., 2010). Particularly, questions remain regarding the extent to which paleotropical temperatures varied throughout the waxing and waning of Gondwanan ice sheets (Angiolini et al., 2007; Poulsen et al., 2007; Soreghan et al., 2007a,b; Tabor and Poulsen, 2008). Numerous studies have demonstrated that the isotopic composition of hydroxyl-bearing phyllosilicate minerals that form in lowtemperature sub-aerial weathering environments may provide information related to weathering and environmental conditions at

N.A. Rosenau, N.J. Tabor / Palaeogeography, Palaeoclimatology, Palaeoecology 392 (2013) 382–397

the time of formation, such as the oxygen isotope composition of local meteoric water and crystallization temperature (Savin and Epstein, 1970; Lawrence and Taylor, 1971, 1972; Bird and Chivas, 1988, 1989; Lawrence and Rashkes-Meaux, 1993; Delgado and Reyes, 1996; Stern et al., 1997; Tabor et al., 2002, 2004; Vitali et al., 2002; Tabor and Montañez, 2005). Consequently, authigenic, soil-formed (pedogenic) phyllosilicate minerals taken from paleosol profiles have the potential to provide proxies of terrestrial paleoenvironmental conditions, provided that the original isotopic signatures have not been diagenetically-altered. Tabor and Montañez (2005) estimated crystallization temperatures of mixtures of kaolinite and 2:1 phyllosilicate minerals in Upper Pennsylvanian (Virgilian)–Lower Permian (Leonardian) paleosols in north-central Texas, U.S.A. To date, this study represents one of the few quantitative terrestrial paleotemperature estimates for the Pennsylvanian tropics (see also Lawrence and Rashkes-Meaux, 1993; Tabor, 2007). This contribution expands upon this limited data set by providing paleotropical Middle–Late Pennsylvanian terrestrial paleotemperature estimates from the Illinois basin (U.S.A.). This paper discusses the spatial and stratigraphic variations in the mineralogic, chemical and oxygen- and hydrogen-isotope compositions of 2:1 phyllosilicate and kaolinite mixtures from 11 Middle–Late Pennsylvanian (Desmoinesian–Missourian) paleosol profiles across the Illinois basin. Following the methodological approach discussed in Tabor and Montañez (2005), this work evaluates the utility of the mineralogic, chemical and isotopic properties of Illinois basin paleosol phyllosilicates as proxies for paleotemperature and Pennsylvanian soil water δ18O values. In addition, the influence of burial history and tectonic activity is considered as a potential mechanism for diagenetic alteration of the mineralogic and isotopic composition among some of these phyllosilicate mixtures. Finally, the data set provides best estimates of Pennsylvanian phyllosilicate crystallization temperatures from a shallowly-buried part of the Illinois basin. These data (1) demonstrate the isotopic composition of paleosol phyllosilicates that have experienced shallow burial and minimal post-depositional heating likely retain their original pedogenic values and provide a reliable source of paleoenvironmental proxy information and (2) provide the first quantitative paleotropical terrestrial temperature estimates across the Desmoinesian–Missourian boundary, an interval in the Pennsylvanian marked by significant paleoenvironmental changes, but whose mechanistic links to LPIA glacioeustasy remain debated (e.g., Frakes et al., 1992; Fielding et al., 2008a,b,c; Bishop et al., 2010; Gulbranson et al., 2010; Falcon-Lang et al., 2011). 2. Geologic background 2.1. Basin history The Illinois basin is an intracratonic basin located in Illinois, Indiana, and Kentucky (Fig. 1). The basin began subsiding in the late Precambrian–Early Cambrian in response to extension in the Reelfoot rift and Rough Creek graben associated with late Paleozoic Appalachian–Ouachita tectonic activity (Kolata and Nelson, 1991, 1997). Structural features separating the Illinois basin from adjacent provinces include the Kankakee Arch to the northeast, the Wisconsin Arch to the north, the Ozark Dome to the southwest, and the Cincinnati Arch to the east (Fig. 1; Willman et al., 1975). During the Pennsylvanian, the Illinois basin was located in western equatorial Pangaea (~0 to 5°N; Scotese et al., 1999; Blakey, 2007). Pennsylvanian biostratigraphic zonation in the basin, based on fusulinids (Dunbar and Henbest, 1942; Douglass, 1987), ostracods (Thompson et al., 1959; Thompson and Shaver, 1964), spores (Winslow, 1959; Peppers, 1964, 1970, 1996) and conodonts (Heckel and Baesemann, 1975; Swade, 1985; Heckel, 1991; Heckel and Weibel, 1991), permit regional and global correlations to Middle–Upper Pennsylvanian strata in contemporaneous basins (Ritter et al., 2002; Barrick et al., 2004; Heckel et al., 2007; Falcon-Lang et al., 2011). The sedimentary section in the Illinois basin preserves a nearly continuous record of Cambrian–Pennsylvanian strata. Pennsylvanian strata

383

comprise the youngest outcrops in the basin, except for abundant Quaternary sediments and Mesozoic strata at the southern margin of the basin (Willman et al., 1975). The greatest degree of Pennsylvanian subsidence and sediment accumulation occurred in the southeastern part of the Illinois basin, resulting in a maximum stratigraphic thickness of approximately 760 m (Fig. 1; Willman et al., 1975). 2.2. Tectonic and burial history The upper contact of Pennsylvanian strata is erosional (Willman et al., 1975) indicating maximum burial depths in excess of modern (note that strata studied herein are from bore-hole cores). Additionally, vitrinite reflectance values for the Middle Pennsylvanian (upper Carbondale Formation) Herrin Coal suggest significantly higher thermal maturities than can be satisfactorily explained by current burial depths (Fig. 1; Damberger, 1971, 1974). To account for this inconsistency, it is estimated that 1500 to 3000 m of southeastward-thickening Permian to Lower Cretaceous strata were deposited upon the Pennsylvanian strata, and later removed by erosion, in the thickest areas of the basin (Damberger, 1971, 1974; Zimmerman, 1986; Buschbach and Kolata, 1991; Gharrabi and Velde, 1995; Rowan et al., 2002). Permian strata associated with the Rough Creek graben in western Kentucky support this postPennsylvanian burial scenario (Fig. 1; Schwalb, 1982; Kehn et al., 1982). From the Cambrian through the Cretaceous the Illinois basin was intermittently affected by magmatic activity in the Reelfoot rift (Fig. 1). Radiometric ages of intrusive igneous rocks reveal a particularly substantial interval of magmatism during the middle Permian (~270–280 Ma). This activity was concentrated in southern Illinois and resulted in the emplacement of numerous ultramafic dikes and sills in the southern part of the basin and northern Reelfoot rift (Zartman, 1977; Brannon et al., 1992; Chesley et al., 1994). Contemporaneous fluorite mineralization occurred in the Fluorspar district in the vicinity of the cryptovolcanic Hicks dome (Fig. 1; Zartman, 1977; Nelson and Lumm, 1987; Rowan and Goldhaber, 1996). Collectively, these studies indicate that the Illinois basin experienced a complex tectonic and thermal history that was primarily focused in the southern part of the basin. Estimated maximum burial depth and temperatures for basal Pennsylvanian strata in the southern part of the basin are ~3000 m and N 175 °C, respectively (Harris, 1979; Elliott and Aronson, 1993; Grathoff et al., 2001). However, temperature estimates from fluid inclusions in diagenetic sphalerites in Middle Pennsylvanian coals (Damberger, 1971, 1974, 1991; Cobb, 1981; Coveney et al., 1987), vitrinite reflectance values (Barrows and Cluff, 1984; Comer et al., 1994), and smectite to illite mineral transformations (Gharrabi and Velde, 1995; Grathoff et al., 2001; Rowan et al., 2002) indicate that the Pennsylvanian strata in the northern part of the basin experienced a maximum burial depth of ~1500 m (approximately equivalent to eroded post-Pennsylvanian strata) and burial temperatures ranging from 70 to 120 °C. 2.3. Previous work and samples used in this study Pennsylvanian strata of the Illinois basin are characterized by a series of repeated successions of terrestrial, deltaic and marine deposits called cyclothems (Udden, 1912; Wanless and Weller, 1932). Cyclothems are commonly interpreted to result from repeated large-scale (≥100 m), high-frequency (104 to 105 kyr) eustatic sea-level fluctuations associated with the waxing and waning of ice sheets in Gondwananland (Heckel, 1977, 1986, 1994; Soreghan, 1994; West et al., 1997; Soreghan and Giles, 1999; Haq and Schutter, 2008). Pennsylvanian strata contain abundant pedogenically-modified horizons (paleosols; Hopkins and Simon, 1975; Greb et al., 1992; Mastalerz and Shaffer, 2000; Rosenau et al., 2013a,b) which developed on paleolandscapes during periods of glacio-eustatic lowstand and relative landscape stability. Based upon a comprehensive study of the morphological, mineralogical and geochemical content of paleosol profiles preserved within Pennsylvanian (Atokan–Virgilian; Westphalian–Stephanian) strata

384

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Kankakee Arch

0.50 0.55 0.60 `

0.65 0.70 0.55

0.75

Ozark Uplift

Fluorspar district t

Re

e

o lfo

t rif

Rough Creek graben

Cinc. Arch

Fig. 1. Map showing the location of the Illinois basin, adjacent structural features, and the Fluorspar district. Also shown are vitrinite reflectance contours (%Ro) for the Middle Pennsylvanian (Carbondale Formation) Herrin Coal. Filled circles represent the locations of cores from which samples were analyzed for this study. Cinc. Arch = Cincinnati Arch. Figure Modified from Cluff and Byrnes (1991) and Damberger (1971).

of the Illinois basin, Rosenau et al. (2013a) identified seven distinct paleosol types, or pedotypes (Retallack, 1990, 1994) that display significant morphological variability across the basin. The majority of paleosols preserved in strata from the interior of the basin preserve low chroma (gley) colors, abundant fossil plant debris, as well as calcite, sphaerosiderite and pyrite cements and nodules. In contrast, paleosols in strata from the northern margin of the basin display high chroma colors and pedogenic calcite, and considerably less fossil organic matter, sphaerosiderite and pyrite cements. Based on the spatial distribution of paleosol types and redox sensitive mineral assemblages (e.g., siderite, pyrite, iron oxide concentrations and depletions) and morphologies, Rosenau et al. (2013a) suggested that paleosols from the basin interior underwent a multiphase (polygenetic) pedogenic history characterized by an initial well-drained, oxygenated stage under seasonal precipitation, followed by a poorly-drained, reducing stage, while paleosols from the northern margin of the basin experienced overall more oxygenated conditions throughout their pedogenic development. Considering the different inferred pedogenic and burial histories of paleosols across the basin, paleosol phyllosilicate samples were collected from three continuous drill cores representing the northern (LSC core) and interior parts (MAC and ELY cores) of the basin (Figs. 1 and 2). The morphological characteristics of the paleosols from which phyllosilicates were analyzed for this study are discussed below. Twenty phyllosilicate samples were isolated from 11 different paleosols representing six different paleosol types (Tables 1–3). Using the paleosol-specific classification scheme of Mack et al. (1993), these

paleosol types are (1) a gleyed Protosol, (2) a gleyed Vertisol, (3) gleyed calcic Vertisols, (4) a gleyed vertic Calcisol, (5) a gleyed Calcisol, and (6) calcic Vertisols (Rosenau et al., 2013a,b; Tables 1 and 3). The gleyed Protosol is characterized by an abundance of fossil organic matter and weakly-developed horizonation. The gleyed Vertisol, gleyed calcic Vertisol, and gleyed vertic Calcisols preserve morphological properties similar to modern Vertisols, including wedge-shaped aggregate structure and pedogenic slickensides. These features were generated by shrink– swell processes within a climate characterized by seasonal precipitation and/or episodic changes in the position of the water table (Rosenau et al., 2013a,b). These paleosols also contain evidence for post-depositional waterlogging, including the presence of drab-hued mottles, and ferrous (Fe2 +) minerals, such as pyrite and siderite. The gleyed calcic Vertisols, gleyed vertic Calcisol, and gleyed Calcisol preserve similar morphologies to gleyed Vertisols; however, these paleosol types also contain accumulations of pedogenic calcite distributed throughout the profile. The gleyed Calcisol preserves accumulations of pedogenic calcite but lacks evidence for shrink–swell processes. Calcic Vertisols preserve similar morphological features to gleyed calcic Vertisols and the gleyed vertic Calcisol; however, calcic Vertisols preserve high-chroma matrix colors and lack secondary accumulations of siderite and/or pyrite suggesting that these paleosols never experienced extensive periods of waterlogging. The occurrence of pedogenic calcite in these paleosols suggests that these profiles formed on well-drained, stable portions of the Illinois basin landscape under sub-humid to semi-arid climates (Rosenau et al., 2013a,b).

~304 Ma

Fm. MATTOON

SubSystem Global Stage N. Am. Series

N.A. Rosenau, N.J. Tabor / Palaeogeography, Palaeoclimatology, Palaeoecology 392 (2013) 382–397

Member

3. Methods

Lithology

Cohn Coal

306.7 Ma

Flannigan Coal

BASHK. ATO KAN TRAD E WATE R

L

Carthage Ls. Stark Shale

Macoupin Ls. Hushpuckney Sh.

Carlinville Ls. Mound City Shale Chapel Coal

L

E

M

Exline Ls.

M L

Attila Shale

L

Danville Coal Crown Mine Ss. Bankston Fork Ls. Anna Shale Energy Shale Herrin Coal Briar Hill Coal Turner Mine Sh. Springfield Coal Galatia Ss.

Excello Shale Houchin Creek Coal

Elysium Energy Core (ELY)

Faribanks Coal.

Survant Coal Mecca Quarry Sh. Colchester Coal

L

Seelyville Coal

Wise Ridge Coal

314.6 Ma

L E

Lone Star Cement Core (LSC) Macoupin County Core (MAC)

305.7 Ma

P E N N S Y L V A N I A N M O S C O V I A N DESMOINESIAN CARBONDALE SHELBURN PATOKA

KASIMOVIAN MISSOURIAN BOND

Livingston Ls.

Hall/Reel Ls. Quivira Shale Flat Creek Coal Bunje Ls.

L

Murphysboro Coal

Delwood Coal Curlew Limestone

Phyllosilicates

L LSC M MAC E ELY

Murray Bluff Ss.

Reynoldsburg Coal

385

Coal Underclay Shale Siltstone Sandstone Limestone

Fig. 2. Generalized stratigraphic column of the Pennsylvanian sub-system in Illinois illustrating units that are discussed in the text (i.e., upper Tradewater–lower Mattoon formations). Select members are displayed for points of reference. Black dots with inset lettering indicate the position of paleosols from which phyllosilicates were analyzed for this study (L = LSC core; M = MAC core; E = ELY core. Global stage boundary and North American series dates are based on North American cyclothem calibration of Heckel (2008) in conjunction with radiometric (U–Pb) dates of zircons from the Donets Basin of Eastern Europe (Schmitz and Davydov, 2012). Figure modified from Willman et al. (1975) and Nelson et al. (2011).

Bulk paleosol samples were disaggregated by ultrasonic agitation in deionized water and the b2 μm equivalent spherical diameter size fraction was isolated from the matrix by centrifugation (Jackson, 2005). The b2 μm size fraction was processed further to isolate only the b 0.2 μm size fraction because this size fraction is commonly regarded to be dominated by pedogenic clays in paleosol profiles (Stern et al., 1997; Tabor et al., 2002; Vitali et al., 2002; Tabor and Montañez, 2005). The b0.2 μm size fraction was then treated in a set of selective dissolution procedures to remove non-phyllosilicate material: (1) 10% acetic acid solution for at least 24 h to remove calcite (Savin and Epstein, 1970; Lawrence and Taylor, 1971), (2) sodium citrate–bicarbonate–dithionite solution to remove secondary iron oxy-hydroxides, and (3) 30% H2O2 solution for at least 25 days to remove organic matter. Each aliquot of the b 0.2 μm size fraction was prepared on filter membranes and transferred to glass slides as oriented aggregates. A set of chemical and heat treatments were performed upon different aliquots of the size fractions, including (1) K+ saturation at room temperature, (2) Mg2+ saturation at room temperature, (3) Mg2+ saturation and glycerol solvation at room temperature, and (4) K+ saturation and heating at 500 °C for at least 2 h. Step-scan analyses of each treatment were performed on a Rigaku Ultima III X-ray diffractometer in the Huffington Department of Earth Sciences at Southern Methodist University (SMU) using Cu-Kα radiation over a range of 2° to 30°-2θ, a step size of 0.04°-2θ, and 1 second count time per step. Mineralogical identification from X-ray diffraction (XRD) spectra follows the procedures outlined in Moore and Reynolds (1997). To facilitate characterization of mixed-layer phases (e.g., illite–smectite; I/S) and permit an estimate of the proportion of each phase in the mixed-layer minerals, representative samples prepared as oriented aggregates were solvated in a vapor of ethylene glycol at 60 °C for at least 8 h. Estimates of the relative abundance (wt.%) of 2:1 phyllosilicates and kaolinite in the b0.2 μm size fraction were determined using the area of the 001 d(hkl) peak on XRD patterns of Mg2+-saturated samples with the background removed. The estimated uncertainty associated with this method is approximately ±5 wt.% (Moore and Reynolds, 1997). Samples were then split into three aliquots for chemical and stable isotope analysis. One aliquot of the b 0.2 μm fraction was prepared as a pressed powder pellet using 0.25 μm polished hardened steel dies and then coated with ~150-Å layer of high-purity carbon in a carbon evaporator. The pressed pellets were analyzed for major and minor elemental composition using a Cameca SX50 electron microprobe housed in the Department of Geology and Geophysics at Texas A&M University. Several Source Clay mineral standards, whose chemical compositions are well characterized (KGa 1-b and SAz-2; Van Olphen and Fripiat, 1979), were prepared as pressed pellets to assess the precision of the pressed powder method we follow herein. Normalized wt.% elemental oxides measured on the pressed powders are within 1% of the values reported in Van Olphen and Fripiat (1979). A second aliquot of the b0.2 μm fraction was heated in Nickel-rod bombs connected to gas extraction lines. Samples were heated at 100–150 °C overnight in order to remove sorbed and interlayer water. Samples were then reacted with BrF5 at 560 °C overnight to produce O2 gas following the methods of Clayton and Mayeda (1963). The O2 gas was quantitatively converted to CO2 using heated graphite rods in high-vacuum glass extraction lines and cryogenically collected to determine CO2 yield via a mercury manometer. The δ18O values of CO2 were measured on a Finnigan MAT 252 isotope ratio mass spectrometer (IRMS) operating in dual inlet mode at SMU. Repeated oxygen isotope analyses of NBS-28 over the course of study yield an analytical uncertainty of 0.2‰. The third aliquot was analyzed for stable hydrogen isotope composition following the methods of Savin and Epstein (1970). Samples were initially outgassed at ~ 125 °C for at least 75 min under

N.A. Rosenau, N.J. Tabor / Palaeogeography, Palaeoclimatology, Palaeoecology 392 (2013) 382–397

Kaolinite (wt. %)

1 1 7 1 4 2 3 28 45 17 52 47 65 49 90 55 21 7 14 32 80 60 65 55 60 55 40 40 30 15 20 15 10 35 20 20 b10 b10 b10 b10

0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 22 19 26 41

open-system conditions to remove sorbed and interlayer water. The samples were then dehydroxylated at ~ 825 °C under closed-system conditions in 0.16 bar of O2 gas. The low- and high-temperature water fractions were then quantitatively converted to H2 gas by passage over depleted uranium metal at ~ 760 °C as two separate fractions. H2 gas yields were measured on a mercury manometer with an uncertainty of ± 1 μmol, and the δD values of H2 were measured on a Finnigan MAT 252 IRMS operating in dual inlet mode at SMU. Replicate hydrogen isotope analyses of phyllosilicate samples over the course of study yield an analytical uncertainty of ± 4‰. Both the oxygen and hydrogen isotopic composition of the b0.2 μm fraction are reported in the conventional delta notation (δ) in parts per thousand (‰) relative to Vienna Standard Mean Ocean Water (V-SMOW; Gonfiantini, 1984) where  2
18 16 3 O =Osample 4  18 16  −15  1000 δ O¼ O =Ostd 18

99 99 93 99 96 98 97 72 55 83 48 53 35 51 10 45 57 74 60 27

ð1Þ

2


 3 2 1 H =Hsample δD ¼ 4  2 1  −15  1000: H =Hstd

ð2Þ

d

c

b

Chemical data represents the average of 5 individual analyses of each sample. Classification scheme of Mack et al. (1993). Meter levels refer to the stratigraphic distance as measured from the top of the core (0 m = core top). Clay mineralogy of the b0.2 μm phyllosilicate size fraction organized in relative order of abundance; K = kaolinite; I = illite, I/S = illite-smectite.

3.1. Phyllosilicate oxygen and hydrogen isotope fractionation factors

a

I/S, K I/S, K I/S, K I/S, K I/S, K I/S, K I/S, K I/S, K I/S, K I/S, K K, I/S I/S, K K, I/S K, I/S K, I/S K, I/S I/S, K, I I/S, K, I I/S, K, I I/S, K, I 95.16 95.39 95.07 94.08 94.18 92.52 93.89 93.53 89.60 90.29 92.25 92.34 91.92 91.79 90.12 91.22 93.74 91.96 91.63 92.76 0.17 0.34 0.20 0.18 0.05 0.04 0.06 0.11 0.19 0.15 0.14 0.08 0.05 0.05 0.11 0.11 0.10 0.07 0.05 0.04 3.05 2.88 3.11 2.38 2.28 2.20 2.02 1.52 1.11 1.09 1.45 1.50 0.99 1.01 0.85 1.45 2.60 1.82 2.18 2.29 Missourian Missourian Missourian Missourian Missourian Missourian Missourian Desmoinesian Desmoinesian Desmoinesian Desmoinesian Desmoinesian Desmoinesian Desmoinesian Desmoinesian Desmoinesian Missourian Missourian Missourian Desmoinesian LSC17 N LSC16 N LSC15 N LSC13 N LSC12 N LSC11 N LSC5 N LSC2 N LSC1 N LSC24 N LSC25 N LSC21 N LSC62 N LSC56 N LSC54 N LSC52 N ELY1-45 ELY1-72 MAC1-18 MAC1-21

Bond Bond Bond Bond Bond Bond Patoka Shelburn Shelburn Shelburn Shelburn Shelburn Carbondale Tradewater Tradewater Tradewater Bond Patoka Patoka Shelburn

calcic Vertisol (P5) calcic Vertisol (P5) calcic Vertisol (P5) calcic Vertisol (P4) calcic Vertisol (P4) calcic Vertisol (P4) calcic Vertisol (P2) gleyed Protosol (P1) gleyed Protosol (P1) calcic Vertisol (P7) calcic Vertisol (P7) calcic Vertisol (P7) gleyed vertic Calcisol (P16) gleyed Vertisol (P15) gleyed Vertisol (P15) gleyed Vertisol (P15) gleyed Calcisol (P9) gleyed Calcisol (P14) gleyed calcic Vertisol (P2) gleyed vertic Calcisol (P3)

3.7 4.5 5.1 12.2 12.8 14.6 27.5 30.2 31.1 35.4 35.6 36.7 124.7 128.1 128.8 130.3 159.7 252.7 47.1 69.1

0.24 0.51 0.35 0.25 1.89 1.89 1.67 0.23 1.45 0.93 0.86 1.15 0.12 0.09 0.24 0.31 0.29 0.34 0.24 0.20

25.67 24.63 24.45 23.75 25.47 25.58 25.43 30.11 31.17 29.78 29.42 30.51 32.46 31.48 33.45 26.24 27.81 29.12 27.84 27.82

59.13 58.46 57.69 56.40 54.82 53.00 54.09 53.57 50.02 49.02 50.50 50.64 51.99 51.96 50.00 51.93 53.95 52.16 52.39 53.15

2.52 3.81 3.66 4.11 3.69 3.65 4.40 2.45 1.22 4.55 4.75 4.06 2.10 2.38 1.36 4.42 4.50 4.89 4.55 4.99

3.81 4.19 4.95 6.70 5.35 5.50 5.56 4.66 2.84 3.72 3.98 3.64 2.60 3.58 2.49 5.25 3.42 2.64 3.66 3.59

0.58 0.58 0.66 0.31 0.63 0.65 0.66 0.88 1.59 1.04 1.15 0.77 1.60 1.25 1.63 1.51 1.07 0.93 0.72 0.68

Mineralogyd CaO K2O SiO2 Al2O3 MgO Na2O Depth (m)c Paleosol type (no.)b Formation North American Series Sample

Table 1 Electron microprobe and X-ray diffraction data for the b0.2 μm size fraction of Pennsylvanian phyllosilicates in the Illinois basin.a

Fe2O3

TiO2

Total

I/S (wt.%)

Smectite in I/S (%)

Illite (wt. %)

386

3.1.1. Kaolinite Based on revisions of existing empirical and experimental data, Sheppard and Gilg (1996) proposed the following oxygen and hydrogen isotope fractionation equations between kaolinite and water, which are employed herein: 18

6

2

ð3Þ

D

6

2

ð4Þ

1000 ln αkaolinite–water ¼ 2:76  10 =T –6:75 1000 ln αkaolinite–water ¼ −2:2  10 =T −7:7;

where 18α and Dα are the oxygen and hydrogen isotope fractionation factors between kaolinite and water, respectively, and temperature is in degrees Kelvin. 3.1.2. 2:1 phyllosilicates At present, there is a lack of accurately determined oxygen isotope fractionation factors for most 2:1 phyllosilicate minerals. This is due in large part to (1) the difficulty in isolating pure end-members from most naturally occurring phyllosilicate mixtures, and (2) the highly-variable chemical composition of naturally occurring 2:1 phyllosilicates that form in low-temperature environments, which has been shown to significantly affect 2:1 phyllosilicate-water oxygen and hydrogen isotope fractionation (Savin and Epstein, 1970; Lawrence and Taylor, 1971, 1972). An alternative approach for determining the oxygen isotope fractionation among phyllosilicate minerals and water is the bond-model technique proposed by Savin and Lee (1988). This approach considers phyllosilicate oxygen isotope fractionation to depend only upon the type of element to which the oxygen is bonded in the crystal lattice, not the physical structure of the mineral, and by extension, the oxygen isotopic fractionation between phyllosilicates and water may be expressed as the weighted sum of oxygen isotope fractionation values of the different oxygensharing bonds in that mineral (Savin and Lee, 1988). Herein, the bondmodel method of Savin and Lee (1988) is used to calculate mixed-layer 2:1 phyllosilicate–water oxygen isotope fractionation factors (Table 3). Hydrogen isotope fractionation factors for 2:1 phyllosilicates are likewise poorly constrained (Yeh, 1980; Capuano, 1992; Sheppard and Gilg, 1996). Studies of naturally occurring samples indicate that the hydrogen isotope fractionation factor for the 2:1 phyllosilicate, smectite, is largely controlled by its octahedral cation chemistry (Gilg and Sheppard, 1995). Considering this, as well as the work by Suzuoki and Epstein

N.A. Rosenau, N.J. Tabor / Palaeogeography, Palaeoclimatology, Palaeoecology 392 (2013) 382–397 Table 2 Calculated chemical formulae for end-member 2:1 phyllosilicates in phyllosilicate mixturesa. Sample

Depth (m)b

Chemical Formulae

LSC-52N LSC-54N LSC-56N LSC-62N LSC-21N LSC-25N LSC-24N LSC-1N LSC-2N LSC-5N LSC-11N LSC-12N LSC-13N LSC-15N LSC-16N LSC-17N ELY1-45 ELY1-72 MAC1-18 MAC1-21

130.3 128.8 128.1 124.7 36.7 35.6 35.4 31.1 30.2 27.5 14.6 12.8 12.2 5.1 4.5 3.7 159.7 252.7 47.1 69.1

(K0.71Na0.7Ca0.015)(Al1.13Fe0.50Mg0.27Ti0.14)(Si3.21Al0.79)O10(OH)2 (K0.43Na0.12Ca0.03)(Al1.43Fe0.46Mg0.31Ti0.30)(Si1.88Al2.12)O10(OH)2 (K0.33Na0.02Ca0.01)(Al1.64Fe0.29Mg0.16Ti0.10)(Si3.10Al0.90)O10(OH)2 (K0.37Na0.03Ca0.01)(Al1.61Fe0.27Mg0.20Ti0.17)(Si2.88Al1.12)O10(OH)2 (K0.56Na0.24Ca0.01)(Al1.50Fe0.30Mg0.24Ti0.06)(Si3.04Al0.96)O10(OH)2 (K0.72Na0.20Ca0.02)(Al1.35 Fe0.36Mg0.26Ti0.10)(Si3.00Al1.00)O10(OH)2 (K0.46Na0.14Ca0.01)(Al1.65Fe0.22Mg0.13Ti0.06)(Si3.25Al0.75)O10(OH)2 (K0.16Na0.30Ca0.02)(Al1.63Fe0.23Mg0.17Ti0.13)(Si3.07Al0.93)O10(OH)2 (K0.26Na0.04Ca0.01)(Al1.62Fe0.29Mg0.19Ti0.06)(Si3.33Al0.67)O10(OH)2 (K0.38Na0.22)(Al1.49Fe0.28Mg0.20Ti0.03)(Si3.54Al0.46)O10(OH)2 (K0.31Na0.25)(Al1.50Fe0.28Mg0.22Ti0.03)(Si3.51Al0.49)O10(OH)2 (K0.32Na0.25)(Al1.49Fe0.27Mg0.23Ti0.03)(Si3.55Al0.45)O10(OH)2 (K0.34Na0.03Ca0.01)(Al1.47Fe0.33Mg0.23Ti0.02)(Si3.66Al0.34)O10(OH)2 (K0.32Na0.05Ca0.01)(Al1.48Fe0.25Mg0.31Ti0.03)(Si3.67Al0.33)O10(OH)2 (K0.31Na0.06Ca0.02)(Al1.54Fe0.20Mg0.27Ti0.03)(Si3.70Al0.30)O10(OH)2 (K0.20Na0.03Ca0.01)(Al1.60Fe0.18Mg0.29Ti0.03)(Si3.71Al0.29)O10(OH)2 (K0.45Na0.04Ca0.01)(Al1.53Fe0.20Mg0.30Ti0.06)(Si3.43Al0.57)O10(OH)2 (K0.44Na0.05Ca0.01)(Al1.70Fe0.14Mg0.19Ti0.05)(Si3.43Al0.57)O10(OH)2 (K0.44Na0.03)(Al1.60Fe0.21Mg0.24Ti0.04)(Si3.44Al0.56)O10(OH)2 (K0.57Na0.03)(Al1.50Fe0.24Mg0.30Ti0.05)(Si3.38Al0.62)O10(OH)2

a Chemical formulae were calculated from the oxide data reported in Table 1, based on the presence of twelve oxygen atoms and two hydrogen atoms in each unit-cell (Moore and Reynolds, 1997). b Meter levels refer to the stratigraphic distance as measured from the top of the core (0m = core top).

(1976), Tabor and Montañez (2005) proposed that hydrogen isotope fractionation factors for naturally occurring smectites encompassing a range of chemical compositions could be approximated by the following equation:
 D 6 2 1000 ln α ¼ −2:2  10 =T –7:7 þ 2XAl –4XMg –68XFe

ð5Þ

where XAl, XMg, and XFe are the mole fractions of aluminum, magnesium, and iron, respectively, present in the octahedral layer of the 2:1 phyllosilicate mineral. Given the lack of better theoretical or experimental data, this relationship is used herein to calculate the hydrogen isotope fractionation factors of mixed layer 2:1 phyllosilicates (Table 3). 3.1.3. Illite Discrete illite is limited to four samples from the interior of the basin (MAC and ELY cores; Fig. 3). Illite is typically considered to be an inherited phase from the soil or a product of diagenesis (Southard and Miller, 1966; Yemane et al., 1996), but it nevertheless may also form through mineralogic transformations in soil profiles (Wilson, 1999). Similar to other 2:1 phyllosilicates, the chemical variability of illite is expected to have a significant effect on illite–water oxygen and hydrogen isotope fractionation factors (Suzuoki and Epstein, 1976; Savin and Lee, 1988; Bechtel and Hoernes, 1990). Accordingly, the bond-model approach of Savin and Lee (1988) and hydrogen isotope fractionation factor equation of Tabor and Montañez (2005) are used in this study to calculate the oxygen and hydrogen isotope fractionation factors for illite in the mixtures of 2:1 phyllosilicate and kaolinite (Table 3).

387

mineralogy of Pennsylvanian paleosols in the Illinois basin can be found in Rosenau et al. (2013a). The samples are composed of mixtures of 2:1 phyllosilicates and 1:1 kaolinite (Table 1; Figs. 3 and 4). 2:1 phyllosilicates include interlayered illite–smectite (I/S) with variable percentages of interlayer smectite, and rarely, discrete illite. Discrete illite is defined herein as a 100% nonexpansible 2:1 mica-like mineral (sensu Środoń, 1999) that displays a basal (001) 10 Å reflection on an X-ray diffraction spectrum (using Cu-Kα as a radiation source; Moore and Reynolds, 1997). Remarkably, different phyllosilicate–mineral assemblages are preserved in the b 0.2 μm fractions across the Illinois basin (Figs. 3 and 4). The b0.2 μm fractions of the upper two, stratigraphically-shallowest paleosols (Missourian; Bond Formation) from the northern part of the basin (LSC core; Fig. 4) are dominated by smectite-rich (55–80% smectite) I/S and relatively minor amounts of kaolinite (b 10 wt.%; Fig. 4; Table 1). Stratigraphically-deeper in the LSC core (uppermost Desmoinesian; uppermost Shelburn Formation), the relative amount of kaolinite in the b0.2 μm fraction in paleosols increases, ranging from 17 to 52% and I/S becomes less smectite-rich (~ 15–40% smectite; Fig. 4). The stratigraphically-deepest samples (Desmoinesian; Carbondale and Tradewater formations) in the LSC core are enriched in kaolinite (49–90 wt.%) and illite-rich I/S (~ 10–35% smectite; Fig. 4; Table 1). The mineralogic composition of samples from the interior of the basin (MAC and ELY cores; Fig. 3) consists of variable amounts of illite-rich (b10% smectite) I/S, kaolinite (27–74 wt.%) and discrete illite (19–41 wt.%). Except for the presence of discrete illite, the mineralogy and I/S composition of phyllosilicate samples from the MAC and ELY cores is similar to the mineralogy and I/S composition observed in the stratigraphically-deepest Desmoinesian (Carbondale and Tradewater formations) paleosols in the LSC core (Table 1; Figs. 3 and 4). The chemical composition of end-member 2:1 phyllosilicates in the 2:1 phyllosilicate/kaolinite mixtures were calculated by considering that the wt.% kaolinite, as calculated from X-ray diffraction spectra, in each sample corresponds to a fraction of the Al2O3 and SiO2 in the measured chemical compositions of the b 0.2 μm phyllosilicate mixtures (Table 1). Owing to differences in the relative proportions of Al2O3 and SiO2 in the crystal structure of kaolinite and 2:1 phyllosilicates, it is necessary to consider the different mole fraction contributions of Al and Si from kaolinite and 2:1 phyllosilicates in the samples. The mole fraction of Al2O3 and SiO2 were calculated from the wt.% oxide data in Table 1 and the estimated mole fraction of Al and Si contributed from kaolinite were subtracted from the phyllosilicate mixtures to calculate an end-member 2:1 phyllosilicate chemical formula for each sample (Tabor and Montañez, 2005; Table 2). 4.2. Phyllosilicate oxygen and hydrogen isotopes The measured δ18OV-SMOW and δDV-SMOW values of paleosol phyllosilicates range from 17.2‰ to 23.0‰ and −56‰ to −27‰, respectively (Table 3). Phyllosilicate δ18O and δD values from laterally equivalent paleosols across the basin display distinct isotopic trends. Specifically, δ18O values progressively increase, and δD values progressively decrease, from the shallowly buried basin margin (LSC core) to more deeply-buried interior parts of the basin (MAC and ELY cores; Figs. 1 and 5). 5. Discussion

4. Results 4.1. Mineralogy and chemical composition

5.1. Stable isotope composition of paleosol phyllosilicates as paleoenvironmental proxies

Sample locations and the elemental and mineralogic composition of the b0.2 μm phyllosilicate fraction of the samples are presented in Table 1. A detailed discussion of the b 2 μm and b 0.2 μm phyllosilicate

The phyllosilicate fraction in paleosols can be mixtures of pedogenic clays, detrital and burial authigenic clays, as well as diagenetically altered pedogenic clays.

388

Table 3 Measured phyllosilicate oxygen and hydrogen-isotope compositions, oxygen and hydrogen isotope fractionation equations and calculated phyllosilicate crystallization temperatures. a

103ln18α2:1-water

103lnDα2:1-water = −2.2*106*T -2 +

b

c

103ln18αmix-water

d 103lnDαmix-water = −2.2⁎106⁎T -2 +

North American Series

Formation

Paleosol Type (No.)

Horizon

LSC–17N LSC–16N LSC–15N LSC–13N LSC–12N LSC–11N LSC–5N LSC–2N LSC–1N LSC–24N LSC–25N LSC–21N LSC–62N LSC–56N LSC–54N LSC–52N

Missourian Missourian Missourian Missourian Missourian Missourian Missourian Desmoinesian Desmoinesian Desmoinesian Desmoinesian Desmoinesian Desmoinesian Desmoinesian Desmoinesian Desmoinesian

Bond Bond Bond Bond Bond Bond Patoka Shelburn Shelburn Shelburn Shelburn Shelburn Carbondale Tradewater Tradewater Tradewater

calcic Vertisol (P5) calcic Vertisol (P5) calcic Vertisol (P5) calcic Vertisol (P4) calcic Vertisol (P4) calcic Vertisol (P4) calcic Vertisol (P2) gleyed Protosol (P1) gleyed Protosol (P1) calcic Vertisol (P7) calcic Vertisol (P7) calcic Vertisol (P7) gleyed vertic Calcisol (P16) gleyed Vertisol (P15) gleyed Vertisol (P15) gleyed Vertisol (P15)

Bss1 Bss2 Bkss Bkss Bkssg Bkg Bkss ABwg Bwg Bk1 Bk2 Bw2 Bwg ABssg Bssg1 Bssg4

23.0 21.9 22.0 22.4 20.8 20.6 20.6 21.2 20.5 19.4 19.6 18.9 21.0 21.3 20.7 21.0

−38 −41 −43 −54 −43 −50 −51 −48 −51 −45 −44 −49 −53 −56 −51 −53

1 1 7 1 4 2 3 28 45 17 52 47 65 49 90 55

2.89*106*T-2–6.90 2.88*106*T-2–6.90 2.85*106*T-2–6.83 2.85*106*T-2–6.94 2.87*106*T-2–7.13 2.87*106*T-2–7.20 2.88*106*T-2–7.16 2.89*106*T-2–7.50 2.91*106*T-2–7.84 2.94*106*T-2–7.78 2.82*106*T-2–7.96 2.87*106*T-2– 7.94 2.88*106*T-2–8.21 2.90*106*T-2–7.87 2.80*106*T-2–9.92 2.73*106*T-2–7.44

−12.66 −13.52 −15.24 −17.73 −15.89 −16.11 −16.27 −16.00 −13.96 −13.85 −19.17 −16.67 −15.33 −16.02 −21.29 −25.00

2.88*106*T-2–6.90 2.88*106*T-2–6.90 2.84*106*T-2–6.82 2.85*106*T-2–6.94 2.87*106*T-2–7.12 2.87*106*T-2–7.19 2.87*106*T-2–7.15 2.85*106*T-2–7.27 2.84*106*T-2–7.32 2.90*106*T-2–7.59 2.79*106*T-2–7.28 2.81*106*T-2–7.34 2.80*106*T-2–7.22 2.82*106*T-2–7.29 2.76*106*T-2–7.02 2.75*106*T-2–7.04

−12.52 −13.35 −13.87 −17.44 −15.00 −15.63 −15.54 −11.54 −9.52 −11.50 −10.30 −10.10 −8.85 −9.84 −8.16 −11.26

ELY1–45 ELY1–72 MAC1–18 MAC1–21

Missourian Missourian Missourian Desmoinesian

Bond Patoka Patoka Shelburn

gleyed Calcisol (P9) gleyed Calcisol (P14) gleyed calcic Vertisol (P2) gleyed vertic Calcisol (P3)

Bssg Bkg Bkg2 Bkg2

19.2 18.2 19.2 19.0

−27 −34 −35 −38

21 7 14 32

2.87*106*T-2–7.22 2.94*106*T-2–7.40 2.90*106*T-2–7.35 2.86*106*T-2–7.40

−13.52 −11.08 −13.45 −14.86

2.84*106*T-2–7.11 2.93*106*T-2–7.35 2.88*106*T-2–7.25 2.83*106*T-2–7.17

−10.94 −10.46 −11.57 −10.64

Wt. % Kaolinite (± 5%)

e T°C (±3 °C)

f 18

δ Ometeoric water V-SMOW

Surface Domain

29 33 31 25 37 34 34 29 28 39 35 36 23 22 24 23

−1.9 −2.1 −2.1 −2.9 −2.1 −2.8 −2.9 −3.0 −3.7 −3.0 −2.7 −3.4 −3.9 −4.1 −3.8 −3.5

WESD WESD WESD MSD WESD WESD WESD MSD MSD non-sd WESD non-sd MSD MSD MSD MSD

51 55 48 44

−0.9 −1.8 −1.7 −2.1

outside SD's outside SD's outside SD's outside SD's

a Calculated oxygen isotope fractionation equations between the b0.2 μm 2:1 phyllosilicate-size fraction and water. Fractionation equations were calculated using the bond-model data of Savin and Lee (1988) in conjunction with the calculated chemical composition the end-member 2:1 phyllosilicates presented in Table 2. b Calculated hydrogen isotope fractionations between the b0.2 μm 2:1 phyllosilicate-size fraction and water. Fractionation equations were calculated based on the molar fraction of Al, Mg and Fe in the octahedral layer (Eq. (5)) as proposed by Gilg and Sheppard (1995) and Tabor and Montañez (2005). c Calculated oxygen isotope fractionation factors between mixed phyllosilicate mineralogies in the b0.2 μm size fraction and water. Fractionation factors were calculated based on the molar fraction of oxygen contributed from kaolinite and 2:1 phyllosilicates in each sample (Sheppard and Gilg,1996). d Calculated hydrogen isotope fractionation factors between mixed phyllosilicate mineralogies in the b0.2 μm size fraction and water. Fractionation equations were calculated based on the molar fraction of hydrogen contributed from kaolinite and 2:1 phyllosilicates in each sample (Sheppard and Gilg, 1996). e Calculated equilibrium crystallization temperatures for each phyllosilicate mixture (Eq. (7)). f Calculated 618O of meteoric water in equilibrium with the phyllosilicate mixtures (Eq. (8)).

N.A. Rosenau, N.J. Tabor / Palaeogeography, Palaeoclimatology, Palaeoecology 392 (2013) 382–397

δ18Omix V- δDmix VSMOW SMOW (±0.2%) (±4%)

Sample

N.A. Rosenau, N.J. Tabor / Palaeogeography, Palaeoclimatology, Palaeoecology 392 (2013) 382–397

Fulfillment of condition 1 is accomplished by isolating and analyzing only the b 0.2 μm size fraction because this size fraction is commonly regarded to be dominated by pedogenic clays in paleosol profiles (Stern et al., 1997; Tabor et al., 2002; Vitali et al., 2002; Tabor and Montañez, 2005). Conditions 2 and 3 are specific to the mineral of interest and that mineral's paragenetic history. Considering the processes governing the global meteoric water line (Craig, 1961; Rozanski et al., 1993; Yapp, 2000) and the low solubility of silicate minerals (Stumm and Morgan, 1981; Gregory, 1991), it is likely that conditions 4 and 5 have been constant for well-developed soils throughout geologic history. Nevertheless, even with satisfaction of the above conditions, these data may be assessed in several other ways. For example, it is possible that all of the phyllosilicates have been chemically- and isotopically altered from their original compositions during burial and lithification (i.e., diagenesis). Additionally, it is possible that authigenic phyllosilicates from a given paleosol may be composed of several different fractions that crystallized under different climate regimes throughout the development of the soil; several lines of evidence are presented in the following sections that likely discount the former possibility. However, the latter scenario cannot be absolutely ruled out, and therefore, phyllosilicate crystallization temperatures, and soil–water δ18O estimates determined from the phyllosilicate δ18O and δD values, may be best considered as representing the average paleoenvironmental conditions that persisted throughout active pedogensis. Ultimately, based upon the lines of evidence presented herein, pedogenesis and diagenesis are considered as the dominant processes responsible for the paleosol phyllosilicate δ18O and δD values reported here. 5.2. Suitability of interlayered illite–smectite (I/S) as a paleotemperature proxy The 2:1 phyllosilicate component of the mineralogic mixtures preserved in both the b2 μm and b0.2 μm size fractions are dominated by interlayered I/S (Figs. 3 and 4). The occurrence of I/S in soils and paleosols is not unusual (Grim, 1935; Barnhisel, 1977; Wilson, 1987; Moore and Reynolds, 1997), and I/S exhibiting a range of expansible layers and interlayer stacking patterns has been described from rocks, soils and paleosols that were never deeply buried or exposed to elevated temperatures (Schultz, 1978; Moore, 1982; Rimmer and Eberl, 1982; Deconinck et al., 1988; Turner and Fishman, 1991; Kirsimäe et al., 1999; Środoń, 1999; Huggett and Cuadros, 2003; Velde et al., 2003). Furthermore, it has been shown that (1) the isotopic composition of mixed-layer clay minerals are not altered during the physical and chemical pretreatments required to remove non-phyllosilicate materials (VanDeVelde and Bowen, 2013) and (2) mixed-layer clay minerals may preserve δ18O and δD values consistent with crystallization in low-temperature (b40 °C), near-surface weathering environments (Vitali et al., 2002; Gilg et al., 2003; Velde et al., 2003) 5.3. Calculation of oxygen and hydrogen isotope fractionation factors Oxygen isotope fractionation equations for end-member 2:1 phyllosilicate minerals in each sample were calculated following the bond-model approach of Savin and Lee (1988), while hydrogen isotope fractionation equations were calculated using the equation

I n t e n s it y

I/S

I

3.6Å

10Å

I/S

7.2Å

11.6Å

I/S

K

P9

ELY-45 I/S I/S

I K

I/S

R e la t iv e

(1) phyllosilicates analyzed from a given paleosol are authigenic, (2) phyllosilicates crystallized in isotopic equilibrium with soil water, (3) the phyllosilicates have not been diagenetically altered since the time of formation, (4) knowledge of the relationship between the oxygen and hydrogen isotope compositions of soil–water, and (5) phyllosilicates crystallized in a water-dominated system.

30Å

As such, in order to utilize the chemical and stable isotope composition of paleosol phyllosilicates as a paleotemperature proxy, several conditions are required:

389

P14

ELY-72

I I/S

K I

MAC-18

P2

MAC-21

P3

K

2

4

6

8

10

12

14

16

18

20

22

24

26

28

30

°2θ Fig. 3. X-ray diffraction spectra of Mg2+-saturated oriented aggregates of the b 0.2 μm phyllosilicate fraction of samples from the MAC and ELY cores that are discussed in the text. Spectra belonging to each paleosol are denoted by encircled letter–number combinations and follow the designation presented in Tables 1 and 3. Note the presence of illite in these samples. I/S = illite/smectite; K = kaolinite; I = illite.

proposed by Tabor and Montañez (2005; Table 3). Given that all samples analyzed for this study are mineralogic mixtures of 2:1 phyllosilicates and kaolinite, unique 2:1 phyllosilicate–water hydrogen and oxygen isotope fractionation equations may be calculated for each mineralogic mixture based on the molar fraction of oxygen and hydrogen contributed from 2:1 phyllosilicates and kaolinite (Tables 1–3). Savin and Epstein (1970) suggested that, under conditions of chemical equilibrium, the relationship between the oxygen and hydrogen-isotope compositions of hydroxyl-bearing minerals and meteoric water could be described by the following equation: "

# # "" # D αm‐w 18 αm‐w D δD ¼ 8 18 δ Om þ 1000 8  18 −6:99 αm‐w −1 αm‐w αm‐w D

ð6Þ

where Dαm-w and 18αm-w represent the temperature-dependent hydrogen and oxygen isotope fractionation factors between a specific mineral and water. If the temperature-dependent fractionation factors are relatively well known, then measurements of δDm and δ18Om of pedogenic hydroxylated minerals may provide an estimate of mineral crystallization temperature (Delgado and Reyes, 1996; Yapp, 1987, 1993, 2000; Tabor and Montañez, 2005; Tabor, 2007). Herein, this well-known relationship is utilized to calculate phyllosilicate crystallization temperatures. Assuming that the phyllosilicate minerals in each sample analyzed in this study crystallized under conditions of chemical equilibrium with meteoric waters that plot upon the global meteoric water line (GMWL, Craig, 1961) and considering the calculated temperaturedependent oxygen and hydrogen isotope fractionation equations, the following relationship may be used to relate the measured δ18O and δD values of the phylloslicate mixtures to crystallization temperatures:
 D 18 18 1000 ln αmix‐water –δDmix ¼ 8 1000 ln αmix‐water −δ Omix –10

ð7Þ

where Dαmix-water and 18αmix-water denote the temperature-dependent hydrogen and oxygen isotope fractionation, respectively, between the phyllosilicate samples and waters that reside upon the GMWL, and

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meteoric water from which the phyllosilicate crystallized (Table 3) through the following relationship:

LSC-17N

3.6Å

I/S

5Å

I/S

7.2Å

11.5Å

15Å


 18 18 18 18 δ Owater ¼ 1000 þ δ Omix –1000 αmix‐water = αmix‐water :

P5

ð8Þ

5.4. Calculation of phyllosilicate crystallization temperatures

I/S I/S

LSC-16N

P5

LSC-15N

P5

LSC-13N

P4

LSC-12N

P4

LSC-11N

P4

LSC-5N

P2

LSC-2N

P1

LSC-1N

P1

LSC-24N

P7

LSC-25N

P7

LSC-21N

P7

LSC-62N

P16

LSC-56N

P15

LSC-54N

P15

LSC-52N

P15

I/S

Relative Intensity

I/S

I/S

I/S

I/S

I/S I/S

I/S K I/S K I/S K I/S

K

I/S K I/S

K

I/S K

2

4

6

8

10 12 14 16 18 20 22 24 26 28 30

The measured δ18O and δD values of the b0.2 μm phyllosilicate fraction in combination with estimated oxygen and hydrogen isotope fractionation equations calculated for each sample results in phyllosilicate crystallization temperatures that range from 22 ± 3 °C to 55 ± 3 °C (Table 3, Fig. 6). The uncertainty of the temperaturedependent oxygen and hydrogen isotope fractionation factors between the phylloslicate minerals and water makes a precise estimate of the error around these paleotemperature estimates difficult. The uncertainty of ±3 °C reported herein for phyllosilicate crystallization temperatures reflects the analytical uncertainty of the measured phyllosilicate oxygen (±0.2‰) and hydrogen (±4‰) isotope values. Calculated crystallization temperatures of phyllosilicates preserved in laterally equivalent paleosols across the basin reveal a trend towards progressively higher crystallization temperatures from the northern margin of the basin (LSC core) to the interior parts of the basin (MAC and ELY cores; Table 3; Fig. 6). Phyllosilicate crystallization temperatures from the northern part of the basin (LSC core) range from 22 ± 3 °C to 37 ± 3 °C, with an average temperature of 29 ± 3 °C (Fig. 6). LSC phyllosilicate crystallization temperatures define a long-term stratigraphic trend from cooler average temperatures (23 ± 3 °C) in the Desmoinesian (upper Tradewater–lower Carbondale formations) to warmer average temperatures (31 ± 3 °C) in the uppermost Desmoinesian–lower Missourian (uppermost Shelburn–Bond formations; Fig. 6). Phyllosilicate crystallization temperature estimates from a locality in the central part of the basin (MAC) range from 44 ± 3 °C to 48 ± 3 °C, whereas temperatures estimates from the southernmost sampling locality in the basin (ELY) range from 51 ± 3 °C to 55 ± 3 °C (Table 3; Fig. 6). For context, mean annual soil temperatures in modern low-altitude (b 1000 m), low-latitude soil-forming environments range from 26 °C to 32 °C and these mean annual soil temperatures are typically ~ 2 °C warmer than mean annual surface-air temperatures (Chang, 1958).

δ18Omix and δDmix are the measured oxygen and hydrogen isotope values of the phyllosilicate samples (Table 3). The mineralogic and chemical variability of the phyllosilicate mixtures (Tables 1 and 2) result in a range of oxygen and hydrogen isotope fractionation equations for the studied samples, precluding the application of a single phyllosilicate mineral geothermometer equation, such as that proposed by Delgado and Reyes (1996) for a particular smectite mineralogy. In conjunction with their respective oxygen isotope fractionation equations, the calculated phyllosilicate crystallization temperatures also provide an independent estimate of the oxygen isotope composition of the

WL

: δD

=8

*δ 1 8 O+

10

10 5 0 -5 -10 -15 -20 -25 -30 -35 -40 -45 -50 -55 -60 -65

GM

Fig. 4. X-ray diffraction spectra of Mg2+-saturated oriented aggregates from the b 0.2 μm phyllosilicate fraction of samples from the LSC core that are discussed in the text. Spectra are organized with respect to their relative stratigraphic position (i.e., LSC-17N = stratigraphically highest sample). Spectra belonging to each paleosol are denoted by letter–number combinations and follow the designation presented in Tables 1 and 3. Light-gray and dark-gray spectra correspond to Desmoinesian samples and Missourian samples, respectively. The patterns reveal a down-section increase in the relative intensity of the 7.2 Å kaolinite peak. I/S = illite/smectite; K = kaolinite.

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5.5. Assessment of diagenesis Utilization of the hydrogen- and oxygen-isotope compositions of pedogenically-derived phyllosilicate minerals as a source of information about near-surface paleotemperatures requires that (1) the original isotopic compositions have been preserved, and (2) the phyllosilicate is authigenic and formed in chemical equilibrium with waters that have stable oxygen and hydrogen isotope compositions that reside upon the GMWL (Craig, 1961). The unrealisticallywarm phyllosilicate crystallization temperature estimates (44 to 55 ± 3 °C) from the MAC and ELY localities (Table 3; Figs. 1 and 6) in conjunction with their illite-rich I/S and discrete illite mineralogy (Fig. 4), suggest that the original pedogenic suites of mineralogic and isotopic compositions of the b0.2 μm phyllosilicate minerals in these paleosols have not been preserved and therefore environmental estimates made from them (T °C and δ18Ometeoric water; Table 3) cannot be taken seriously. Previous studies have recognized the importance of diagenesis in the formation of clay mineral assemblages, sediments (Bethke, 1985; Gharrabi and Velde, 1995; Grathoff et al., 2001) and coal maturity (Damberger, 1971; Zimmerman, 1986; Cluff and Byrnes, 1991) in the Illinois basin, and attributed this diagenetic alteration to burial by ~1.5 to 3 km of southward-thickening, post-Pennsylvanian strata in conjunction with middle Permian (270–280 Ma) elevated heat flow associated with magmatic intrusions in the southern part of the basin. As such, the mineralogical and stable isotope composition of the b0.2 μm phylloslicate samples in the MAC and ELY cores are likely altered by diagenesis, do not accurately reflect near-surface ancient Climate sensitive indicators

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Temperature °C Fig. 6. Calculated phyllosilicate crystallization temperatures in equilibrium with meteoric water from the LSC (open black circles), MAC (open black squares), and ELY (open black diamonds) cores plotted relative to their stratigraphic position within the LSC core. Data points represent the average phyllosilicate crystallization temperature for individual paleosols. The black line passes through the data points that plot within their respective modern surface domain (MSD) or warm earth surface domain (WESD). The open gray circle (upper Shelburn Formation) represents the average crystallization temperature of phyllosilicates that did not plot in either their MSD or WESD. The horizontal dashed line denotes the position of the Desmoinesian–Missourian boundary. Also shown is a summary of low-latitude paleotropical climate variability based on the distribution of climate sensitive indicators, including fossil flora (Phillips et al., 1974; Pfefferkorn et al., 2008) coal quality (Cecil et al., 2003), and paleosol morphology and mineralogy (Cecil et al., 1985; Rosenau et al., 2013a,b).

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weathering environments, and therefore are not appropriate for paleotemperature estimates. While samples from the MAC and ELY cores are considered to have undergone diagenetic alteration, several observations suggest Pennsylvanian phyllosilicates in the shallowly-buried LSC core may provide reliable estimates of soil crystallization temperatures. First, the clay mineral assemblages preserved in LSC samples are consistent with a pedogenic origin (Rosenau et al., 2013a). Second, as noted previously, numerous studies suggest the northern part of the Illinois basin experienced a relatively mild burial and diagenetic history (Damberger, 1974, 1991; Cobb, 1981; Barrows and Cluff, 1984; Coveney et al., 1987). Finally, the isotopic trends in the LSC core indicate cooler temperatures (~23 ± 3 °C) at greater stratigraphic depth and significantly warmer temperatures (~31 ± 3 °C) at shallower depths (Fig. 6; Table 3). This stratigraphic trend within the LSC core is inconsistent with a diagenetic pattern affected by burial depth (Stern et al., 1997). 5.6. Meteoric waters and phyllosilicate surface domain arrays Key factors that affect the δ18O and δD values of naturally occurring phyllosilicates are the oxygen and hydrogen-isotope composition of the water from which the phyllosilicates precipitated and crystallization temperature. Considering this, the global IAEA database of meteoric water δD and δ18O values and mean annual surface temperatures in conjunction with hydrogen and oxygen isotope fractionation equations for a particular mineral may be used to produce an array of possible mineral δ18O and δD values that would be expected at modern surface and warm earth surface temperatures (Fig. 7). This approach assumes that the global meteoric water line (Craig, 1961) accurately represents the relationship between the δD and δ18O values of meteoric waters through geologic history. Such arrays of hypothetical δ18O and δD values define a modern surface domain (MSD) and warm earth surface domain (WESD), respectively, for a given mineral (Yapp, 1993, 2000). For the MSD scenario, only mean annual temperatures in excess of 0 °C are used to develop the array of oxygen and hydrogen isotope values, as it is anticipated that liquid water is a requirement for the crystallization of hydroxylated minerals. For a mean annual temperature range of 0 to ~30 °C, this treatment of the IAEA database results in 184 data points (Rozanski et al., 1993; Yapp, 2000; Tabor and Montañez, 2005). For the WESD case, a temperature range of 0 to ~35 °C is chosen and it is assumed that average δ18O value of the oceans may have been ~1‰ more negative than the modern oceans as a result of deglaciation (Savin, 1977). A detailed explanation of the MSD and WESD model is presented in Yapp (2000). It is important to note, however, that an important assumption in the calculation of the WESD is that the equation of the modern GMWL (δD = 8 * δ18O + 10; Craig, 1961) is applicable to the hypothetical WESD. In general, the δD and δ18O values of the modern ice sheets are approximately coincident with the GMWL (Taylor, 1974). Furthermore, the negative shift of oceanic δD and δ18O values, which would accompany the melting of the ice sheets, would plot along a mixing line with a slope of approximately 7.7, nearly parallel to the meteoric water line (Yapp, 2000). Therefore, considering that there are no significant variations in the isotopic fractionation associated with nonequilibrium evaporation from the oceans, small decreases (~1‰) of ocean δ18O values resulting from the melting of the ice sheets would likely have a minimal effect on the intercept of the GMWL (Yapp, 2000). Yapp (1993, 2000), Savin and Hsieh (1998), and Tabor and Montañez (2005) utilized this approach to define surface domain arrays for goethite and phyllosilicate minerals. While it is not necessary that phyllosilicate δD and δ18O values plot within their respective MSD or WESD to demonstrate or negate its formation under modern surface or warm earth surface conditions, the concept, nevertheless, provides a means by which to assess the likelihood that oxygen and hydrogen isotope composition of a particular hydoxylated mineral preserves information about ancient climate (Yapp, 1993, 2000; Savin and Hsieh, 1998; Tabor and Montañez, 2005). To further assess the possibility that

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phyllosilicate δ18O and δD values from the LSC core preserve information about temperatures in the low-latitude Pennsylvanian weathering environment, we adopt the approach used for the assemblage of mineral surface domain arrays (e.g., Yapp, 1993, 2000; Tabor and Montañez, 2005). Following this approach, seven of the sixteen phyllosilicate samples from the LSC core plot within their respective MSD, seven samples plot within the respective WESD, and two samples, from the same paleosol, plot outside both their respective MSD and WESD (Fig. 7; Table 3). Those samples plotting within their respective MSD may preserve isotopic compositions consistent with formation at modern earth-surface conditions (e.g., Savin and Hsieh, 1998). Alternatively in the WESD model, it is assumed that global temperatures are 5 °C warmer and that the Earth is ice-free (Yapp, 2000). The WESD is more inclusive of higher phyllosilicate crystallization temperatures but still may be consistent with phyllosilicate crystallization at earth-surface temperatures, given that mean annual soil temperatures are typically ~2 °C warmer than the mean annual air temperatures (Chang, 1958). Following the organization of phyllosilicate δ18O and δD values within the surface domain model, the two upper Desmoinesian (Shelburn Formation) samples from the same upper Desmoinesian paleosol that plot outside either of their respective surface domains (Fig. 7; Table 3; LSC-21N and LSC-24N), are not considered further, as it is possible that these samples do not preserve information about the low temperature, supergene weathering environment. Nevertheless, it should be noted, that given the uncertainties in δ18O and δD values, as well as the established offset between average annual air temperature and mean annual soil temperature, these samples may preserve a record of near-surface climate conditions. It is difficult to unequivocally document whether ancient phyllosilicate minerals preserve their original isotopic composition. However, considering the aforementioned lines of evidence, it is proposed that, with the exception of one upper Desmoinesian paleosol, phyllosilicates preserved in paleosols in the shallowlyburied LSC core from the northern part of the IB have retained their original oxygen and hydrogen isotope composition, and therefore may provide meaningful information about Pennsylvanian weathering environments. The following discussion considers the paleoclimatic implications of the stable isotope composition, mineralogic composition, and mineral-crystallization temperatures of phyllosilicates from the LSC core that plot within their respective MSD or WESD (Table 3). 5.7. Pennsylvanian phyllosilicate crystallization temperatures and soil water δ18O values (LSC core) Some of the paleotemperature estimates from the LSC core may seem unrealistic, considering that several paleotemperature estimates are higher than those typically observed at modern tropical latitudes (e.g., Chang, 1958). However, as mentioned, mean soil temperatures in modern low-latitude soil-forming environments are typically 1–2 °C warmer than mean annual air temperatures (Barron and Moore, 1994; Barron and Fawcett, 1995; Buol et al., 2003) and due to variations in local canopy cover and soil moisture, it is possible for average annual soil temperatures to be up to 4 °C warmer than average annual air temperatures (Tang et al., 2005; Passey et al., 2010). Furthermore, the specific timing of phyllosilicate crystallization and its relation to seasonally-biased temperature estimates is unknown (Passey et al., 2010; Quade et al., 2013). Nevertheless, the variations in temperature estimates should reflect changes in the environment in which these phyllosilicates crystallized. In this regard, LSC crystallization temperatures provided herein are best considered as maximum estimates, or slight overestimates (~ 2–4 °C), of mean annual surface air temperatures. Numerous studies have recognized significant paleoenvironmental changes near the Middle–Late Pennsylvanian (~ Desmoinesian–

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Fig. 7. δD vs. δ18O plots showing the hypothetical oxygen and hydrogen isotope composition of phyllosilicate samples LSC-62N, LSC-24N, and LSC-15N (open circles) that were calculated using their respective mixed phyllosilicate hydrogen and oxygen isotope fractionation equations (Table 3) in conjunction with mean annual δ18O, δD, and temperature information in the IAEA database (Rozanski et al., 1993). The limits of the Modern Surface Domain (MSD) and corresponding Warm Earth Surface Domain (WESD) for each sample are defined by the solid polygons and dashed polygons, respectively. The solid black squares in each pane represent the measured phyllosilicate δ18O and δD values. (A) Plot of the LSC-62N MSD and WESD along with the measured phyllosilicate δD and δ18O value of LSC-62N showing that the measured isotopic composition is located within the MSD. (B) Plot of the LSC-15N MSD and WESD along with the measured phyllosilicate δD and δ18O value of LSC-15N showing that the measured isotopic composition is located within the WESD. (C) Plot of the LSC-24N MSD and WESD along with the measured phyllosilicate δD and δ18O value of LSC-24N showing that the measured isotopic composition is located outside both the MSD and WESD. See text for discussion.

Missourian) boundary. Particularly well-documented are the profound changes in tropical wetland vegetation that occur during the Middle–Late Pennsylvanian transition, where there is a rapid change from wetland ecosystems dominated by arborescent lycopsids, seed ferns and tree ferns, to ecosystems dominated only by tree ferns

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5.8. Similarities to the Pennsylvanian–Permian transition Interestingly, the ~9 °C temperature increase in phyllosilicate crystallization temperatures from the Desmoinesian to the Missourian

is comparable to the increase in low-latitude pedogenic mineral crystallization temperatures observed across the Pennsylvanian–Permian boundary (Tabor and Montañez, 2005; Tabor, 2007). Akin to the Desmoinesian– Missourian interval, significant changes in tropical vegetation occur near the Pennsylvanian–Permian boundary, wherein there is an overall shift from a humid, seasonally-dry wetland dominated ecosystem in the Late Pennsylvanian to a semi-arid, seasonally-dry dominated ecosystem in the Early Permian (DiMichele et al., 2006, 2001, 2004). These observations, in conjunction with many others (e.g., Mack and James, 1986; Patzkowsy et al., 1991; Kessler et al., 2001; Gibbs et al., 2002; Tabor and Montañez, 2004; DiMichele et al., 2006; Schneider et al., 2006; Montañez et al., 2007; Tabor et al., 2008; Gulbranson et al., 2011) reveal the Pennsylvanian–Permian transition was a period of tremendous lowlatitude environmental change marked by a significant drying of the equatorial Early Permian landscape (Mack and James, 1986; Patzkowsy et al., 1991; Kessler et al., 2001; Gibbs et al., 2002; Tabor and Montañez, 2004; DiMichele et al., 2006; Schneider et al., 2006; Montañez et al., 2007; Tabor et al., 2008). The factors responsible for paleoenvironmental changes across both the Pennsylvanian–Permian (Tabor and Poulsen, 2008, and references therein) and Desmoinesian–Missourian remain unresolved. However, the similar responses observed at low-latitudes across both intervals of apparent warming (i.e., equatorial drying and increased phyllosilicate crystallization temperature) suggest that the commonly-proposed mechanisms responsible for environmental changes across the Pennsylvanian–Permian, such as increasing atmospheric pCO2 (Ekart et al., 1999; Horton et al., 2007; Montañez et al., 2007) and/or Gondwanan deglaciation (Ziegler et al., 1987; Perlmutter and Matthews, 1989; Cecil, 1990; Miller and West, 1993) may also have driven tropical environmental and ecosystem changes across the Desmoinesian–Missourian boundary.

Climate sensitive indicators

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(Peppers, 1964, 1996, 1997; Phillips et al., 1974; Pfefferkorn and Thomson, 1982; Phillips and Peppers, 1984; Phillips et al., 1985a,b; Kosanke and Cecil, 1996; Cleal and Thomas, 2005). This significant change in the composition of wetland vegetation has been attributed to a shift in low-latitude Pennsylvanian climate towards warmer and drier conditions in the Late Pennsylvanian (Missourian; Phillips, 1979; DiMichele and Phillips, 1996; Blake et al., 1999; Falcon-Lang and DiMichele, 2010; Sahney et al., 2010). Numerous other studies of Carboniferous strata support the aridification trend (Cecil et al., 1985; Cecil, 1990; Cecil et al., 2003; Bertier et al., 2008; DiMichele et al., 2009; Bishop et al., 2010; Gulbranson et al., 2010; Rosenau et al., 2013a,b). However, it is not clear whether the aridification was the result of a particularly intense glacial interval on Gondwanaland (Heckel, 1991; Frakes et al., 1992; Falcon-Lang et al., 2011) or, alternatively, an interval of greenhouse warming (Bertier et al., 2008; Bishop et al., 2010; Gulbranson et al., 2010) and attendant retreat of Gondwanan ice sheets. It is during this interval that perhaps the greatest inconsistencies in estimates of ice volume exist between near-field high-latitude records in Gondwanaland and far-field, lowlatitude (cyclothem) records in Euramerica. Estimated phyllosilicate crystallization temperatures from the LSC core reveal significant variations in paleotropical Middle–Late Pennsylvanian soil temperatures. This is significant, because with the exception of a few notable studies (Soreghan et al., 2007a,b, 2008; Tabor et al., in press), it has generally been presumed that Pennsylvanian paleoequatorial climate remained warm throughout the waxing and waning of Gondwana ice sheets (Schopf, 1975; Joachimski et al., 2006; Angiolini et al., 2007; Came et al., 2007). Considering only those phyllosilicates that plot within their respective MSD or WESD, there is a long-term temporal trend towards higher average crystallization temperatures from an average temperature of 23 ± 3 °C in the lower Desmoinesian (uppermost Tradewater–lower Carbondale formations) to an average temperature of 32 ± 3 °C in the Missourian (Patoka–Bond formations; Fig. 6). This temperature increase is coincident with a change from phyllosilicate δ18O and δD values that plot within their respective MSD, to those that plot within their respective WESD (Table 3). Across this same interval, the calculated δ18O values of soil water from which the phyllosilicates crystallized display a + 1.7‰ shift from an average value of − 3.8 ± 0.2‰ in the lower Desmoinesian to an average value of −2.1 ± 0.2‰ in the Missourian (Fig. 8). This positive shift in soil water δ18O values could be attributed to increased soil temperature, greater soil water evaporation, or a combination of both of these factors. Concurrently, there is a shift from a kaolinite-rich b0.2 μm and b2 μm phyllosilicate mineral assemblage in the Desmoinesian (Tradewater– Carbondale formations) to a kaolinite-poor, I/S-rich b 0.2 μm and b2 μm phyllosilicate mineralogy in the Missourian (Rosenau et al., 2013a; Fig. 4; Table 1). This mineralogic trend may reflect a climatic shift from relatively humid conditions in the Desmoinesian to more seasonal precipitation patterns in the Missourian (Rosenau et al., 2013b). Collectively, these trends in phyllosilicate crystallization temperatures, clay mineralogy, and soil water δ18O values in conjunction with a shift from phyllosilicates that plot within their respective MSD in the Desmoinesian to those that plot within their respective WESD in the Missourian, reveal a dynamic Middle–Late Pennsylvanian tropical climate during an icehouse world (LPIA), characterized by a shift in low-latitude near-surface temperatures towards warmer, and possibly drier, conditions from the Desmoinesian into the Missourian. Therefore, these observations lend support to the hypothesis that aridification near the Desmoinesian–Missourian boundary was coincident with a drawdown, rather than a build-up, of Gondwanan ice volume (Bertier et al., 2008; Bishop et al., 2010; Gulbranson et al., 2010).

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δ18Owater(‰V-SMOW) Fig. 8. Average calculated δ18O values of meteoric water in equilibrium with phyllosilicates from the LSC core (open black circles). The black line passes through the calculated meteoric water δ18O values in equilibrium with phyllosilicates that plot within their respective modern surface domain (MSD) or warm earth surface domain (WESD). The open gray circle (upper Shelburn Formation) represents the calculated δ18O value of meteoric water in equilibrium with phyllosilicates that did not plot in either their MSD or WESD. The horizontal dashed line denotes the position of the Desmoinesian–Missourian boundary. See text for discussion.

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6. Summary The b 0.2 μm phyllosilicate fractions of Middle–Upper Pennsylvanian paleosols sampled from three different localities within the Illinois basin are composed of mineralogic mixtures of interstratified illite–smectite and kaolinite, and rarely, illite. Thermodynamic data in conjunction with the chemical composition and mineralogy of the phyllosilicate mixtures were used to calculate unique oxygen and hydrogen isotope fractionation factors for each sample. Assuming the phyllosilicates preserve a record of isotopic equilibrium with Pennsylvanian meteoric waters, the measured oxygen and hydrogen isotope values of these samples correspond to crystallization temperatures that range from 22 ± 3 °C to 55 ± 3 °C. Average phyllosilicate crystallization temperatures from a shallowly-buried locality in the northern part of the basin (LSC) are 29 ± 3 °C, and are considered consistent with formation in the low-temperature Pennsylvanian weathering environment. Average crystallization temperatures from two deeply-buried localities in the interior (MAC) and southern part (ELY) of the Illinois basin are 46 ± 3 °C and 53 ± 3 C, respectively. Phyllosilicate temperature estimates from the MAC and ELY cores are inconsistent with a pedogenic origin and instead indicate post-pedogenic, diagenetic alteration of the original isotopic composition of the phyllosilicates. A similar spatial trend in clay mineralogy is observed across the basin, in which I/S becomes more illite-rich from the margin to the interior of the basin. These isotopic and mineralogic trends are considered to be the result of diagenetic recrystallization of pedogenic minerals in more deeply-buried parts of the basin (MAC and ELY) in response to burial by approximately 1.5 to 3 km of southward-thickening, post-Pennsylvanian strata, and an interval of middle Permian elevated heat flow associated with magmatic activity in the southern part of the basin. Phyllosilicate crystallization temperatures from the shallowlyburied LSC core reveal a long-term warming from an average value of 23 ± 3 °C in the Desmoinesian to an average value of 32 ± 3 °C in the Missourian. With the exception of one uppermost Desmoinesian paleosol, the phyllosilicate δ18O and δD values plot within their respective Modern Surface or Warm Earth Surface Domain arrays, suggesting that these phyllosilicates preserve a record that is consistent with low-latitude Pennsylvanian Earth-surface conditions. These significant variations in low-latitude Middle–Late Pennsylvanian temperatures are unparalleled on modern Earth and reveal a dynamic Late Paleozoic paleoequatorial icehouse climate. Moreover, the increase in phyllosilicate crystallization temperatures from the Desmoinesian to the Missourian is similar in magnitude to pedogenic mineral crystallization temperatures observed across the Pennsylvanian–Permian boundary in western equatorial Pangaea, suggesting that similar mechanisms may have been responsible for the accompanying paleoenvironmental changes across both intervals. The oxygen-isotope composition of meteoric water from which the LSC phyllosilicates crystallized ranges from − 4.1 ± 0.2‰ to − 1.9 ± 0.2‰ and displays a + 1.7‰ shift from an average value of − 3.8‰ in the lower Desmoinesian to an average value − 2.1‰ in the Missourian. This trend is consistent with a long-term warming of the low-latitude Pangaean tropics. The complex history of pedogenesis and diagenetic alteration in the Illinois basin brings to light the need for cautious interpretation of the stable isotope composition of phyllosilicates from paleosols, but ultimately indicates that reliable paleoenvironmental information may be retained in paleosol phyllosilicates that have experienced a shallow burial history and minimal post-depositional heating. Acknowledgments The authors would like to thank two anonymous reviewers for their constructive comments that greatly improved the quality of this manuscript. We thank Dr. Ray Guillemette of Department of the Department of Geology and Geophysics at Texas A&M University for assistance in the

collection of electron microprobe data. We also thank Phil Ames of Peabody Energy, Barry Sargeant of Knight Hawk Coal, Vigo Coal Operating Company, and Triad Mining, Inc. for providing access to their mines and permitting collection of samples, as well as the Illinois State Geological Survey for providing access to cores, core descriptions, and core samples. N. Rosenau was supported by a National Science Foundation Graduate Research fellowship, a Roy M. Huffington fellowship provided by the Roy M. Huffington Department of Earth Sciences (SMU), and grants-in-aid from the Geological Society of America and Clay Minerals Society. N. Tabor was supported by NSF-EAR 0545654 and NSF-EAR 0844147. References Angiolini, L., Gaetani, M., Muttoni, G., Stephenson, M.H., Zanchi, A., 2007. Tethyan oceanic currents and climate gradients 300 m.y. ago. Geology 35, 1071–1074. Barnhisel, R.I., 1977. Chlorites and hydroxyl interlayered vermiculite and smectite. In: Dixon, J.B., Weed, S.B. (Eds.), Minerals in Soil Environments. Soil Science Society of America, Madison, Wisconsin, pp. 331–356. Barrick, J.E., Lambert, L.L., Heckel, P.H., Boardman, D.R., 2004. Pennsylvanian conodont zonation for Midcontinent North America. Rev. Esp. Micropaleontol. 36, 231–250. Barron, E.J., Fawcett, P.J., 1995. The climate of Pangea: a review of climate model simulations of the Permian. In: Scholle, P.A., Peryt, T.M., Ulmer-Scholle, D.S. (Eds.), The Permian of Northern Pangaea. Paleogeography, Paleoclimates, Stratigraphy, vol. 1. Springer, New York, pp. 37–52. Barron, E.J., Moore, G.T., 1994. Climate Model Application in Paleoenvironmental Analysis, 33. Society for Sedimentary Geology (SEPM) Short Course. Barrows, M.H., Cluff, R.M., 1984. New Albany Shale Group, (Devonian–Mississippian) source rocks and hydrocarbon generation in the Illinois basin. In: Demaison, G., Murris, R.J. (Eds.), Petroleum Geochemistry and Basin Evaluation. 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