The Pennsylvania State University. The Graduate School. College of Earth and Mineral Sciences THE EFFECT OF TEMPERATURE AND PRECIPITATION ON SODIUM

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1 The Pennsylvania State University The Graduate School College of Earth and Mineral Sciences THE EFFECT OF TEMPERATURE AND PRECIPITATION ON SODIUM DEPLETION FRONTS IN SOILS DEVELOPED ON PEORIA LOESS A Thesis in Geosciences by Jennifer Zan Williams 2008 Jennifer Zan Williams Submitted in Partial Fulfillment of the Requirements for the Degree of Master of Science August 2008

2 The thesis of Jennifer Zan Williams was reviewed and approved* by the following: Susan L. Brantley Professor of Geosciences Director, Earth and Environmental Systems Institute Thesis Advisor Eric Kirby Associate Professor of Geosciences Lee R. Kump Professor of Geosciences David Pollard Senior Scientist, Earth and Environmental Systems Institute Katherine H. Freeman Professor of Geosciences Head of the Department of Geosciences Graduate Program *Signatures are on file in the Graduate School ii

3 ABSTRACT A north-south transect along the Mississippi River valley provides an opportune environmental gradient across which to investigate chemical weathering. Soil profiles along this transect are interpreted to have developed from a uniform parent material, the Peoria Loess, with pedogenesis commencing between C ka. We examined mineral evolution in these soils using X-ray fluorescence (XRF) and X-ray powder diffraction (XRD) mineralogical analysis. Results indicated quartz, feldspars, and clays dominate the loess and soils. In each pedon the Na concentrations generally decrease from the deepest collected sample to the surface, defining depletion profiles. These depletion profiles document the chemical weathering of plagioclase. Consistent with this interpretation, kaolinite and illite concentrations are greatest in the surface horizons and decrease with depth. In addition, the montmorillonite concentrations are observed greatest with depth, consistent with primary mineralogy inherited from the source region. In order to interpret the concentration versus depth profiles as a function of climate variation along this transect, we simulated climate at 10 ka and 6 ka, as well as for modern day using the GENESIS v2 Global Climate Model (GCM). The GCM calculated annual averages for surface air temperature, precipitation, and downward flux of soil pore water. Uniform soil texture was assumed to be constant throughout this transect from 28º to 50º N latitude and 94º to 88º W longitude along the Mississippi River valley. The Na profiles were fit to mathematical model equations for profile development using the GCM outputs to determine values of a kinetic parameter for each pedon. The fraction of albite dissolved (f) varied from approximately 0% in the north to 26% in the south as determined through integration of the model equations. Variation in f iii

4 is observed to largely depend upon the variation in annual rainfall from north (0.5 m / y) to south (1.2 m / y). The kinetic parameters were corrected for porefluid advection along the transect and plotted against 1/T. From these plots, the best estimate of the apparent activation energy for Na-plagioclase dissolution is observed to be 95 ± 24 kj/mol. Such quantitative interpretations of soil profiles can provide ways to understand and predict the effects of climate on soil chemistry. iv

5 Table of Contents List of Figures...vi List of Tables...vii Acknowledgements..viii Introduction..1 Soil Pedons..3 Mineralogical Analysis 7 Soil Profiles 10 Mass balance calculations..10 Characterization of profiles 14 Results 20 Mineralogical analysis...20 Bulk density...33 Fitting the concentration gradient..35 Kinetic parameter...40 Porefluid advection velocity..43 Activation energies of dissolution. 45 Extent of weathering..50 Discussion..55 Conclusions 61 Bibliography..62 Appendix A...67 Appendix B...80 Appendix C...90 Appendix D...98 Appendix E.111 v

6 List of Figures Figure 1. Map of central United States..4 Figure 2. Plots of τ i,j for Pedon 13 versus depth...12 Figure 3. Plots of τ Zr,j for j =Ca, Fe, and Mn for Pedon 8 assuming different compositions for parent 13 Figure 4. Graphical representations of the different reaction front types observed in the soils along this Mississippi valley transect 16 Figure 5. Diagrams for Na concentrations normalized to Zr (τ Zr,Na ) plotted versus depth 17 Figure 6a. X-ray diffractograms for Pedon 1 (northeast Iowa)...21 Figure 6b. X-ray diffractograms for Pedon 13 (Tennessee) 22 Figure 6c. X-ray diffractograms for Pedon 22 (Louisiana).23 Figure 7. X-ray diffractogram for clay separate from Pedon Figure 8. Normalized Na concentrations plotted versus depth for four pedons..39 Figure 9. Values for K plotted versus corresponding pedon latitude..41 Figure 10. GENESIS v2 Global Climate Model quantities for the Mississippi River transect..46 Figure 11. GCM calculated porefluid advection velocities for modern day plotted versus depth..47 Figure 12. Arrhenius plots depicting the variation in ln Kv versus 1/T..50 Figure 13. Arrhenius plot depicting the weighted means of GCM model output plotted as ln Kv versus 1/T 51 Figure 14. Percent Na depletion plotted versus latitude for soils along transect 53 Figure 15. Calculated depletion trends plotted versus latitude...54 vi

7 List of Tables Table 1. Published activation energies for albite dissolution 2 Table 2. Summary of Pedons 5 Table 3. Characterization of soil profiles along Mississippi transect..19 Table 4. Mineral presence or absence based on XRD.25 Table 5. Volumetric mineral abundance from XRD...26 Table 6. Calculated volumetric mineral abundance from normative mineral model for Pedon 1 31 Table 7. Statistical analyses of model equations 37 Table 8. GENESIS v2 Global Climate Model output for Mississippi River transect 48 Table 9. Calculated parameters which describe the lumped kinetic parameter..58 vii

8 Acknowledgements I wish to thank my advisor, Sue Brantley, for her guidance in my scientific growth. I would like to thank my coauthors, J. Bandstra and D. Pollard for their modeling expertise and contributions to the project. I would also like to thank Daniel Muhs (USGS; Denver, CO) for providing soil samples and insightful dialogues, Alex Blum (USGS; Boulder, CO) for providing analyses and interpretation and Art White (USGS; Menlo Park, CA) for original motivation. I would like to thank the Brantley group for their intellectual and emotional encouragement. Additionally, I would like to thank all of the people here at Penn State who have contributed to this research or the learning process: H. Albrecht, M. Angelone, H. Buss, B. Fambrough, T. Fischer, P. Fulton, H. Gong, D. Greene, E. Hausrath, B. Kimball, L. Liermann, K. Morell, J. Moore, A. Navarre-Sitchler, A. Wall, and N. Wonderling. Funding was provided by the following sources: Bunton-Waller Graduate Award, Penn State Biogeochemical Research Initiative for Education (BRIE) sponsored by NSF (IGERT) grant DGE , Charles E. Knopf, Sr. Memorial Scholarship and by the National Science Foundation under Grant No. CHE (CEKA). And finally, I am most grateful to my parents, James P. and Mary R. Williams, who have always encouraged me to pursue my interests and to enjoy the journey. viii

9 Introduction Mineral dissolution studies have often focused on feldspars since they are the most common minerals in crustal rocks (Blum and Stillings, 1995). It is well documented that mineral dissolution rates are temperature-dependent when observed in the laboratory (Helgeson et al., 1984; Chou, 1985; Knauss and Wolery, 1986; Rose, 1991; Chen, 1994; Hellman, 1994; Stillings et al., 1995; Blum and Stillings, 1995; Chen and Brantley, 1997; Brantley, 2008). However, laboratory dissolution rates are consistently measured at 2-5 orders of magnitude faster than field measured rates for the same minerals (White and Brantley, 2003). This laboratory to field rate discrepancy and the inability to extrapolate rates of dissolution from one system to another has been attributed to such differences as the rate-controlling mechanisms between laboratory and field, problems in assessing surface area, and the complex hydrology of natural systems (Kump et al., 2000; White, 2008). Although this rate discrepancy exists, a direct comparison between laboratory and field measurements can be accomplished by measuring energies of activation for mineral dissolution. For example, a compilation of granitoid watershed weathering data for small, monolithologic watersheds worldwide documents that the activation energy for Na plagioclase dissolution inferred from watershed measurements is 78 kj/mol (White and Blum, 1995). This value is comparable to laboratory measurements of 60 kj/mol ± 10 (Blum and Stillings, 1995). Table 1 summarizes all literature reported activation energies for both laboratory and field studies of Na-plagioclase dissolution to date. 1

10 Table 1. Published Activation Energies for Albite Dissolution Measurement E app (kj/mol) Reference Laboratory Temperature Range º C 84 Helgeson et al., º C 59 Chou and Wollast, º C 68 Knauss and Wolery, 1986 nr 58 Sverdrup, º C 71 Rose, º C 63 Chen, º C 81 Hellman, º C 44 Stillings et al., º C 60 Blum and Stillings, º, 50º, and 90º C 65 Chen and Brantley, 1997 Average 65 ± 12 Field Location metamorphic watershed 77 Velbel, 1993 Hawaiian basalt flow 109 Dorn and Brady, 1995 granitoid watersheds 78 White and Blum, 1995 granitoid watershed 80 White et al., 1999 Average 86 ± ± 24 This study 2

11 Here, we analyze a published geochemical data set for soils developed from Peoria loess spanning 13º of latitude to infer the influence of temperature and precipitation on chemical weathering (Muhs et al., 2001). We examine the application of proposed model equations used to describe reaction fronts developed during weathering (Murphy et al., 1998; Lichtner, 1988; Brantley et al., 2008). These fronts are defined by concentration gradients that document weathering rates at the pedon scale. Further, we use a global climate model (GCM) to simulate temperature, precipitation, and porefluid velocities during the last 10 ky in order to calculate rates of plagioclase dissolution along this transect as a function of these climate variables. Soil Pedons Extent of weathering was investigated by analyzing published elemental data as a function of depth for 22 pedons developed on deep Peoria Loess (Muhs et al., 2001; see Appendix A). These 22 pedons (Figure 1), classified as very deep, well-drained, noneroded, upland soils, are members of the Fayette, Alford, or Memphis series (Ruhe, 1984a; Muhs et al., 2001; Soil Survey Staff, 2004). This sequence of soils represents a climosequence: a climosequence is defined as a set of soils that developed on relatively identical parent material and was weathered for approximately identical duration under different regimes of precipitation and temperature (White, 1995). Summarized in Table 2 are the pedon numbers, location, soil series, and mean annual precipitation for each of these sites (Muhs et al., 2001). Based upon radiocarbon dating, Peoria loess was deposited between ~23 10 ka and pedogenesis commenced between ka (Muhs et al., 2001). The classic model 3

12 Figure 1. Map of central United States. Stars correspond to pedon site locations where the modern soils developed on Peoria Loess were sampled from well-drained, non-eroded uplands. Longitude and latitude coordinates for each pedon from Muhs et al. (2001) are summarized in Table 2. 4

13 Table 2. Summary of Pedons 1 Mean Annual Location Soil Precipitation Pedon City & State Latitude Longitude Series (mm/yr) 1 Clayton County, IL 42º59.00' 91º12.66' Fayette Galena, IL 42º26.22' 90º30.46' Fayette Mt. Carroll, IL 42º11.11' 89º59.72' Alford Morrison, IL 41º49.02' 89º87.94' Alford Rapids City, IL 41º34.02' 90º20.85' Alford Mammoth, IL 40º57.84' 90º39.02' Alford Quincy, IL 39º57.52' 91º21.11' Fayette Greenbay Hollow, IL 38º59.02' 90º36.42' Fayette Ellis Grove, IL 38º00.79' 89º54.48' Fayette Murphysboro, IL 37º46.44' 89º21.52' Fayette Anna, IL 37º22.50' 89º07.50' Alford Arlington, KY 36º46.93' 89º02.57' Memphis Lauderdale County, TN 35º51.26' 89º33.05' Memphis Millington, TN 35º25.73' 89º48.37' Memphis Senatobia, MS 34º37.22' 90º04.03' Memphis Batesville, MS 34º14.31' 89º56.29' Memphis Greenwood, MS 33º30.14' 89º57.12' Memphis Yazoo County, MS 32º46.89' 90º22.38' Memphis Mechanicsville, MS 32º38.88' 90º31.54' Memphis Vicksburg, MS 32º24.46' 90º49.30' Memphis Natchez, MS 31º34.27' 91º18.97' Memphis St. Francisville, LA 30º47.27' 91º22.44' Memphis All data previously published by Muhs et al. (2001). 5

14 for formation of glacial loess such as the Peoria is described by a two-part depositional history. First, silt-sized particles are produced by the Laurentide ice sheet and carried in meltwaters to the floodplain of the Mississippi and Missouri Rivers in North America (Chamberlin, 1897; Leighton and Willman, 1950; Lugn, 1965; Norton et al., 1988; Bettis et al., 2003; Mahowald et al., 2006). Subsequently, the dried fines were picked up by northwesterly and westerly winds (Pye and Johnson, 1988; Muhs and Bettis, 2000) and redeposited on the adjacent uplands (Chamberlin, 1897; Leighton and Willman, 1950). The thickest loess deposits are proximal and east of the Mississippi River, thinning and fining with increasing distance from the river source region (Chamberlin, 1897; Udden, 1898; Leighton and Willman, 1950; Fehrenbacher et al., 1965a; Norton et al., 1988). Samples along this north-south transect were collected at each pedon from the surface to the C horizon by Muhs et al. (2001). This C horizon was identified as the depth at which prismatic, columnar, or subangular blocky structure was no longer evident in the profile, and was interpreted to represent unaltered parent material (Muhs et al., 2001). Bulk mineralogy for the parent loess (and modern soils) was reported previously on the basis of observations using XRD (Ruhe, 1984a; Pye and Johnson, 1988). Pye and Johnson (1988) concluded that the parent (unleached) loess in the south (at depths greater than 2 m) near Vicksburg, MS, and Natchez, MS (corresponding to Pedon 20 and 21, respectively), contains quartz, dolomite, and feldspars with minor amounts mica, chlorite, kaolinite and calcite. In contrast, kaolinite and illite were identified as the predominant surface and shallow subsurface clay phases in soils along this Mississippi River transect. Expandable clays such as montmorillonite and vermiculite comprise the major clay constituents at 6

15 depth (Ruhe, 1984a). The leached soils are therefore composed mainly of quartz, plagioclase, smectite, and mixed-layer illite-smectite with virtually no calcite or dolomite (Pye and Johnson, 1988). Textural measurements have been reported for each of the soil sequences represented along this transect. For the northern pedons, occurring from Iowa to Illinois (pedons 1, 2, 7, 8, 9, and 10), the Fayette Soil series is reported to contain sand/silt/clay ratios, which range from % (sand), % (silt), and % (clay) (Soil Survey Staff, 2004). Similarly, for the Alford Soil series in southern Illinois and southwestern Indiana (pedons 3 through 6 and 11), Fehrenbacher et al. (1965b) reported particle size distributions that range from ( %) / ( %) / ( %); however, bulk density was not reported for either of these soil series. From a specific investigation of grain size distribution of the Memphis Soil sequence, Lindbo et al. (1994) reported ranges in bulk density from 1.38 to 1.86 g cm -3 and sand/silt/clay ratios that range from (0.3 54%) / ( %) / ( %). This Memphis soil sequence (pedons 12 through 22) occurs from Kentucky to Louisiana on terraces and uplands of the Coastal Plain (Soil Survey Staff, 2004). Mineralogical Analysis In work reported here, we largely rely upon published chemical analyses from Muhs et al. (2001). As stated by Muhs et al. (2001), soils from the southern part of this transect were sampled from hand dug pits or deep road cuts, while soils in the northern part of this transect were sampled from cores collected with a hydraulic drilling rig. Total soil splits from each horizon were pulverized in a shatterbox; major oxide 7

16 chemistry and trace element chemistry were determined using wavelength dispersive and energy dispersive x-ray fluorescence, respectively (Muhs et al., 2001). We selected a subset of samples from pedons 1, 13, and 22 (Muhs et al., 2001) for further analysis to compare with published analyses. Bulk samples corresponding to the deepest C horizon, the Bt horizon at ~0.75 m below surface, and the surface A horizon were ground with a zirconium mortar and pestle to pass through 325 mesh (an equivalent particle size of 45 μm). Characterization was accomplished by X-ray powder diffraction (XRD) on zero-background quartz slides. The Scintag PAD V diffractometer was operated at 35 kv voltage and 30 ma current using Cu-Kα1 and Cu-Kα2 radiation and a Ge solid-state detector. Diffraction patterns were collected from a 2º - 70º continuous scan with a speed of 2º per minute 2θ. One more extensive XRD analysis of clays was performed on sample IA305 (the Bt horizon for the northernmost pedon). Following procedures outlined by White and Dixon (2003), organic matter and carbonates were removed first. Next, 150 mg of sample was saturated with 25 ml of 1N MgCl 2 solution, sonicated, and centrifuged for 5 minutes at 1000 rpm. Clear supernatant was removed and the step was repeated twice. Next, 25 ml of distilled water was added to the centrifuge tube, the sample was mixed and centrifuged for 5 minutes at 1000 rpm and decanted as before. This step was repeated three times (until clays began to disperse), and finally centrifuged at 2500 rpm for 10 minutes. The supernatant was poured off, and using a pipette, the sample was placed on a Vycor slide, dispersed and allowed to air dry. Diffraction patterns were collected from a 2º - 20º continuous scan with a speed of 1º per minute 2θ using the same Scintag PAD V diffractometer described above. Each collected diffractogram was 8

17 interpreted using MDI Jade Version +8.5, available from Materials Data Inc., Livermore, CA, USA. Mineral abundance measured in the same subset of soils was accomplished using the computer program, RockJock (Eberl, 2003), where quantitative estimates of mineralogy are calculated from powder XRD data. Samples were prepared and analyzed following procedures by Eberl (2003) at the United States Geological Survey, Boulder, CO by Alex Blum and described as follows. For each analysis, 00 g of sample was mixed with g of ZnO, passed through the McCrone sieve and then ground with 4 ml of methanol in a McCrone micronizing mill for 5 minutes. Next, the mixture was dried at 85º C then ground to pass through 500 μm McCrone sieve. The final mixture was stirred to ensure uniformity. A random XRD mount was prepared by side packing an XRD holder, using tapping on hard surface to ensure randomness is maintained during compaction. Samples were X-rayed from 5º - 65º 2θ using a step size of 2º 2θ and a count time of two seconds per step. Each collected diffractogram was analyzed by RockJock where pattern fitting is used to model the observed XRD pattern with measured patterns of pure mineral standard. The program iteratively varies the fraction of each standard to minimize the degree of fit between the calculated and measured pattern. Integrated intensities for each mineral were then compared to the integrated intensity of the standards and weight percents of the minerals were calculated. 9

18 Soil profiles Mass balance calculations The interpretation of elemental losses or gains within a soil profile to determine extent of weathering is accomplished with mass balance models (Brimhall and Dietrich, 1987; Anderson et al., 2002; White, 2002). In this study, mass changes relative to parent material were determined by calculating the dimensionless element-mass-transfer coefficient, τ i,j (Brimhall and Dietrich, 1987; Anderson et al., 2002): c j, w ci, p τ i, j = 1 (1) c c j, p i, w where c x,y is the concentration (mass element / mass soil) of mobile (x = j) or immobile (x = i) elements in the weathered (y = w) or parent (y = p) material. Values of this parameter can be interpreted as follows: τ i,j < 0 represents relative depletion of element j with respect to element i in the parent composition; τ i,j = 0 represents no change from parent composition; and τ i,j > 0 represents relative enrichment with respect to the parent composition. To calculate τ i,j using equation (1), we must assume a parent composition and an immobile element. To determine an immobile element, we assumed that the deepest sample collected in each pedon was representative of parent composition and we calculated τ i,j for several elements likely to be immobile: Ti, Zr, and Nb (White, 1995). If an assumed element, i, is immobile and if leaching has occurred for most other elements j, then most of the τ i,j values will demonstrate depletion instead of enrichment 10

19 (Figure 2). Relative depletion can then be inferred for each element: for example, as documented in Figure 2, Zr is less mobile than Ti and Al in Pedon 13. Similar plots for other pedons also documented Zr to be the least mobile element; therefore, Zr was used as the immobile element for all calculations here. Assessment of a parent composition is challenging for depositional materials such as loess (i.e. when no bedrock is present). In general, it is assumed that the specimens collected at greatest depth on depositional material are representative of parent (White, 2002). Here, three potential parent compositions were investigated: i) the deepest, least weathered sample of the entire soil transect (see White, 2002), ii) the average of four specimens stratigraphically correlated to basal Peoria Loess (Pye and Johnson, 1988), and iii) the deepest sample of each pedon. Following approach (i), a sample from Pedon 1, the most northerly location (northeastern Iowa), extracted from a depth of 390cm, was tested as parent. For approach (ii), the analyses of four deep Peoria Loess specimens were averaged (Pye and Johnson, 1988). Approach (iii) was considered because the composition of Peoria Loess had been observed to vary along this soil sequence (Fehrenbacher et al., 1965b); therefore, the deepest collected sample of each pedon was tested as parent for the overlying profile. Testing these three parent composition assumptions using equation (1), we observed differences in the apparent behavior of several elements within the soil profiles. As shown in Figure 3 for Pedon 8 (Greenbay Hollow, IL), the results of each test produced characteristic profiles. We analyzed the curves against a conceptual model wherein we expected little to no addition of elements to the soil in the last 13 ky: i.e., we 11

20 A τ Ti, j Depth (cm) Al Zr 250 B 0 τ Zr, j Depth (cm) Al Ti 250 Figure 2. A) Plot of τ Ti, j for pedon 13 (data from Muhs et al., 2001) versus depth. This plot demonstrates that Al is depleted and Zr enriched in the shallow soils relative to Ti. Zr is inferred to be more immobile than either Ti or Al. B) Plot of τ Zr, j for pedon 13 versus depth. Both Ti and Al are depleted in the shallow soils relative to Zr. Pedon 13 is characteristic of most other pedons. Vertical dashed line indicates τ = 0 (soil identical to parent composition). The assumed parent for these plots was the deepest collected specimen of this profile. 12

21 A τ Zr, j Depth (cm) Calcium Iron Manganese B τ Zr, j Depth (cm) Calcium Iron Manganese C τ Zr, j Depth (cm) Calcium Iron Manganese Figure 3. Plots of τ Zr, j for j = Ca, Fe, and Mn for Pedon 8 assuming different compositions for parent. A) The deepest, least weathered specimen (Pedon 1) of this transect was assumed parent. B) The average of four stratigraphically correlated deep Peoria loess samples from Mississippi was assumed parent. C) The deepest sample of the profile was assumed parent. Vertical dashed line represents parent concentration. 13

22 assumed that each soil should show only depletion or total retention of major elements. In addition, we assumed that τ i,j would approach 0 at depth. In contrast, when the deepest, least weathered northern-most sample is assumed as parent following approach (i) (Figure 3a), the concentrations of calcium, iron, and manganese show significant relative enrichment (τ i,j > 0.5) at depth and do not return to parent concentrations observed in the C horizon of the soils. When the averaged, deep Peoria Loess composition is assumed to be parent according to approach (ii) (Figure 3b), the same three cations still show enrichment at depth relative to the parent concentration, only to a lesser extent. Although such enrichment might be possible with significant lateral flow that brings these cations into the system and causes deposition, the upland location of these soil profiles led us to reject these possible choices of parent. However, when the deepest collected sample from each pedon is assumed as parent (iii) (Figure 3c), the concentrations of calcium, iron and manganese demonstrate depletion profiles (Ca, Mn) and depletion-enrichment profiles (Fe) (further discussed below). Therefore, we argue in support of Fehrenbacher et al. (1965b) and Kleiss (1973) that loess composition varies across the climosequence. This variability requires that the best choice of parent is the deepest sample from each pedon along this transect. Characterization of profiles With the assumptions of parent composition and immobile element, the extent of chemical weathering in these loess-derived soils can be calculated. The resultant reaction fronts document the integrated history of mineral reaction kinetics over the last 13 ky. The slope of these reaction fronts can be interpreted with respect to the weathering 14

23 environment: for systems that experienced similar porefluid velocities, a concentrationdepth curve exhibiting a shallow slope is indicative of a faster weathering rate (Murphy et al., 1998; White, 2002). Five characteristic τ i,j -profile types have been identified: immobile, depletion, depletion-enrichment, addition, and biogenic (Brantley et al., 2007). For the immobile profile, elements commonly considered unreactive in the soil environment, such as Zr, Ti, and Nb, exhibit parent concentration throughout the profile (Brimhall and Dietrich, 1987; White, 1995; see Figure 2). In the Peoria soils, only Zr can be considered largely immobile, while Ti and Al demonstrate depletion profiles when Zr is assumed immobile. Depletion profiles, those that depict continually decreasing concentrations from depth to surface, are often observed for elements such as Na (White, 1995; Figure 4a). For many soils, Na depletion fronts are interpreted to reflect the dissolution of plagioclase because it is the principal and most reactive primary mineral containing Na (White, 1995; Muhs et al., 2001). Along this climosequence, the distinct and consistent depletion of Na, as observed in Figure 5 is interpreted to represent the depletion of plagioclase, further explored in this study. Other elements that largely display depletion profiles in the Peoria loess soils are Ca and Mg. For these two elements essentially complete depletion fronts were observed in pedons 2, 3, 5, 8, and 20. The remaining pedons show depletion profiles for Ca and depletion-enrichment profiles (see below) for Mg as shown in Appendix B. A combination of eluviation (the downward movement of soluble or suspended material by groundwater percolation in a soil) and illuviation (the accumulation of soluble or suspended material that was transported by eluviation) (Brady and Weil, 2002) 15

24 A 0 τ Zr, Na C τ Zr, Mn Depth (cm) Depth (cm) Pedon 9 Illinois Pedon 22 Louisiana B τ Zr, Fe D τ Zr, P Depth (cm) Depth (cm) Pedon 9 Illinois Pedon 18 Mississippi Figure 4. Graphical representations of the different reaction front types observed in the soils along this Mississippi valley transect, as determined by τ Zr,j (Equation 1) where the parent is the deepest sample from the specified pedon. A) A depletion profile is characterized by a generally continual decrease in concentration from parent substrate at depth to soil surface. B) A depletionenrichment profile is characterized by surface depletion and subsurface enrichment of an element, such as Fe shown here. C) An addition profile is characterized by the continual increase in concentration with decreasing depth. D) A biogenic profile is characterized by depletion at depth, grading into enrichment at the soil surface due to biotic uptake, as depicted here by P. Vertical dashed line represents parent concentration. 16

25 A 0 τ Zr, Na Depth (cm) P1 P2 P3 P4 P5 P6 P7 P8 P9 P10 P11 Pedons B 0 τ Zr, Na Depth (cm) P12 P13 P14 P15 P16 P17 P18 P19 P20 P21 P22 Pedons Figure 5. Diagrams for Na concentrations (from Muhs et al., 2001) normalized to Zr (τ Zr, Na ) plotted versus depth. A) Northern pedons, from northeastern Iowa to southern Illinois. Pedons 3, 4, 5, 6 and 11 are members of the Alford Soil Series (Muhs et al., 2001), which consists of an anthropogenic plow layer resultant of farming practices (Soil Survey Staff, 2004). Enrichment of Na concentrations depicted to the right of dashed line may be resultant of fertilizer additions or soil amendments (Bettis et al., 2003). B) Southern pedons, from Kentucky to Louisiana. We assumed parent composition was consistent with the deepest sample of each pedon. 17

26 is documented by the depletion-enrichment profile. Often, Al and Fe show such depletion- enrichment profiles (Brantley et al., 2007; Figure 4b). These profiles are characterized by depletion at the surface combined with enrichment at depth. As shown in Appendix B, in addition to Al and Fe, Mg and P also exhibit this type profile in pedons 10, 11, 12, 14, and 22. Wet or dry deposition from the atmosphere produces addition profiles such as those depicted by Mn in the Peoria loess, where the concentration is highest at the surface and decreases to parent concentrations with depth (Brantley et al., 2007; Figure 4c). Addition profiles also characterize elements such as C that are fixed by plants at the soil surface (Brantley et al., 2007). As shown in Appendix B, Mn is the only element to demonstrate the characteristics of an addition profile in a few of these soils. Finally, a biogenic profile results from the extraction of nutrients such as K at depth and enrichment at the surface due to biotic processes (Brantley et al., 2007; Figure 4d). In the Peoria loess soils, P, K, and Fe are elements that exhibit biogenic profiles in a limited number of the pedons, specifically 5, 6, and 18 (Appendix B). All of the elements investigated in this study were classified into profile types based upon the morphology of their concentration-depth curves (Table 3 and Appendix B). Some elements show mixed profile characteristics that cannot be attributed to one profile type. For the Peoria soils, elements that exhibit mixed profiles in some pedons are Mg, K, Fe, P, Mn and Si (Appendix B and Table 3). 18

27 Table 3. Characterization of Soil Profiles from Mississippi Transect Soil Profile Types Element * Immobile Depletion Depletion-Enrichment Addition Biogenic Zr Na, Ca, Mg, Ti, Si, K, Mn Al, Fe, Mg, P, Si Mn P, K, Fe *Note: For elements listed more than once, they are characterized differently in different pedons. 19

28 Results Mineralogical Analysis XRD analyses documented that quartz, albite, potassium feldspar, illite, mica and clays were observed in the nine subsamples studied from the surface, the Bt, and the C horizons of the most northern, most southern, and mid-latitude pedons. Our XRD results (Table 4) confirm general compositional uniformity of the soils along this transect as discussed in previous research (Ruhe, 1984b; Pye and Johnson, 1988; Muhs et al., 2001). Quantitative XRD results (personal communication, A. Blum; Table 5) support our observations. Only one sample, the deepest from the northernmost pedon (Figure 6a) contained dolomite. Calcite was not identified in any of the samples by XRD, consistent with observations by Pye and Johnson (1988) for the loess soils. A clay separate of the Bt horizon from Pedon 1 was analyzed. The Bt horizon represents the horizon containing the maximum amount of clays accumulated through several processes: a) mineral dissolution in the overlying horizons followed by clay precipitation from the downward flux of soil pore waters, b) in situ mineral weathering and clay formation, or c) the accumulation of suspended clay particles transported downward by porefluid advection (Birkeland, 1999; Buol et al., 2003). After Mg saturation, the < 2 μm fraction of the clay separate (IA305) documented montmorillonite, kaolinite and illite were present (Figure 7), consistent with analyses reported in the literature (Ruhe, 1984a; Pye and Johnson, 1988). The Na depletion trend exhibited from north to south along this transect (Figure 5) is consistent with the loss of a Na-containing mineral. The mineralogy interpreted 20

29 Intensity Q = Quartz D = Dolomite A = Albite K = K-feldspar I = Illite M = Muscovite C = clays C I C Q A M A A C Q A M A A C Q Q A K Q A K Q K K D IA 301 Q Q Q M Q Q Q M Q Q Q IA 305 Q IA C I C M A A AC A K K Q Q Q Q Q M D D θ Figure 6a. X-ray diffractogram for Pedon 1 (northeast Iowa). IA 301 represents the surface, IA 305 represents the Bt horizon at ~ 0.75m depth, and IA 311 represents the deepest collected sample ~ 3.77m depth. In this diffractogram Q = quartz, D = dolomite, A = albite, K = K-feldspar, I = illite, M = muscovite, and C = additional clays. All samples were ground with zirconium mortar and pestle to an equivalent particle size of 45μm prior to analysis. Otherwise none of these samples were pre-treated. 21

30 Intensity Q = Quartz A = Albite K = K-feldspar I = Illite C = clays C C I I I 10 C C C I I Q I Q I Q I 20 A A A K A A C A A A C A 2 θ Q K Q A K Q A A K K K K 30 C C C Q Q Q Q Q Q A Q Q Q A Q Q Q I 40 TN 11 Q TN 16 Q A TN 19 Figure 6b. X-ray diffractogram for Pedon 13 (Tennessee). TN 11 represents the surface, TN 16 represents the Bt horizon at ~ 0.75m depth, and TN 19 represents the deepest collected sample ~ 2.10m depth. In this diffractogram Q = quartz, A = albite, K= K-feldspar, I = illite, and C = additional clays. All samples were ground with zirconium mortar and pestle to an equivalent particle size of 45μm prior to analysis. Otherwise none of these samples were pre-treated.. Q A 50 22

31 Intensity Q = Quartz A = Albite K = K-feldspar I = Illite C = clays C C I I 10 Q Q Q I I I 20 A A A A A AC A A AC 2 θ Q K K A K Q A K Q A K K K 30 C C C Q Q Q Q Q Q Q Q Q 40 A Q Q Q LA 1 Q LA 5 Q LA 14 Figure 6c. X-ray diffractogram for Pedon 22 (Louisiana). LA 1 represents the surface, LA 05 represents the Bt horizon at ~ 0.75m depth, and LA 14 represents the deepest collected sample ~ 2.53m depth. In this diffractogram Q = quartz, A = albite, K = K-feldspar, I = illite, and C = additional clays. All samples were ground with zirconium mortar and pestle to an equivalent particle size of 45μm prior to analysis. Otherwise none of these samples were pre-treated. Q 50 23

32 2000 M M = Montmorillonite K = Kaolinite I = Illite 1500 Intensity I K M M IA θ Figure 7. X-ray diffractogram for Pedon 1, sample IA 305, the Bt horizon at ~ 0.75m depth. This diffractogram was measured after Mg saturation as described in the main text. The particle size analyzed was the < 2 μm particle fraction. In this diffractogram, K = kaolinite, M = montmorillonite, and I = illite. 24

33 Table 4. Mineral Presence or Absence Based on XRD Pedon Location Horizon Quartz Dolomite Albite 1 K-feldspar 1 Illite Mica Clays 2 1 Clayton County, IL -9 A x x x x x Bt2 x x x x x x C x x x x x x x 13 Lauderdale County, TN -7 A x x x x x Bt1 x x x x x C x x x x x 22 St. Francisville, LA -8 A x x x x x Bt3 x x x x x C x x x x x 1 Albite and K-feldspar were the only feldspars identified using MDI Jade Version Clays could include both kaolinite and montmorillonite, as determined from the < 2 μm particle size fraction after Mg saturation as described in the text. 25

34 Table 5. Volumetric Mineral Abundance from Quantitative XRD 1 Pedon Location Horizon Quartz Dolomite Plagioclase 2 K-feldspar 3 Kaolinite 4 I/M/S 5 Total 1 Clayton County, IL -9 A C Lauderdale County, TN -7 A Bt C St. Francisville, LA -8 A Bt C Analyses and interpretation performed by Alex Blum, USGS, Boulder, CO using Rockjock software following procedures outlined by Eberl (2003). 2 Plagioclase includes the following phases, which are represented as a sum total here: albite, oligoclase, and andesine. 3 K-feldspar includes the following phases, which are represented as a sum total here: orthoclase, ordered microcline and intermediate microcline. 4 Kaolinite totals include halloysite and kaolinite fractions. I/M/S refers to the clay fraction, which could include illite, micas and smectites

35 from the X-ray diffractograms (Figures 6 and 7) confirms albite (NaAlSi 3 O 8 ) is present in these soils. As previously stated by Ruhe (1984a), the kaolinite and illite contents in the soils are greatest in the surface and near surface horizons of these soils and decrease with depth, consistent with secondary formation during weathering. In contrast, the expandable clay content is lowest in the near surface and surface horizons and increases with depth, consistent with the interpretation that expandable clays are present as primary minerals (Ruhe, 1984a). To quantify the composition of these pedons for modeling purposes, we developed a normative mineral model by calculating the volumetric concentrations of minerals as a function of elemental concentrations. The XRF analyses of the soil samples (Muhs et al., 2001) were reported as weight percent oxides for the major elements and parts per million for the trace elements (Appendix A). Based on our XRD observations, we assumed the following minerals were present in these soils (see also Ruhe, 1984; Pye and Johnson, 1988): albite (NaAlSi 3 O 8 ), K-feldspar (KAlSi 3 O 8 ), dolomite (CaMg(CO 3 ) 2 ), montmorillonite (Ca 0.5 MgAlSi 4 O 10 (OH) 2 ), kaolinite (Al 2 Si 2 O 5 (OH) 4 ), and quartz (SiO 2 ). Although no calcite (CaCO 3 ), chlorite (Mg 3 Fe 3 AlSi 3 O 10 (OH) 8 ), or ferrihydrite (Fe(OH) 3 ) were identified in any of our samples, to complete allocation of elements into the model required the inclusion of these minerals at minor concentration (< 5%), with maximum concentrations occurring for ferrihydrite in the Bt horizon. Consistent with this assumption, Pye and Johnson (1988) report minor amounts of calcite, chlorite, and heavy minerals such as goethite in the unleached loess (parent material). In addition to the montmorillonite, illite, and kaolinite, Pye and Johnson (1988) reported interlayer illite-smectite in these soils. For simplicity, however, we have 27

36 assumed that montmorillonite is the only primary smectite present and that it weathers to form only chlorite. In effect, the chlorite in the model may be considered a proxy for the interstratified illite-smectite-chlorite that forms during weathering of these soils. Alternatively, the minimal concentrations of chlorite can be considered as error in our model allocations related to the assumption of a constant montmorillonite composition. Stoichiometries for montmorillonite and chlorite were set equal to ideal formulas presented by Moore and Reynolds (1997). These stoichiometries correspond generally to the phases inferred from XRD using the analytical software program MDI Jade Version To calculate the mineralogical percentages in parent loess, we first assumed that Na and K are only found in feldspar; therefore, all Na and K concentrations were allocated into albite and K-feldspar, respectively. Allocating Mg was more challenging because Mg is presumed to be present or have been present in dolomite and montmorillonite. We used the following observation to constrain the allocation of Mg in the parent: in the northernmost pedon (pedon 1) and in a southern pedon (pedon 20), we observed constant values of τ Zr,Mg for the C and B horizons (over depths from m and m, respectively). These constant concentrations, depicted in Figure B3 in the Appendix, were interpreted to indicate that 60-70% of the Mg in northern Pedon 1 and 80-90% of the Mg in Pedon 20 had been lost throughout the profiles during dissolution of a very soluble phase. Pedons 2, 3, 5, and 8 also show evidence for the loss of a very soluble Mg-containing phase (a depletion profile for τ Zr,Mg ). Since dolomite was observed in the deepest sample in Pedon 1, we assumed the loss of Mg could largely be attributed to dolomite dissolution and dolomite concentrations varied among parent 28

37 loess samples. Therefore, enough Mg and Ca were allocated to dolomite to produce the 12 ± 1% dolomite by volume quantified by XRD in Pedon 1. Next, all remaining Mg was allocated to montmorillonite. The remaining Ca in the parent after allocation to dolomite and montmorillonite was allocated to calcite. Although not observed by XRD in the parent or soils, the total calcite volume percent was < 1.5%, well below detection limits for calcite using XRD, for all samples. We then allocated Fe and Al that is not present in any of the other soil minerals to ferrihydrite and kaolinite respectively. All remaining Si was allocated to quartz. Similar to the allocations for parent mineralogy described above, to estimate mineralogical content of the soils, we first allocated all the Na and K into albite and K- feldspar, respectively. For some of the soil samples, Ca was present at higher concentrations than Mg, and in others, the reverse was true. However, as indicated in Appendix A, the Ca and Mg concentrations in the soils are generally very low. Since we only observed dolomite in the deepest sample of the loess in the northernmost pedon with XRD, for the soils we assumed montmorillonite to be the dominant mineral containing Ca and Mg. Therefore, we assumed that the following relation describes the fraction of dolomite (f dolomite ) present in soils that demonstrated τ Zr,Mg depletion profiles: f dolomite 1 τ Zr, Mg = (2) The use of this relation to allocate Mg is therefore based on the assumption that the observed Mg depletion represents dolomite loss in the profiles (see Figure B3). All remaining Mg was allocated to montmorillonite. Again, if any Ca remained after 29

38 allocation to dolomite and montmorillonite, it was assumed to be present as calcite. In all cases, the total calcite volume percent for the soils was < 1%. At depths where all Ca was allocated before all Mg was allocated, the residual Mg concentrations after dolomite and montmorillonite were allocated to chlorite. Chlorite, although not observed with XRD, was found by this allocation to be present at < 1.2% in all soils. Such an allocation is consistent with the XRD patterns in that such small concentrations of a mineral would not be observed in the diffractogram. Finally, after the Na, K, Mg, and Ca were allocated, the remaining Fe, Al, and Si were allocated to ferrihydrite, kaolinite, and quartz, respectively. The mineral abundances estimated by this model are shown for Pedon 1 in Table 6. While not uniquely constrained by XRD or elemental analysis, this model allocation shows acceptable agreement with the quantitative XRD mineral abundance (Table 5). In contrast, the quantitative XRD results indicated the presence of oligoclase and andesine, as well as albite. Similarly, the total feldspars reported included ordered microcline and intermediate microcline in addition to orthoclase. The total clays quantified represent the literature observations well and indicate our normative model overpredicts kaolinite and under predicts the smectite and quartz fractions. In addition, our mineral model is consistent with reported mineralogy during weathering documented in the literature for a variety of systems. As reported by Meunier (2005), albite weathering in temperate environments frequently produces kaolinite or halloysite. In the same temperate environments, K-feldspar weathers to illite (Eggleton, 1986). Kaolinite can also result from the weathering of Na-Ca-K feldspars (Dixon, 1989), micas (Garrels and Christ, 1965) and smectites (Borchardt, 1989). The quantity of 30

39 Table 6. Calculated Volumetric Mineral Abundance from Normative Model for Pedon 1. Quartz Albite K-feldspar Montmorillonite Dolomite Calcite Chlorite Kaolinite Total Totals reported here do not include ferrihydrite. 31

40 kaolinite present in the soil environment is not only a function of parent mineralogy, but also a function of the extent of leaching, the ph of the soils, and duration of weathering (Dixon, 1989). According to the Goldich (1938) stability series, potassium feldspar shows an increased resistance to weathering in relation to albite. Therefore, the presence of kaolinite in the Peoria soils is assumed to be predominantly due to weathering of albite. Chlorite, common in metamorphic, igneous, and sedimentary rocks (Barnhisel and Bertsch, 1989), ultimately transforms into smectite; however this process generally proceeds through a series of mixed-layer minerals (Banfield and Murakami, 1998). Smectite, including montmorillonite, is commonly inherited directly from the parent material for loess-derived soils (Kleiss and Fehrenbacher, 1973). However, as the weathering intensity increases and leaching becomes pronounced, montmorillonite transforms to pedogenic chlorite, a process resulting in the interstratification of clays (Borchardt, 1989). The inferred mineralogy and evidence from the literature (Ruhe, 1984a; Pye and Johnson, 1988; Kleiss and Fehrenbacher, 1973; Eggleton, 1986; Barnhisel and Bertsch, 1989; Borchardt, 1989; Banfield and Murakami, 1998; Meunier, 2005) leads to the conclusion that a simple model for the weathering reactions in these soils can be summarized as: Albite Kaolinite K-feldspar Illite Montmorillonite Pedogenic chlorite + interstratified smectite 32

41 Based entirely on chemistry, we believe our simple model provides a reasonable representation for mineral weathering in the loess and soils. Bulk density As summarized in Appendix A, Na 2 O concentrations for each pedon generally decrease from the parent (in the C horizon) upward through the B horizon. We have hypothesized that Na concentrations (g Na / g soil) (see Appendix A and C) change due to Na depletion (i.e. albite dissolution). However, two other possible phenomena could contribute to decreases in Na concentration upward in the soil profiles: i) other elemental gains within the profile such that Na concentration appears to decrease or ii) the soil is expanded resulting in nonisovolumetric weathering and apparent decrease in Na concentrations. To solve (i), we seek to model variations in Na concentration in the soil using mass Na / volume of soil. For these soils, no bulk density values were reported. Assuming mass balance on the immobile element Zr for a given mass of soil weathering from a parent with volume V p and bulk density ρ p to a soil with volume V w and bulk density ρ w, we can write mass balance with the following relation: V V p w ρ ρ p w = m m Zr, w Zr, p (3) Here, m Zr,w and m Zr,p are the mass of Zr in weathered soil and mass of Zr in parent material, respectively. If these soils weathered isovolumetrically, then we can assume 33

42 V p = V w (4) Although not explicitly stated in the literature, we assume the upland, non-eroding status of these Peoria loess pedons and the minimal time for pedogenic processes to proceed (Muhs et al., 2001) can be interpreted to represent isovolumetric weathering conditions. If isovolumetric weathering has occurred, bulk density changes throughout the profiles can be expressed by: ρ w Zr, p ρ p m = (5) m Zr, w Bulk densities for typical loess have been reported to range from g cm -3 (Pye, 1987). For unaltered Peoria Loess, published bulk density along this transect was reported to range from g cm -3 (Bettis et al., 2003). We utilize an average bulk density of 1.43 g cm -3 to represent the parent loess along the Mississippi River transect (personal communication, D. Muhs). Following equation (5), we calculated normalized albite concentrations (mol albite / m 3 soil) within the profiles relative to changes in density and Zr concentrations (see Appendix C). All further calculations utilize these recalculated albite concentrations on a volume basis (C, mol Na / m 3 soil) to investigate how concentration varies as a function of depth within the soil profiles along this transect. 34

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