Chemical Geology 290 (2011) 89–100

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Chemical Geology

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Research paper

Soil profiles as indicators of mineral weathering rates and organic interactions for aPennsylvania diabase

E.M. Hausrath a,⁎, A.K. Navarre-Sitchler c,1, P.B. Sak b, J.Z. Williams c, S.L. Brantley c

a University of Nevada, Las Vegas, Las Vegas, NV, 89154, United Statesb Department of Earth Sciences, Dickinson College, Carlisle, PA 17013, United Statesc Earth and Environmental Systems Institute, The Pennsylvania State University, University Park, PA 16802, United States

⁎ Corresponding author. Tel.: +1 702 895 1134; fax:E-mail address: Elisabeth.Hausrath@unlv.edu (E.M. H

1 Present address: Civil and Environmental Engineer1500 Illinois St., Golden, CO 80401, United States.

0009-2541/$ – see front matter © 2011 Elsevier B.V. Alldoi:10.1016/j.chemgeo.2011.08.014

a b s t r a c t

a r t i c l e i n f o

Article history:Received 22 December 2010Received in revised form 17 August 2011Accepted 27 August 2011Available online 3 September 2011

Editor: J.D. Blum

Keywords:BasaltDiabaseWeatheringTrace elementsBiosignature

Basaltic bedrock dissolves quickly, and its weathering rate is therefore important towards controlling the com-position of natural waters, soil formation, and CO2 concentrations in the atmosphere. Despite its importance,however, few reports of basalt or diabase and gabbro weathering rates exist in the literature, and most havebeen measured in laboratory dissolution experiments or based on watershed studies. Here, using elementalprofiles measured through regolith on a Jurassic diabase dike in south-central Pennsylvania, we calculatetime-integrated log dissolution rates (mol m−2 s−1) of the primaryminerals plagioclase (−14.9 s−1) and au-gite (−14.8), and of smectite (−17.6), a secondary clay mineral formed in the soil. Characteristic patterns inelemental profiles are consistent with preserved signatures of corestone formation. Elemental and mineralsignatures of the soils relative to the parent rock are compared to predictions from citrate-containing basaltcolumn dissolution experiments. Depletion of apatite and of Al, Fe, Mn, Ti, P, Y, Ni, Cr, Sc, V, Ga, Cu, Zn, andLa are observed in the upper meter of the profile relative to the parent rock.

+1 702 895 4064.ausrath).

ing, Colorado School of Mines,

rights reserved.

© 2011 Elsevier B.V. All rights reserved.

1. Introduction

Silicate weathering is a very important process on Earth, contrib-uting to soil formation, water chemistry, and controlling atmosphericCO2 concentrations over geologic time scales. Basalt, as a quickly dis-solving rock, plays a particularly important role in the global carboncycle — although basalt makes up less than 5% of the Earth's subaeri-ally exposed surface, approximately 30–35% of the carbon dioxide re-moved from the atmosphere by silicate weathering is estimated to beremoved by basalt (Dessert et al., 2003). In addition, basalt is current-ly being studied as a substrate useful for geologic CO2 sequestrationand mineral carbonation (Matter et al., 2009). The dissolution ratesof basalt therefore have important implications globally.

Despite its importance, however, relatively few basalt dissolutionrates exist in the literature. Basalt dissolution rates have been previ-ously measured in the laboratory using batch and flow-through ex-perimental setups (Gislason and Eugster, 1987; Eick et al., 1996a;Eick et al., 1996b; Oelkers and Gislason, 2001; Gislason and Oelkers,2003; Hausrath et al., 2009; Hausrath and Brantley, 2010). Field basaltweathering has been studied in watersheds (Gislason et al., 1996;

Louvat and Allegre, 1997; Louvat and Allegre, 1998; Dessert et al.,2001; Dessert et al., 2003; Gaillardet et al., 2003; Das et al., 2005), insoil profiles (Chadwick et al., 2003), and on weathering rinds (Porter,1975; Colman and Pierce, 1981; Oguchi and Matsukura, 1999; Saket al., 2004; Hausrath et al., 2008a; Hausrath et al., 2008b; Pelt et al.,2008; Sak et al., 2010). Mineral dissolution rates can be calculatedfrom gradients that extend through soils into the weathered andthen unweathered parent rock e.g. (White, 2002; Jin et al., 2010).However, few such profiles formed on basaltic rocks or rocks of basal-tic composition (e.g. diabase, gabbro) exist in the literature (Gardneret al., 1981; Rice et al., 1985; Pavich et al., 1989; Buol and Weed,1991; Oh and Richter, 2005). In addition, dissolution rates of morecrystalline rocks such as diabase may be slower than less crystallinerocks such as basalt, although the rates have been shown to be similarbetween glass and minerals of similar composition (e.g.Hamiltonet al., 2000; Wolff-Boenisch et al., 2006).

Here we examine the chemical and mineralogical depth profilesthrough a 4 m thick soil formed on top of a 200 ma Pennsylvania dia-base. From weathering profiles of Na, Mg and Si in the soil formed onthis diabase, we calculate dissolution rates of the primary mineralsplagioclase and augite using the method of White (2002). We alsotreat the bulk diabase with a similar approach and report a weather-ing rate for the diabasic material itself. Primary minerals such as pla-gioclase and augite often weather in soils to form secondary clayminerals such as smectites and kaolinite. Such secondary mineralscan impact the capacity for cation exchange and water-retention.

90 E.M. Hausrath et al. / Chemical Geology 290 (2011) 89–100

However, few dissolution rates of clay minerals have been reported inthe literature, and, as reviewed by Nagy (1995), most have been mea-sured in the laboratory. Here we calculate smectite dissolution ratesfrom the Mg depletion profile in the biologically impacted uppermostmeter of the soil where augite is no longer present.

In previous experiments, Hausrath et al. (2009) have documen-ted that significantly enhanced mobility of Zr, Sc, Mn, Y, La, Ce, Th,Ti, Al, P, Pb, Ni, and Fe as well as dissolution of apatite, Fe oxides, Fesulfides, and augite occurs in basalt dissolution experiments in thepresence of organic acids. Such observations could be useful astrace element biosignatures or as an “organomarker” of organicacids in paleosols on early Earth, or other solar system bodies. Herewe examine elemental profiles through the weathering profile in-cluding Si, Al, Ca, Mg, Na, K, Fe, Mn, Ti, P, Cr, Rb, Sr, Y, Zr, Nb, Ba, Ni,Sc, V, Ba, Ga, Cu, Zn, Pb, La, Ce, and Th as a field test of these previouslaboratory experiments. The examination of weathering profiles inthis work yields valuable new insights about the dissolution of pri-mary and secondary minerals, and a field test of previous experi-ments containing organic acids.

2. Methods

2.1. Field site and sample collection

The study soil developed on a dike of basaltic composition which ispart of the 200 ma Gettysburg diabase sheet (Mesozoic basin) (Sutter,1988). Samples were collected from a forested ridge-top site along theAppalachian trail (40°13′14.0″N/77°6′14.2″W) (Fig. 1) to avoid contri-butions from downslope mass movements.

Two soil cores were collected by hand using a 4 in. (10 cm) diam-eter stainless steel auger to the depth of refusal at approximately 4 m.Samples of largely unweathered rock material, inferred to be similarto the parent lithology, were collected from a nearby road cut(40°13′16.0″N/77°6′11.2″W) (Fig. 1). In addition, one rock samplewas recovered from ~260 cm during augering. Samples from theweathering profile were distinguished in the field by color, texture,and abundance of rock fragments. Core 1 was collected at intervalsof ~10 cm for bulk chemistry and mineralogy and water content.Core 2 was collected at intervals of 6–10 cm, placed in sterile Whirl-pak bags and stored at 4 °C until analysis to measure bulk density, or-ganic acids, water content and pH. Samples for bulk densitymeasurements were collected using a bulk density sampler fromcore 2 (soil moisture sampler (Soil Moisture Equipment)) at 8 cm,43 cm, 71 cm, 111 cm, 167 cm, 208 cm, 300 cm, and 358 cm depths.Samples from core 1 were dried and powdered (8510 Spex) for scan-ning electron microscopy (SEM) and X-ray fluorescence (XRF) ana-lyses. Similarly, the parent rock was powdered for chemical analysisby XRF, and a chip was mounted for analysis by back-scattered elec-tron microscopy.

2.2. Bulk chemistry

The major, minor and trace element bulk chemistry (Si, Al, Ca, Mg,Na, K, Fe, Mn, Ti, P, Cr, Rb, Sr, Y, Zr, Nb, Ba) of the solid material fromthe dried and powdered core (core 1) was determined using XRF(XRAL Laboratories, SGS Canada) (Appendix Table 3). Loss on ignition(LOI) was also reported. Bulk chemistry was determined for ~10 cm ofcore at depth intervals of ~20 cm (every other sample collected) aswell as for the sample of outcrop inferred to be similar to parent rock.Additionally, every sample (~10 cm apart) below 290 cmwas dissolvedusing lithiummetaborate fusion and analyzed on a Perkin-Elmer Ultima5300 ICP-AES (Inductively Coupled Plasma Atomic Emission Spectrom-eter in the Materials Research Institute at The Pennsylvania State Uni-versity) for the elements Al, Ba, Ca, Fe, K, Mg, Mn, Na, P, Si, Sr, and Ti(Appendix Table 8). Trace elements (Ni, Sc, V, Ba, Ga, Cu, Zn, Pb, La,Ce, and Th), were analyzed using XRF (XRAL Laboratories, SGS Canada)

at 4 depths (32 cm, 155 cm, 233 cm, and 396 cm) (Appendix Table 4),as well as for the rock collected at ~260 cm depth that was interpretedto represent parent material.

2.3. Scanning electron microscopy

Polished samples were made of powdered material from core 1,and a chip of the outcrop inferred to be similar to parent. Mineralgrain size in the parent material was measured in backscattered elec-tron micrographs to estimate surface area. Mineral compositionswere also measured using Energy Dispersive Spectroscopy (EDS) onpolished sections made of the parent material and in grain mountsfrom core 1 at depths of 0 cm, 71 cm, 155 cm, 245 cm, 325 cm, and396 cm. SEM and EDS analyses were performed on an FEI Quanta200 Environmental Scanning Electron Microscope (Materials Charac-terization Laboratory, Pennsylvania State University).

2.4. X-ray diffraction (XRD)

Mineralogy of the bulk and clay size fraction of the dried andpowdered sample was determined by XRD. Two sets of samplesfrom the same depths as the SEM analyses were ground in anagate mortar to b45 μm for XRD. One set was analyzed for bulk min-eralogy and one set prepared for analysis of the clay size fraction.Forty to forty-five milliliters of 1% sodium pyrophosphate solutionwere added to the b45 μm size powdered sample to suspend thisfraction. The soil slurry was ultrasonicated for 30 min to dispersethe clay size fraction. Samples were then centrifuged at 1000 rpmfor 1 min at 25 °C to remove the N2 μm fraction from solution. Thesupernatant was decanted and centrifuged again for 30 min at4000 rpm to concentrate the b2 μm fraction for XRD analysis. Theslurry was dropped onto a glass slide to prepare an oriented claymount for XRD identification of the clay mineralogy. The remainingfraction was air dried and re-weighed to determine the weight per-cent of the b2 μm fraction (clay size fraction) removed from the sam-ple. The samples were analyzed with a SCINTAG diffractometer usingCuKα radiation, and scans were performed at 2θ/min with a 0.2° stepsize over the range of 2–70° 2θ.

2.5. Water content and bulk density

In the laboratory, field-wet samples were weighed, oven-dried at50 °C for 48 h, and then re-weighed to determine gravimetric watercontent. Water content is reported as kg water/kg dry soil, and wascalculated as the wet weight minus the dry weight, divided by thedry weight. Bulk density was calculated for depths sampled withthe bulk density sampler (see above), by dividing the dry weight ofthe soil sampled within the bulk density sampler by the volume ofthe sampler (68.7 cm3).

2.6. Pore water concentrations and pH

The pH of the soil pore water was estimated by adding 2 ml ofdeionized water to 1 g of soil, and measuring the pH of that slurry(Brady and Weil, 2002). Although this method is commonly used, itcan yield measured pH values that are 0.2 to 0.5 units higher thanwould be measured in undiluted soil pore water (Brady and Weil,2002). Small variations in salt concentrations can additionally leadto variations of up to 0.5 pH units.

Pore waters were also extracted for analysis by adding 10 ml ofMilliQ deionized water to 30 g of field-wet samples, and then equili-brating at ~4 °C for 6 h. The slurry was then centrifuged at 4500 rpmfor 35 min at 5 °C, and then filtered (0.2 μm). Chemical composition ofpore water can vary based on the method of extracting the pore water(Lloret et al., 2011). Thefiltered solutionswere analyzed by ion chroma-tography (IC) using a Dionex 100 with an IonPac AS-18 column and an

Fig. 1. The figure indicates the location of the soil cores, road cut and quarry samples. In each case the sampling location is indicated with a white point, and the diabase and dike inred. The inset in the upper left indicates the position of the maps within the state of Pennsylvania. Panel A indicates the position of all three sampling points: the road cut (northernpoint), soil cores (middle point) and quarry samples (southern point), in each case the point is shown in a box; the scale bar is 10 km. Panel B is a higher resolution map indicatingthe position of the soil cores (southern point) and the road cut (northern point); the scale bar is 1 km. Panel C is also a higher resolution map indicating the position of the quarrycorestones; the scale bar is 1 km.

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IonPac AG-18 guard column for organic acids (acetate, formate+propionate, oxalate, and citrate) and anions (chloride, bromide, andphosphate).Measured concentrationswere then corrected back to orig-inal pore water concentrations. Acetate filters were used in samplepreparation in samples below 250 cm depth. Therefore acetate concen-trations in this depth range are not reported.

3. Results

3.1. Parent material and soil characterization

The entire 4 m thick profile has experienced weathering, and in theaugered samples the O (0–20 cm), A (20–120 cm), B (120–140 cm and

Table 2Average composition with standard deviation of parent estimated, unless otherwise in-dicated, from measured concentrations in outcrop sample, float rock, and previouslymeasured outcrop samples (Davidson, 2010; Davidson and Edwards, 2010).

Value used for parent Standard Deviation

Si02 (wt %) 50.73 0.56Ti02 (wt %) 0.72 0.02AI203 (wt %) 15.38 0.25Fe203 (wt %) 11.27 0.27MnO (wt %) 0.19 0.01MgO (wt %) 7.90 0.14CaO (wt %) 11.17 0.29Na20 (wt %) 2.13 0.21K20 (wt %) 0.40 0.09P205 (wt %) 0.08 0.01Ba (ppm) 103 19Ce (ppm) 31a 32Cr (ppm) 249 31Cu (ppm) 136 123Ga (ppm) 14b 1La (ppm) 7b 8Ni (ppm) 43b 8Sc (ppm) 38b 3V (ppm) 253b 11Y (ppm) 19 1Zn (ppm) 83 NAc

Zr (ppm) 37 11bulk density 3.0 0.15

a Parent value estimated from previously measured values of 8 locations along Iron-stone Ridge, south-central Pennsylvania (Davidson, 2010; Davidsonand Edwards, 2010).

b Parent value estimated from previously measured values of 8 locations along Iron-stone Ridge, south-central Pennsylvania (Davidson, 2010; Davidson and Edwards,2010) averaged with rock recovered during augering.

c Parent composition estimated from rock recovered during augering only.

92 E.M. Hausrath et al. / Chemical Geology 290 (2011) 89–100

C (N140 cm) horizons are all recognized based on color. Soil characteris-tics, including color, texture and rock fragments are reported in Table 1.

3.2. Mass-element transfer coefficient

Elemental concentrations measured in the soil cores can be normal-ized to an immobile element to account formobility ofmultiple elements(Brimhall and Dietrich, 1987; Anderson et al., 2002)

τi;j ¼Cj;w

Cj;p

Ci;p

Ci;w−1 ð1Þ

Here, τi,j is the fraction of mobile element or mineral j lost or gainedassuming that element ormineral i is immobile (w and p refer toweath-ered and parentmaterial respectively) and C is the concentration of theimmobile and mobile elements in the parent and weathered materials.Where τi,jN0, the element is enriched relative to parent, where τi,jb0,the element is depleted relative to parent (τi,j=−1 indicates that theelement is completely depleted relative to parent), and where τi,j=0the element is immobile. Parent was estimated from themeasured con-centrations in outcrop sample (Appendix Table 9), float rock (AppendixTable 9), and previously measured outcrop samples from the same dia-base (Davidson, 2010; Davidson and Edwards, 2010) (see Table 2 forthe parent composition used).

In order to determine which element is least mobile, we calculatedstrain, εi,w, the ratio of volume change to initial volume (Brimhallet al., 1991). Strainwas calculated using the bulk densitymeasurements(Fig. 2A). Ti, Al, Fe, Y and Zrwere all tested to determinewhich elementis least mobile using the following equation (Brimhall and Dietrich,1987):

εi;w ¼ ρpCip

ρwCiw

−1 ð2Þ

where εi,w is the strain calculated under the assumption that element i isimmobile, ρ is the bulk density, C is the concentration, andw and p referto weathered and parent material respectively.

Nb was below detection in the parent material, and was thereforenot used to calculate strain. Since bulk density and elemental chemistrywere measured on different cores, the chemical composition of thesample closest to the depth interval for which bulk density had beenmeasured was used. The value of the strain calculated on the basis ofZr was smaller at all points than the other elements tested for immobil-ity (Fig. 2B). Higher calculated strain values for other elements suggestthe loss of those elements byweathering (White, 1995), indicating thatZr is the least mobile element. Strain calculations for this profile(Fig. 2A) indicate dilation that is generally larger than 50% below150 cm, which is consistent with previous measurements of strainreported for weathered soils formed on diabase in the Appalachians inNorth Carolina (Oh and Richter, 2005). Specifically, in that publication,dilation was reported for soils in the depth range 0–8 m.

When normalized to Zr as the least mobile element using Eq. (1), ele-ments display different profile behaviors. Calcium,Mg, Si and Na all com-prise depletion profiles (Brantley and White, 2009): they decrease fromparent concentrations (τ=0) at depth to nearly complete depletionvalues (τ=−1) (Fig. 3). Approximately 50% of the Ca, Mg, Si and Na

Table 1Soil characteristics.

Depth (m) Color Soil texture Rock fragments

0–1.80 10YR Sandy clay silt loam Few1.80–2.10 2.5Y Loam None2.10–2.45 2.5Y Loam None2.45–3.25 10YR Sandy Few3.25–4.01 2.5Y Silty clay Few

are lost at the base of the profile, and then these elements remain largelyimmobile until approximately 150 cmdepth,where they are then strong-ly depleted. Such profiles that attain near 100% depletion at the surfaceare termed completely developed profiles (Brantley and White, 2009). Incontrast, Si, Fe, Ti andAl are all described by incompletely developed deple-tion profiles because 10% or more of these elements are retained at theland surface. Approximately 20% of the Al, Fe, Ti and Mn is lost at thebase of the profile, and then these elements remain largely immobileuntil approximately 150 cm depth, where they then become depleted(Fig. 4). Ti is less depleted at the surface than Fe or Al, and normalizedMn concentrations vary throughout the profile. Phosphorus and Y simi-larly decrease in the bottom few cm, remain relatively stable, and thendecrease above 200 cm (Fig. 5). Nickel, Sc, Cr, V, Ga, Cu, Zn and La(Fig. 6) are all depleted at the surface relative to parent concentrations.

3.3. SEM

Mineral composition was estimated by EDS of plagioclase(bytownite=Ca0.7Na0.3Al1.7Si2.3O8), and pyroxene (augite=Ca0.33-Mg2.67Fe0.33Si3.01Al0.99O1.0(OH)2) (Fig. 7). Qualitative SEM imagesalso documented epitaxial replacement of minerals by Al–Si richphases (Fig. 8). The chemistry of this Al–Si rich phase is approximate-ly 2:1 Si:Al, which is not consistent with kaolinite or halloysite. Thisratio is consistent with a smectite, however, and is thus consistentwith the alteration of the parent minerals pyroxene and plagioclaseto form a mineral of smectite-like composition.

3.4. Model mineral abundances

To interpret theweathering reactions, amodel composition of the par-ent rock was estimated by allocating elements to appropriate minerals.Specifically, all P was assigned to apatite with corresponding Ca. All Na

Fig. 2. A). Bulk density (g/cm3) versus depth for the soil developed on the Pennsylvaniadiabase (see also Appendix Table 1). For measurements of dissolution rate at ~370 to400 cm, a bulk density value of 1 (indicated on the graph by a line) was used in calcu-lations. B). Strain (εi,w)calculated for Ti, Al, Fe, Y and Zr. Higher strain values for Ti, Al, Feand Y indicate loss during weathering (White, 1995), and therefore Zr is most appro-priate as the least mobile element. C). Fraction of clay-size material in the soil devel-oped on the Pennsylvania diabase (see also Appendix Table 2). Depth of refusaloccurred at approximately 4 m.

Fig. 3. Normalized concentrations, τ, versus depth, for the soil developed on the dia-base in central Pennsylvania. The parameter τ represents concentrations normalizedto Zr, determined to be immobile by calculation of strain (Fig. 2) to account for mobil-ity of multiple elements (Brimhall and Dietrich, 1987; Anderson et al., 2002). Normal-ized concentrations of Na, Mg, Ca and Si all display two depth intervals with rapidelemental loss, one between ~370–400 cm, and one between 150 and 0 cm depth.Depth of refusal occurred at approximately 4 m.

93E.M. Hausrath et al. / Chemical Geology 290 (2011) 89–100

was assigned to bytownite (Ca0.7Na0.3Al1.7Si2.3O8)with corresponding Ca,Si, and Al. All Mg was assigned to augite (Ca0.33Mg2.67Fe0.33Si3.01Al0.99O10

(OH)2)with corresponding Ca, Fe, Si, and Al. All Ti was assigned to ilmen-ite (FeTiO3)with corresponding Fe, and remaining Fe assigned tomagne-tite. Assuming 5% porosity, these mineral abundances indicate that therock is ~66% bytownite, 24% augite, ~5% magnetite and 1% ilmenite(Table 3).

3.5. Mineralogical profiles

X-ray diffractograms of samples with depth document that plagio-clase was present throughout the profile while pyroxene was present

400

300

200

100

0

-1.0 -0.9 -0.8 -0.7 -0.6 -0.5 -0.4 -0.3 -0.2 -0.1 0.0τ Zr,j

Dep

th (

cm)

FeMnAlTi

Fig. 4. Normalized concentrations, τ, versus depth, for the soil developed on the dia-base in central Pennsylvania. The parameter τ represents concentrations normalizedto Zr determined to be immobile by calculation of strain (Fig. 2) to account for mobilityof multiple elements (Brimhall and Dietrich, 1987; Anderson et al., 2002). Normalizedconcentrations of Fe, Al and Ti are relatively immobile below ~100 cm, at which pointapproximately 60% (Ti) to 80% (Al and Fe) of the elemental concentration present inthe parent material is lost. This effect is most likely due to decreasing pH or increasingconcentrations of organic acids. Mn concentrations show periodic increases at depth,most likely due to oxidation effects, and increases at the surface, most likely due todust inputs. Depth of refusal occurred at approximately 4 m.

Fig. 7. Backscattered electron micrograph of a polished chip of the parent material(outcrop) of the Pennsylvania diabase. The dark gray minerals are plagioclase, the ligh-ter gray minerals are pyroxene, and the white minerals are iron oxides.

Fig. 5. Normalized concentrations, τ, versus depth, for the soil developed on the dia-base in central Pennsylvania. The parameter τ represents concentrations normalizedto Zr calculated to be the most immobile element based on strain (Fig. 2) to accountfor mobility of multiple elements (Brimhall and Dietrich, 1987; Anderson et al.,2002). Normalized concentrations of P and Y display depletion most likely due to dis-solution of apatite, and P shows a slight enrichment at the surface, most likely due tobiological cycling. Depth of refusal occurred at approximately 4 m.

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below a depth of ~150 cm (Fig. 9). Quartz was observed under XRD tobe present only at depths shallower than ~70 cm (Fig. 9), and was notdetected by XRD elsewhere. Kaolinite/chlorite and goethite were pre-sent throughout the profile, and smectite was observed at all depthsdeeper than ~70 cm. The only two primary minerals identified byXRD were pyroxene and plagioclase (Fig. 9).

3.6. Water content and pore water chemistry

Water content decreases with depth largely in two depth ranges, be-tween 0 and 200 cm, and at ~400 cm (Fig. 10A; Appendix Table 5). Porewater pH increases steadily with depth (Fig. 10B; Appendix Table 6)below ~30 cm; however, the shallowest sample shows a relatively highpH. Oxalate concentrations were measured to be highest at 125 cm

Fig. 6. Normalized concentrations, τ, versus depth, for the soil developed on the dia-base in central Pennsylvania. The parameter τ represents concentrations normalizedto Zr calculated to be the most immobile element based on strain (Fig. 2) to accountfor mobility of multiple elements (Brimhall and Dietrich, 1987; Anderson et al.,2002). All normalized concentrations are relatively immobile at depth, and then de-crease at the surface. Depth of refusal occurred at approximately 4 m.

(Fig. 10C). Acetate and formate–propionate (these were not distinguish-able by the IC) both were highest in concentrations at approximately175 cm.

3.7. Clay fraction percentage and bulk density

The weight percent of clay-sized particles (measured as describedabove) generally decreases with depth. However, two peaks are ob-served in the clay-size fraction, one at ~71 cm, where clay percentageis approximately 16%, and one at about 325 cm depth where the clayfraction reaches 5 wt.% (Fig. 2C; Appendix Table 2). The bulk density(ρw) of the soil varied from 0.65 to 1.29 g cm−3 over the entire depthof the soil column; at depths greater than approximately 1 m the bulkdensity values were very close to 1 g cm−3 (Fig. 2A; AppendixTable 1). Bulk density (Fig. 2A) shows a similar pattern with depth asclay fraction percentage.

Fig. 8. Representative backscattered electron micrograph of a polished grain mount ofthe dried and powdered soil from core 1 (325 cm depth). Bright areas indicate mineralsthat are higher in heavy elements, and darker areas indicate minerals that are less richin heavy elements. This image and similar images document epitaxial replacement ofthe primary minerals by Al–Si rich phases. The primary mineral augite (A) is sur-rounded by weathered material (arrows). The chemistry of the Al–Si rich phase is ap-proximately 2:1 Si:Al as measured by EDS, which is not consistent with kaolinite orhalloysite, but is consistent with smectite.

Table 3Model mineral abundances in parent.

Mineral Formula Element Model volume (%)

Apatite Ca5(PO4)3(OH) P 0.09Bytownite Ca0.7Na0.3Al1.7Si2.3O8 Na 65.7Augite Ca0.33Mg2.69Fe0.33Si3.01Al0.99O10(OH)2 Mg 23.4Ilmenite FeTiO3 Ti 0.85Magnetite Fe3O4 Fe 4.87

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4. Discussion

4.1. Elemental and mineralogical weathering profiles

From the chemistry of the entire depth profiles through the soil(Figs. 3–5), two dissolution reaction fronts are identified. Dissolution re-action fronts represent depth intervals over which normalized elementalconcentrations decrease due to mineral dissolution (White, 2002;Brantley and White, 2009). Such fronts generally document infiltratingsolutions that are corrosivewith respect to themineral of interest because

Fig. 9.Mineral presence throughout the soil profile identified by XRD. Plagioclase and augitepyroxene below a depth of ~70 cm. Kaolinite and goethite are present throughout the soilquartz is present only at depths b~70 cm. Depth of refusal occurred at approximately 4 m.

they are undersaturated, sometimes due to thepresence of carbonic or or-ganic acids or other agents that enhance dissolution. Reaction fronts havebeen identified inmanyweathering systems (Lichtner, 1988; Steefel et al.,1990; White, 2002; Sak et al., 2004; Steefel et al., 2005; Hausrath et al.,2008a; Maher et al., 2009; Navarre-Sitchler et al., 2009; Sak et al., 2010).The two reaction fronts identified in these profiles extend from~370 cm to point of refusal (~4 m), and from ~0 cm to ~150 cm(Figs. 3–5). These fronts are also largely consistentwith zones of increasedclay size fraction content (Fig. 2C). The lower reaction front occurs belowthe visible zone of active root formation and the depth interval that con-tains high concentrations of organic acids (Fig. 10C; Appendix Table 7).The upper reaction front occurs within the zone of higher organic acidconcentrations and active root formation, indicating that biological activ-ity may influence weathering strongly in this zone.

Normalized concentrations of the major elements Ca, Mg, Si and Nashow decreasing τi,j values in both reaction fronts: from ~370 cm to400 cm and ~0–150 cm (Fig. 3). In addition to the decreases in concen-tration observed in the two identified reaction fronts, between the reac-tion fronts some elemental concentrations also vary (Fig. 3). In this zone,we have also identified the presence of significant concentrations of the

are present in the parent material. Plagioclase is present throughout the soil profile, andprofile; smectite is present throughout the soil profile below a depth of ~70 cm, and

Fig. 10. A). Water content plotted versus depth for the soil developed on the Pennsyl-vania diabase for cores one and two (described in text) (Appendix Table 5). B). Soilpore water pH versus depth (Appendix Table 6). C). Organic acid concentrations versusdepth in soil pore waters (Appendix Table 7). Pore waters were extracted by equili-brating soil samples as described in text. Measured concentrations were correctedback to original pore water concentrations using the measured weight percent water(Fig. 10A). Acetate filters were inadvertently used in the sample preparation in samplesbelow 250 cm depth, and therefore acetate concentrations in these samples are notreported. Oxalate concentrations are multiplied by a factor of 10. Depth of refusal oc-curred at approximately 4 m.

96 E.M. Hausrath et al. / Chemical Geology 290 (2011) 89–100

clay mineral, smectite. Some of the variations in element concentrationsin this zone are therefore presumed to be related to variations in smectitechemistry. However, we propose here that this variation may also be aresult of spheroidal weathering of the bedrock as explained in the nextparagraph. Spheroidal weathering has been previously observed in aquarry near Gettysburg, PA,where the diabase outcrops and suchweath-ering is inferred to possibly be occurring in the subsurface where theaugering was completed (Fig. 11A).

Spheroidal weathering is the process by which corestones weatherconcentrically due to spheroidal fracturing. These fractures promoteweathering that occurs in “onion-skin”-like annular rock layers calledrindlets. As fracturing creates rindlets, weathering progresses into theinner corestone with each new spheroidal fracture. Outermost rindletsare thus more extensively weathered than inner rindlets which aremuch more weathered than the center of the corestone. In general, itis possible that the outermost rindlets of corestones located at severaldepths are extensively — and similarly —weathered, while the interiorof the corestones are equally pristine.

Spheroidal weathering has been previously observed in quarry out-crops of the Pennsylvania diabase near Gettysburg, PA (Fig. 11A), sug-gesting that it may be occurring beneath the soil sampled by augeringin this study. Chemical variation through chemically weathered cores-tones has been documented (Patino et al., 2003). Previous work sug-gests that when such spheroidally weathered bedrock forms soil, thechemical variation due to rindletting can persist as a chemical signaturein saprolite. Specifically, in a weathered quartz diorite in Puerto Rico,saprolite layers were identified sometimes as “ghost rindlets” due tochemical signatures related toweathering (Buss et al., 2010). Therefore,transformation of spheroidallyweathered rindlets to saprolitemay helpexplain the chemical variation observed in Figs. 3–5 especially between~350 and 150 cm depth. In order to compare variation observed in theelemental profiles to such putative fracturing, the original depth inter-vals were back-calculated by correcting the strain (Egli and Fitze, 2000)

Δz ¼ Δzw1

εi;w þ 1ð3Þ

Here Δz indicates the original depth interval,Δzw is the depth intervalin theweathered profile, and ε,iw is the strain of theweathered profile cal-culated using immobile element i. Normalized concentrations of Na areplotted versus original depth in Fig. 11C. In Fig. 11, the observed fracturingrevealed in a photograph in the quarry wall in the Gettysburg diabase(below the weathered zone, Fig. 11A) and a histogram of corestonesizes (after removal from the weathering zone by the mining operation,Fig. 11B) are compared to the variations in normalized elemental concen-trations plotted versus original depth (Fig. 11C). The histogram docu-ments measurements of corestones found in the weathered section: thelargest measured corestone (90–180 cm) is inferred to be characteristicof the size of the original corestones (=fracture spacing) whereas thesmaller measured corestones are all that is left after extensive spheroidalweathering. Thefigure also includes a schematic of how such a depth pro-file of corestone weathering might progress with time (Fig. 11D). InFig. 11D, the dashed horizontal lines show the position of original frac-tures delineating corestones and the horizontal gray lines show the posi-tion of rindlet-forming fractures. According to this model, the depths thatare more depleted in Na (original depths 160 cm and 79 cm) may corre-spond to the depths of the fracture zones that opened relatively early be-tween corestones. These zones were thus weathered longer and showgreater extent of weathering. In contrast, the material between thesezones may correspond to rindlets that have not been exposed as long toweathering or to the unweathered interior parts of the corestones. In ef-fect, because the corestones weather concentrically inward due to sphe-roidal fractures, some deeper layers of regolith may have consisted ofrindlets that formed early and are thusmore highly weathered than shal-lower material that comprised the corestone centers and that haveremained less highly weathered. Spacing of the original corestone-

Fig. 11. A) Photograph of weathered corestones from the PA diabase, exhumed in a local quarry (Fig. 1). B) A histogram of mean corestone sizes, showing that the largest corestones,most likely corresponding to the original corestone sizes, are approximately 0.9-1.8 m (unpublished data from R. Fletcher) C) Sodium concentrations normalized to the assumedimmobile element Zr, plotted versus original soil depth calculated from the strain as described in text. The upper limit is truncated at 50 cm depth because of changes in bulk den-sity. D) Schematic figure of τ progressing through time. The bottommost figure is the newly fractured bedrock, with the corestones indicated by τ=0, and then the depletions in-dicating chemical weathering between the corestones. As chemical weathering occurs at the soil surface, τ decreases, and volume increases due to the formation of secondaryminerals. The dotted lines indicate the fractures, and the light gray lines indicate areas that experience weathering between the corestones. Squiggles indicate that this processcould be occurring at some depth.

97E.M. Hausrath et al. / Chemical Geology 290 (2011) 89–100

defining fractures in quarry walls was estimated to equal 90–180 cm(Fig. 11B). This value approximately corresponds to the distance betweenhighly depleted zones (Fig. 11C). As observed in the quarrywalls (Fig. 11AandB), however the actual fracture spacing is variable, indicating that cor-estone size varies. Thus, the conceptual model in Fig. 11D does not fullymatch the data in Fig. 11A, B and C. In some cases, relatively highlyweath-ered zones may have been fractured into smaller corestones as observedin the quarry walls (Fig. 11A). Furthermore, exact attribution of positionof corestones in Fig. 11C is difficult due to a variety of factors, includingthe sample spacing of approximately 10 cm, which contributes to appar-ent smoothing of the depletion profiles.

4.2. Surface-area normalized dissolution rates

Surface-area normalized dissolution rates of augite and plagio-clase were calculated from the slope of the concentration normalizedby measured, extrapolated, and interpolated Zr concentrations versusdepth curves, i.e. the weathering gradient (White, 2002), for the dee-pest reaction front:

Rs ¼ 1000ρw

SBETβωbs

ð4Þ

Here Rs (mol m−2 s−1) is the surface-area normalized weatheringrate, ρw (g cm−3) is the density of weathered material, SBET (m2 m−3)is the specific surface area, β (mol mol−1) is the stoichiometric coeffi-cient for the element in amineral,ω (m s−1) is theweathering advancerate, and bs (m kgmol−1) is the weathering gradient. A similar calcula-tion of the dissolution rate of the bulk rock was also made using bulkcomposition as the diabase formula unit.

Average bulk density values were very close to 1 g cm−3 (Fig. 2A;Appendix Table 1), and that value was therefore used in calculations.The stoichiometric coefficientswere determined from the compositions

of the minerals determined by EDS, and were assumed not to vary sig-nificantly. For the bulk rock, the stoichiometric coefficient was assumedto be one Si atom per diabase formula unit. The weathering gradientwas calculated from the linear portion of the zone of decreasing τ valuesbelow 370 cm for eachweathering profile (pyroxene based onMg, pla-gioclase Ca, diabase, Si).

Reactive surface area (SBET) is difficult to constrain even in laborato-ry experiments, and may be the largest source of uncertainty in suchfield calculations (White, 2002). Some have argued that the geometricsurface area is a better approximation of the reactive surface area thanthe BET surface area (e.g. (Gautier et al., 2001). During soil-forming pro-cesses, the surface area changes as particle size changes, and the surfacearea bathed by flowing fluids is affected by hydrologic status which isdifficult to determine (Velbel, 1993). In addition, while BET-measuredsurface area may increase with time of reaction, the reactive surfacearea (area of reactive surface sites) may actually decrease (White andBrantley, 2003). Furthermore, the growth of clayminerals on the augiteand plagioclase mineral grains (Fig. 8) is likely to alter BET-measuredsurface areas during weathering. Therefore, for these calculations, thesurface area was simply estimated from the grain diameter of plagio-clase and pyroxene as measured in back-scattered electron micro-graphs of the parent material (20 grains per mineral), and the diabasesurface area was considered to be a weighted average of those values.Geometric surface area was estimated assuming spherical geometry. ABET surface area (SBET) was then estimated from the geometric surfacearea by Eq. (5):

SBET ¼ Sgeoλ ð5Þwhere λ is the roughness factor (Helgeson et al., 1984) and Sgeo is thegeometric surface area. Values for λ were assumed to be 10 (Whiteand Peterson, 1990). Plagioclase and pyroxene surface areas weretherefore estimated as 24,100 m2/m3 and 10,200 m2/m3 respectively,and smectite surface areas were estimated from the literature as 51

98 E.M. Hausrath et al. / Chemical Geology 290 (2011) 89–100

m2/g (Kuwahara, 2006; Yokoyama et al., 2005; Metz et al., 2005;Bosbach et al., 2000).

The weathering advance rate, ω, was assumed equal to the erosionrate, which is the case when weathering profiles are in steady state. Anumber of observations suggest that steady state is a reasonable as-sumption in the Susquehanna River Basin. For example, White (2009)documented that erosion rates and chemical dissolution rates are ap-proximately equal in the central Appalachian Mountains of Pennsylva-nia. Similarly, in the immediate region of our study, Potter et al.(2009) documented similar suspended and dissolved load denudationrates, also consistent with steady state.

A number of erosion rates have been calculated in the Appalachianmountain range. Reuter (2005) estimated erosion rates of the Sus-quehanna River basin using the cosmogenic nuclide 10Be and assum-ing steady state. Erosion rates were determined from small GISselected basins (4–54 mm ky−1) and large USGS non-glaciated basins(8–22 mm ky−1), a value of 5 mm ky−1 is calculated from the aver-age difference between the inferred values and predicted valuesfrom the GIS selected basins (Reuter, 2005).

Hancock and Kirwan (2007) calculated erosion rates of 5.7 mmky−1for ridgetops in Dolly Sods, West Virginia, an Appalachian paleoper-iglacial plateau. They note that this rate is much lower than similar fluvialincision rates in the same area, suggesting that ridgetops erode moreslowly than incising channels. They also note that this rate is similar topreviously measured erosion rates on summit tops in the Rocky Moun-tains and Sierra Nevada (Small et al., 1997), Georgia (Bierman et al.,1995), Kentucky (Granger et al., 2001) as well as the Virginia Piedmont(Pavich et al., 1985). Although the lithologies are different, the value of5 mm ky−1 we use for this diabase ridge is therefore a reasonablevalue based on its topographic position. Note that if the weathering ad-vance rate is doubled (to 10 mm ky−1), weathering rates double, and ifthe rate is halved (to 2.5 mm ky−1, then weathering rates are similarlyhalved. Uncertainty due to assumptions regarding the weathering rateand the erosion rate is therefore likely no greater than a factor of two.The time-integrated dissolution rates of the diabase and plagioclase andaugite are: diabase (log rate=−14.4 mol diabase m−2 s−1), plagioclase(log rate=−14.9 mol mineral m−2 s−1), and augite (log rate=−14.8 mol mineral m−2 s−1). These rates are ~1–4 orders of magnitudeslower than laboratory dissolution rates at similar pH values.

Similar to our calculation of dissolution rates for pyroxene, plagio-clase, and the bulk diabase, we also calculate dissolution rates forsmectite in the upper meter. For this calculation we used the averagebulk density for the depth range of 1 g cm−3. The upper meter con-tains substantially more clay-sized particles than other depths(Fig. 2C; Appendix Table 2), and at the surface smectite is notdetected by XRD (smectites are otherwise present throughout the en-tire profile) (Fig. 9). From the XRD analyses (Fig. 9), augite is assumedto be absent in the upper 71 cm. Therefore, all of the Mg loss (from~80% to 95% in the top 70 cm), is attributed to dissolution of thesmectite, which is consistent with the loss of Al and Fe over thesame depth range. Surface area normalized smectite dissolutionrates are therefore calculated from the Mg concentrations normalizedwith Zr concentrations in this depth range. Smectite is likely dissol-ving as indicated below:

Ca0:33Mg2:67Fe0:33Si3:01Al0:99O10 OHð Þ2 sð Þ þ 10Hþaqð Þ→0:33Ca2þ aqð Þ

þ 2:67Mg2þ aqð Þ þ 0:33Fe2þ aqð Þ þ 3:01H4SiO4 aqð Þ þ 0:99Al3þ aqð Þ ð6Þ

The logarithm of the smectite dissolution rates (mol smectitem−2 s−1) calculated in this manner is –17.6.

4.3. Trace elements as organomarkers

In addition to the use of elemental profiles to calculate dissolutionrates, we analyze profiles here as a field test of the utility of trace

elements as “organomarkers”. Based on laboratory experiments of or-ganic acids with basalt rock, Neaman et al. (2005a; 2005b) postulatedthat Fe, P, and Y concentrations in regolith can sometimes be indicatorsof the presence of organic acids, and determined evidence consistentwith this effect in paleosols. Hausrath et al. (2009) further determinedthat enhanced mobility of the elements Zr, Sc, Mn, Y, La, Ce, Th, Ti, Al,P, Pb, Ni and Fe related to enhanced dissolution of apatite, Fe-sulfides,Fe-oxides and augite occurred in basalt in the presence of organicacids in laboratory column dissolution experiments. The elemental pro-files in this soil are therefore examined as a field test of those laboratoryexperiments to determine whether any elements act as “organomar-kers”. As discussed above, the top meter of soil has strong indicationsof biological activity (visible root activity, and increased concentrationsof organic acids, Fig. 10C; Appendix Table 7), andwe therefore compareelemental concentrations in this depth interval to the deeper profiles asa test of potential organomarkers.

We first consider Al, Fe, Mn and Ti as possible “organomarkers”, i.e.we consider if these elements are preferentially mobilized from thePA regolith in the presence of such acids as observed in the laboratoryexperiments of Hausrath et al. (2009). Approximately 20% of the Al,Fe, Mn and Ti is lost within a few cm of the point of augering refusalat 4 m. As discussed previously, this zone may be a zone of fracturingthat has channelized fluid flow during regolith and corestone formation.Thus, the significant loss of Al, Fe, Mn and Ti at this depth could simplybe due to the large fluid flux relative to underlying depths. However, itis also possible that microbiota participate in weathering at 4 m. For ex-ample, in spheroidally weathering quartz diorite in Puerto Rico, an eco-system was identified at the corestone–regolith interface at 5 m depththat putatively was based on Fe-oxidizing bacteria as primary producers(Buss et al., 2005;Minyard et al., 2011). Both heterotrophic and autotro-phic bacteria were identified at the bottom of the saprolite in the PuertoRico system. If such an ecosystemwere present at 4 m in the PA diabase,then Fe, Ti, Al andMn could be organomarkers of that ecosystem at thatdepth.

Above this depth, concentrations of Fe, Ti, Al and Mn remain rela-tively immobile until ~150 cm depth, at which point ~60–80% ofthese elements are lost (Fig. 4). Over this depth interval, these ele-ments are presumably retained in both primary minerals and in sec-ondary precipitates of smectite and goethite. The surface loss maybe due to pH, which decreases towards the surface, thereby enhanc-ing Al and Fe mobility (Fig. 10B; Appendix Table 6), and/or to thepresence of organic acids complexing these elements. In addition, itis likely that more fluid flows through the uppermost meter of rego-lith, contributing to mobilization of elements. These field results areconsistent with these elements acting as organomarkers (Hausrathet al., 2009), although contributions to mobility from enhanced fluidflow cannot be separated from contributions due to organic acids.

Goethite is present throughout the profile (Fig. 9), because Fe isweathered from primary minerals, and then reprecipitated as goethite.Therefore, the loss of Fe at the surface presumably documents the loss ofgoethite as well as smectite as discussed above. The variation of Mnthroughout the profile is likely due to redox effects that impact Mnmo-bility. For example, Mn2+ can be easily oxidized in aerated soils to formthe less mobile Mn4+ (McKenzie, 1989). The increase of Mn at the sur-face could also be due either to biocycling of Mn into the upper soillayers or to Mn input in dust (Herndon et al., 2011).

Phosphorus and Y are similarly depleted at the base of the profilenear 4 m. As discussed above, this depletion can be attributed to eitherhighfluid fluxes or organic acids at that depth. Above this zone of deple-tion, these elements are relatively immobile, only decreasing at the sur-face above about 1.5 m depth. Some P enrichment in the uppermostsample is likely due to biological cycling. Both Y and P are likely con-tained in the mineral apatite. Enhanced dissolution of apatite wasnoted in the basalt column experiments of Hausrath et al. (2009) inthe presence of organic acids. Absence of apatite therefore could insome cases be considered a potential “organomarker”. Again, results

99E.M. Hausrath et al. / Chemical Geology 290 (2011) 89–100

reported here may be broadly consistent with the use of apatite as anorganomarker: apatite dissolves from the regolith largely in twozones, both of whichmay havemore concentrated organic acids. Mobil-ity of phosphate in soils may also be impacted by changes in oxidationand reduction if it is present in iron-containing phosphates, as it oftenis (Lindsay et al., 1989).

Like Ti, Al, Fe, Mn, P, and Y, above the zone of initial depletion atdepth 4 m, Ni Cr, Sc, V, Ga, Cu, Zn and La were all relatively immobilethroughout the profile. However, between 1.5 m and the surface,these elements again showed depletion (Fig. 6). Of these elements,scandium, La and Ni were indicated to be strongly mobilized from ba-salts by organic acids in column experiments (Hausrath et al., 2009).

5. Conclusions

Mafic rocks dissolve quickly, and therefore strongly impact the com-position of natural waters, soil formation, and the carbon cycle overgeological time scales. Here we calculate time-integrated dissolutionrates of a diabase and several individual minerals: diabase (log rate=−14.4 mol diabase m−2 s−1), plagioclase (log rate=−14.9 mol mi-neral m−2 s−1), augite (log rate=−14.8 mol mineral m−2 s−1) andsmectite (log rate=−17.6 mol mineral m−2 s−1). These rates are~1–4 orders of magnitude slower than laboratory dissolution rates atsimilar pH values. We attribute the slow rates of mineral reaction toslow erosion in the Piedmont. Slow erosion allows regolith thicknessto increase, limiting the infiltration of corrosive fluids into unalteredrock at depth. At the reaction front, the rate of weathering of these min-erals may be transport-limited (Lebedeva et al., 2010): hence, the pri-mary mineral dissolution rates are similar. The patterns of elementaldepth profiles are broadly consistent with alteration of corestones inthat element concentrations do not increase steadily with depth backto parent, but show depth intervals of higher and lower depletionthat we attribute to rindlet zones and corestone interiors respectively.

In addition, elemental andmineral signatures in this vegetated en-vironment are compared to predictions from organic-containing ba-salt dissolution column experiments. Depletion of apatite and of Al,Fe, Mn, Ti, P, Y, Ni, Cr, Sc, V, Ga, Cu, Zn and La are observed in twozones i) the zone at the base of the augerable saprolite, and ii) theupper meter of the profile. Based on observations of regolith formedon other spheroidally weathering rocks, both of these zones may becharacterized by higher concentrations of organic acids and higherfluid fluxes. Depletion of these elements could therefore be consistentwith these elements acting as organomarkers. Specifically, in previousbasalt dissolution column experiments enhanced mobility of the ele-ments Zr, Sc, Mn, Y, La, Ce, Th, Ti, Al, P, Pb, Ni and Fe and enhanceddissolution of apatite, Fe sulfides, Fe oxides and augite in basaltwere considered signatures of organic acids (Hausrath et al., 2009).Future work is warranted to explore both the role of organic acidsand high fluid fluxes in controlling the behavior of elements in soilprofiles.

Acknowledgements

We thank H. Gong, M. Angelone, J. Moore, J. Catalina, L. Liermann, J.Macalady, B. Peightal, E. Cabret, M. Nyblade, R. Fletcher and B. Thomasfor their help. This work was supported by funding from the NationalScience Foundation IntegrativeGraduate Education and Research Train-eeship grant DGE-9972759, the Penn State Biogeochemical ResearchInitiative for Education. E.M.H. is also grateful for support from the Na-tional Science FoundationGraduate Research Fellowship Program. S.L.B.acknowledges support from the NSF-funded Center for EnvironmentalKinetics Analysis under NSF grant NSF-CHE-0431328 as well as fundingfrom the US Department of Energy grant DE-FG02-05ER15675. PBS ac-knowledges permission from the National Park Service to sample alongthe Appalachian Trail Easement.

Appendix A. Supplementary data

Supplementary data to this article can be found online at doi:10.1016/j.chemgeo.2011.08.014.

References

Anderson, S.P., Dietrich, W.E., Brimhall, G.H., 2002. Weathering profiles, mass balanceanalysis, and rates of solute loss: linkages between weathering and erosion in asmall, steep catchment. Geological Society of America Bulletin 114 (9), 1143–1158.

Bierman, P., Gillespie, A., Caffee, M., Elmore, D., 1995. Estimating erosion rates and ex-posure ages with 36Cl produced by neutron activation. Geochimica et Cosmochi-mica Acta 59 (18), 3779–3798.

Bosbach, D., Charlet, L., Bickmore, B., Hochella, M.F., 2000. The dissolution of hectorite:In-situ, real-time observations using atomic force microscopy. American Mineralo-gist 85, 1209–1216.

Brady, N.C., Weil, R.R., 2002. The Nature and Properties of Soils. Prentice Hall, UpperSaddle River, NJ. 960 pp.

Brantley, S.L., White, A.F., 2009. Approaches to modeling weathered regolith. Reviewsin Mineralogy and Geochemistry 70 (1), 435–484.

Brimhall, G., Dietrich, W.E., 1987. Constitutive mass balance relations between chemi-cal composition, volume, density, porosity, and strain in metasomatic hydrochemi-cal systems: results on weathering and pedogenisis. Geochimica et CosmochimicaActa 51, 567–587.

Brimhall, G.H., et al., 1991. Deformational mass transport and invasive processes in soilevolution. Science 255, 695–702.

Buol, S.W., Weed, S.B., 1991. Saprolite–soil transformations in the piedmont andmoun-tains of North Carolina. Geoderma 51, 15–28.

Buss, H.L., et al., 2005. The coupling of biological iron cycling and mineral weatheringduring saprolite formation, Luquillo Mountains, Puerto Rico. Geobiology 3 (4),247–260.

Buss, H.L., Mathur, R., White, A.F., Brantley, S.L., 2010. Phosphorus and iron cycling in deepsaprolite, Luquillo Mountains, Puerto Rico. Chemical Geology 269 (1–2), 52–61.

Chadwick, O.A., et al., 2003. The impact of climate on the biogeochemical functioning ofvolcanic soils. Chemical Geology 202 (3–4), 195–223.

Colman, S.M., Pierce, K.L., 1981. Weathering rinds on andesitic and basaltic stones as aQuaternary age indicator, western United States. Geological Survey ProfessionalPaper 1210. United States Government Printing Office, Washington, D.C., p. 41.

Das, A., Krishnaswami, S., Sarin, M.M., Pande, K., 2005. Chemical weathering in theKrishna Basin and Western Ghats of the Deccan Traps, India: rates of basalt weath-ering and their controls. Geochimica et Cosmochimica Acta 69 (8), 2067–2084.

Davidson, J.J., 2010. Geochemistry and Mineralogy of Diabase from Ironstone Ridge,South-central Pennsylvania: Implications for Magma Generation Conditions,Phase Assemblages, and Melt Properties. Dickinson College, Undergraduate seniorthesis Carlisle. 39 pp.

Davidson, J.J., Edwards, B., 2010. Preliminary geochemistry and mineralogy of diabasefrom Ironstone Ridge, central Pennsylvania: implications for magma storagedepths. Northeastern Section and Southeastern Section Joint Meeting of the Geo-logical Society of America, p. 158. Baltimore, MD.

Dessert, C., et al., 2001. Erosion of the Deccan Traps determined by river geochemistry:impact on the global climate and the 87Sr/86Sr ratio of seawater. Earth and Plane-tary Science Letters 188, 459–474.

Dessert, C., Dupre, B., Gaillardet, J., Francois, L.M., Allegre, C.J., 2003. Basalt weatheringlaws and the impact of basalt weathering on the global carbon cycle. Chemical Ge-ology 202, 257–273.

Egli, M., Fitze, P., 2000. Formulation of pedologic mass balance based on immobile ele-ments: a revision. Soil Science 165 (5), 437–443.

Eick, M.J., Grossl, P.R., Golden, D.C., Sparks, D.L., Ming, D.W., 1996a. Dissolution kineticsof a lunar glass simulant at 25 degrees C: the effect of pH and organic acids. Geo-chimica et Cosmochimica Acta 60 (1), 157–170.

Eick, M.J., Grossl, P.R., Golden, D.C., Sparks, D.L., Ming, D.W., 1996b. Dissolution of alunar basalt simulant as affected by pH and organic anions. Geoderma 74, 139–160.

Gaillardet, J., Millot, R., Dupré, B., 2003. Chemical denudation rates of thewestern Canadianorogenic belt: the Stikine terrane. Chemical Geology 201 (3–4), 257–279.

Gardner, L.R., Kheoruenromne, I., Chen, H.S., 1981. Geochemistry and mineralogy of anunusual diabase saprolite near Columbia, South Carolina. Clays and Clay Minerals29 (3), 184–190.

Gautier, J.-M., Oelkers, E.H., Schott, J., 2001. Are quartz dissolution rates proportional toB.E.T. surface areas? Geochimica et Cosmochimica Acta 65 (7), 1059–1070.

Gislason, S.R., Eugster, H.P., 1987. Meteoric water–basalt interactions. I: a laboratorystudy. Geochimica et Cosmochimica Acta 51 (10), 2827–2840.

Gislason, S.R., Oelkers, E.H., 2003.Mechanism, rates and consequences of basaltic glass dis-solution: II. An experimental study of the dissolution rates of basaltic glass as a func-tion of pH and temperature. Geochimica et Cosmochimica Acta 67 (20), 3817–3832.

Gislason, S.R., Arnorsson, S., Armannsson, H., 1996. Chemical weathering of basalt insouthwest Iceland: effects of runoff, age of rocks and vegetative/glacial cover.American Journal of Science 296 (8), 837–907.

Granger, D.E., Fabel, D., Palmer, A.N., 2001. Pliocene–Pleistocene incision of the GreenRiver, Kentucky, determined from radioactive decay of cosmogenic 26Al and 10Bein Mammoth Cave sediments. Geological Society of America Bulletin 113 (7),825–836.

Hamilton, J.P., Pantano, C.G., Brantley, S.L., 2000. Dissolution of albite glass and crystal.Geochimica et Cosmochimica Acta 64, 2603–2615.

100 E.M. Hausrath et al. / Chemical Geology 290 (2011) 89–100

Hancock, G., Kirwan, M., 2007. Summit erosion rates deduced from 10Be: implicationsfor relief production in the central Appalachians. Geology 35 (1), 89–92.

Hausrath, E.M., Brantley, S.L., 2010. Basalt and olivine dissolution under cold, salty, andacidic conditions: what can we learn about recent aqueous weathering on Mars?Journal of Geophysical Research 115 (E12), E12001.

Hausrath, E.M., Navarre-Sitchler, A.K., Sak, P.B., Steefel, C.I., Brantley, S.L., 2008a. Basaltweathering rates on Earth and the duration of liquid water on the plains of GusevCrater, Mars. Geology 36 (1), 67–70.

Hausrath, E.M., et al., 2008b. Short- and long-term olivine weathering in Svalbard: im-plications for Mars. Astrobiology 8 (6), 1079–1092.

Hausrath, E.M., Neaman, A., Brantley, S.L., 2009. Elemental release rates from dissolvingbasalt and granite with and without organic ligands. American Journal of Science309, 633–660.

Helgeson, H.C., Murphy, W.M., Aagard, P., 1984. Thermodynamic and kinetic con-straints on reaction rates amoung minerals and aqueous solutions, II. Rate con-stants, effective surface area, and the hydrolysis of feldspar. Geochimica etCosmochimica Acta 48, 2405–2432.

Herndon, E., Jin, L., Brantley, S.L., 2011. Soils reveal widespeard manganese enrichmentfrom industrial. Environmental Science and Technology 45 (1), 241–247.

Jin, L., et al., 2010. Mineral weathering and elemental transport during hillslope evolu-tion at the Susquehanna/Shale Hills Critical Zone Observatory. Geochimica et Cos-mochimica Acta 74 (13), 3669–3691.

Kuwahara, Y., 2006. In-situ AFM study of smectite dissolution under alkaline condi-tions at room temperature. American Mineralogist 91, 1142–1149.

Lebedeva, M.I., Fletcher, R.C., Brantley, S.L., 2010. A mathematical model of steady-stateregolith production at constant erosion rate. Earth Surface Processes and Land-forms 35 (5), 508–524.

Lichtner, P.C., 1988. The quasi-stationary state approximation to coupled mass transitand fluid–rock interaction in a porous medium. Gecochimica et CosmochimicaActa 52, 143–165.

Lindsay, W.L., Vlek, P.L.G., Chien, S.H., 1989. Phosphate minerals. In: Dixon, J.B., Weed,S.B. (Eds.), Minerals in Soil Environments. Soil Science Society of America, Madi-son, p. 1244.

Lloret, E., et al., 2011. Comparison of dissolved inorganic and organic carbon yields andfluxes in the watersheds of tropical volcanic islands, examples from Guadeloupe(French West Indies). Chemical Geology 280 (1–2), 65–78.

Louvat, P., Allegre, C.J., 1997. Present denudation rates at Reunion island determined byriver geochemistry: basalt weathering and mass budget between chemical andmechanical erosions. Geochimica et Cosmochimica Acta 61, 3645–3669.

Louvat, P., Allegre, C.J., 1998. Riverine erosion rates on Sao Miguel volcanic island,Azores archipelago. Chemical Geology 148, 177–200.

Maher, K., Steefel, C.I., White, A.F., Stonestrom, D.A., 2009. The role of reaction affinityand secondary minerals in regulating chemical weathering rates at the Santa CruzSoil Chronosequence, California. Geochimica et Cosmochimica Acta 73 (10),2804–2831.

Matter, J.M., et al., 2009. Permanent carbon dioxide storage into basalt: the CarbFixpilot project, Iceland. Energy Procedia 1 (1), 3641–3646.

McKenzie, R.M., 1989. Manganese oxides and hydroxides. In: DIxon, J.B., Weed, S.B.(Eds.), Minerals in Soil Environments. Soil Science Society of America, Madison,WI, pp. 439–466.

Metz, V., Raanan, H., Pieper, H., Bosbach, D., Ganor, J., 2005. Towards the establishmentof a reliable proxy for the reactive surface area of smectite. Geochimica et Cosmo-chimica Acta 69, 2581–2591.

Minyard, M.L., et al., 2011. Halloysite nanotubes and bacteria at the saprolite–bedrockinterface, Rio Icacos watershed, Puerto Rico. Soil Science Society of America Journal75 (2), 348–356.

Nagy, K.L., 1995. Dissolution and precipitation kinetics of sheet silicates. In: White, A.F.,Brantley, S.L. (Eds.), Chemical Weathering Rates of Silicate Minerals. Reviews inMineralogy. Mineralogical Society of America, Washington, D.C, pp. 173–225.

Navarre-Sitchler, A., Steefel, C.I., Yang, L., Tomutsa, L., Brantley, S.L., 2009. Evolution ofporosity and diffusivity associated with chemical weathering of a basalt clast. Jour-nal of Geophysical Research 114.

Neaman, A., Chorover, J., Brantley, S.L., 2005a. Element mobility patterns record organicligands in soils on early Earth. Geology 33 (2), 117–120.

Neaman, A., Chorover, J., Brantley, S.L., 2005b. Implications of the evolution of organicacid moieties for basalt weathering over geological time 10.2475/ajs.305.2.147.American Journal of Science 305 (2), 147–185.

Oelkers, E.H., Gislason, S.R., 2001. The mechanism, rates and consequences of basalticglass dissolution: I. An experimental study of the dissolution rates of basalticglass as a function of aqueous Al, Si, and oxalic acid concentrations at 25 °C andpH=3 and 11. Geochimica et Cosmochimica Acta 65 (21), 3671–3681.

Oguchi, C.T., Matsukura, Y., 1999. Effect of porosity on the increase in weathering-rindthicknesses of basaltic andesite gravel. Engineering Geology 55 (1–2), 77–89.

Oh, N.-H., Richter, D.D., 2005. Elemental translocation and loss from three highlyweathered soil-bedrock profiles in the southeastern United States. Geoderma126, 5–25.

Patino, L.C., Velbel, M.A., Price, J.R., Wade, J.A., 2003. Trace element mobility duringspheroidal weathering of basalts and andesites in Hawaii and Guatemala. ChemicalGeology 202, 343–364.

Pavich, M.J., Brown, L., Valette-Silver, J.N., Klein, J., Middleton, R., 1985. 10Be analysis ofa Quaternary weathering profile in the Virginia Piedmont. Geology 13 (1), 39–41.

Pavich, M.J., Leo, G.W., Obermeier, S.F., Estabrook, J.R., 1989. Investigations of the char-acteristics, origin, and residence time of the upland residual mantle of the Pied-mont of Fairfax County, Virginia. U.S. Geological Survey Professional Paper, 1352.58 pp.

Pelt, E., et al., 2008. Uranium–thorium chronometry of weathering rinds: rock alter-ation rate and paleo-isotopic record of weathering fluids. Earth and Planetary Sci-ence Letters 276 (1–2), 98–105.

Porter, S.C., 1975. Weathering rinds as a relative-age criterion: application to subdivi-sion of glacial deposits in the Cascade Range. Geology 3 (3), 101–104.

Potter Jr., N., Niemitz, J.W., Sak, P.B., 2009. Long-term field-based studies in teachinggeology. In: Whitmeyer, S.J., Mogk, D.W., Pyle, E.J. (Eds.), Field Geology Education:Historical Perspectives and Modern Approaches: Geological Society of AmericaSpecial Paper, 461. doi:10.1130/2009.2461(16.

Reuter, J.M., 2005. Erosion rates and patterns inferred from cosmogenic 10Be in the Sus-quehanna River Basin. University of Vermont. 160 pp.

Rice Jr., T.J., Buol, S.W., Weed, S.B., 1985. Soil–saprolite profiles derived from maficrocks in the North Carolina Piedmont: I. Chemical, morphological, and mineralog-ical characteristics and transformations. Soil Science Society of America Journal 49,171–178.

Sak, P.B., Fisher, D.M., Gardner, T.W., Murphy, K., Brantley, S.L., 2004. Rates of weather-ing rind formation on Costa Rican basalt. Geochimica et Cosmochimica Acta 68 (7),1453–1472.

Sak, P.B., et al., 2010. Controls on rind thickness on basaltic andesite clasts weatheringin Guadeloupe. Chemical Geology 276 (3–4), 129–143.

Small, E.E., Anderson, R.S., Repka, J.L., Finkel, R., 1997. Erosion rates of alpine bedrocksummit surfaces deduced from in situ 10Be and 26Al. Earth and Planetary ScienceLetters 150 (3–4), 413–425.

Steefel, C.I., Van Capellen, P., Nagy, K.L., Lasaga, A.C., 1990. Modeling water–rock inter-action in the surficial environment: the role of precursors, nucleation, and Ostwaldripening. Chemical Geology Geochemistry of the Earth's Surface and of MineralFormation 84 (1–4), 322–325.

Steefel, C.I., DePaolo, D.J., Lichtner, P.C., 2005. Reactive transport modeling: an essentialtool and a new research program for the Earth sciences. Earth and Planetary Sci-ence Letters 240, 539–558.

Sutter, J.F., 1988. Innovative approaches to the dating of igneous events in the earlyMesozoic basins of the Eastern United States. U. S. Geological Survey Bulletin Re-port: B 1776, 194–200.

Velbel, M.A., 1993. Constancy of silicate-mineral weathering-rate ratios between natu-ral and experimental weathering: implications for hydrologic control of differ-ences in absolute rates. Chemical Geology 105 (1–3), 89–99.

White, A.F., 1995. Chemical weathering rates of silicate minerals in soils. In: White, A.F.,Brantley, S.L. (Eds.), Chemical Weathering Rates of Silicate Minerals. Reviews inMineralogy. Mineralogical Society of America, Washington, D.C., pp. 407–461.

White, A.F., 2002. Determining mineral weathering rates based on solid and soluteweathering gradients and velocities: application to biotite weathering in sapro-lites. Chemical Geology 190 (1–4), 69–89.

White, W.B., 2009. The evolution of Appalachian fluviokarst: competition betweenstream erosion, cave development, surface denudation, and tectonic uplift. Journalof Cave and Karst Studies 71 (3), 159–167.

White, A.F., Brantley, S.L., 2003. The effect of time on the weathering of silicate min-erals: why do weathering rates differ in the laboratory and field? Chemical Geolo-gy 202 (3–4), 479–506.

White, A.F., Peterson, M.L., 1990. Role of reactive-surface-area characterization in geo-chemical kinetic models. In: Melchior, D.C., Bassett, R.L. (Eds.), Chemical Modelingof Aqueous Systems II. : ACS Symposium Series. American Chemical Society, Wash-ington, D.C., pp. 461–475.

Wolff-Boenisch, D., Gislason, S.R., Oelkers, E.H., 2006. The effect of crystallinity on dis-solution rates and CO2 consumption capacity of silicates. Geochimica et Cosmochi-mica Acta 70 (4), 858–870.

Yokoyama, S., Kuroda, M., Sato, T., 2005. Atomic force microscopy study of montmoril-lonite dissolution under highly alkaline conditions. Clays and Clay Minerals 53,147–154.