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Weathering of the Rio Blanco quartz diorite, Luquillo Mountains,Puerto Rico: Coupling oxidation, dissolution, and fracturing

Heather L. Buss a,*, Peter B. Sak b, Samuel M. Webb c, Susan L. Brantley d

a U.S. Geological Survey, 345 Middlefield Road, MS 420, Menlo Park, CA 94025, USAb Department of Geology, Dickinson College, Carlisle, PA 17013, USA

c Stanford Synchrotron Radiation Laboratory, Stanford Linear Accelerator Center, 2525 Sand Hill Road, MS 69, Menlo Park, CA 94025, USAd Earth and Environmental Systems Institute, Penn State University, University Park, PA 16802, USA

Received 13 September 2007; accepted in revised form 10 June 2008; available online 4 July 2008

Abstract

In the mountainous Rio Icacos watershed in northeastern Puerto Rico, quartz diorite bedrock weathers spheroidally, pro-ducing a 0.2–2 m thick zone of partially weathered rock layers (�2.5 cm thickness each) called rindlets, which form concentriclayers around corestones. Spheroidal fracturing has been modeled to occur when a weathering reaction with a positive DV ofreaction builds up elastic strain energy. The rates of spheroidal fracturing and saprolite formation are therefore controlled bythe rate of the weathering reaction.

Chemical, petrographic, and spectroscopic evidence demonstrates that biotite oxidation is the most likely fracture-inducing reaction. This reaction occurs with an expansion in d (001) from 10.0 to 10.5 A, forming ‘‘altered biotite”.Progressive biotite oxidation across the rindlet zone was inferred from thin sections and gradients in K and Fe(II). Usingthe gradient in Fe(II) and constraints based on cosmogenic age dates, we calculated a biotite oxidation reaction rate of8.2 � 10�14 mol biotite m�2 s�1. Biotite oxidation was documented within the bedrock corestone by synchrotron X-raymicroprobe fluorescence imaging and XANES. X-ray microprobe images of Fe(II) and Fe(III) at 2 lm resolution revealedthat oxidized zones within individual biotite crystals are the first evidence of alteration of the otherwise unalteredcorestone.

Fluids entering along fractures lead to the dissolution of plagioclase within the rindlet zone. Within 7 cm surroundingthe rindlet–saprolite interface, hornblende dissolves to completion at a rate of 6.3 � 10�13 mol hornblende m�2 s�1: thefastest reported rate of hornblende weathering in the field. This rate is consistent with laboratory-derived hornblende dis-solution rates. By revealing the coupling of these mineral weathering reactions to fracturing and porosity formation we areable to describe the process by which the quartz diorite bedrock disaggregates and forms saprolite. In the corestone, bio-tite oxidation induces spheroidal fracturing, facilitating the influx of fluids that react with other minerals, dissolving pla-gioclase and chlorite, creating additional porosity, and eventually dissolving hornblende and precipitating secondaryminerals. The thickness of the resultant saprolite is maintained at steady state by a positive feedback between the denu-dation rate and the weathering advance rate driven by the concentration of pore water O2 at the bedrock–saproliteinterface.� 2008 Elsevier Ltd. All rights reserved.

1. INTRODUCTION

Rock weathering is the initial step in the soil-formingprocess and the primary contributor of chemical solutesto the hydrosphere. The boundaries at which weatheringtransforms intact bedrock into disaggregated regolith

0016-7037/$ - see front matter � 2008 Elsevier Ltd. All rights reserved.

doi:10.1016/j.gca.2008.06.020

* Corresponding author.E-mail addresses: [email protected] (H.L. Buss), sakp@

dickinson.edu (P.B. Sak), [email protected] (S.M.Webb), [email protected] (S.L. Brantley).

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Geochimica et Cosmochimica Acta 72 (2008) 4488–4507


(such as saprolite) are complex interfaces controlled bycoupled chemical, physical, and microbial processesdriven by the flux of reactants into the bedrock. The rateof rock weathering determines the advance rate of thebedrock–saprolite interface and, in a steady state profile,the rate of formation of soil. In turn, the rates ofindividual geochemical reactions contribute to the bed-rock weathering rate, the flux of mineral nutrients tothe biosphere, and the flux of solutes to the hydro-sphere.

Soils developed on granite, sandstone, and metamor-phosed crystalline bedrock have often been presumed tobe in steady state with respect to thickness (e.g., Brownet al., 1995; Heimsath et al., 1999; von Blanckenburget al., 2004), regolith formation, erosion, weathering (e.g.,Pavich, 1989; Brown et al., 1995; Riebe et al., 2003; Turneret al., 2003; Dosseto et al., in press), elemental concentra-tions (e.g., Riebe et al., 2003), and yearly averaged biomassand organic carbon (Buss et al., 2005). If indeed the totalmass of regolith (including saprolite) in a profile is constantin time, then conservation of mass requires that the rate offormation of regolith equals the total denudation rate, andthus the sum of the rates of chemical weathering and phys-ical erosion. In an isovolumetric weathering profile (such asa saprolite), steady state mass is equivalent to steady statethickness. In the granitic Rio Icacos watershed in PuertoRico, Brown et al. (1995) estimated a long-term average to-tal denudation rate of 43 m Ma�1 ± 35% based on cosmo-genic 10Be accumulations in the stream sediments,consistent with later estimates by White et al. (1998) andRiebe et al. (2003) using combinations of cosmogenic nu-clide and geochemical mass balance methods. White et al.(1998) calculated the rate of regolith formation (also knownas the weathering advance rate) as 58 m Ma�1 based on thewatershed solute flux normalized to the geographical sur-face area (expressed as unit regolith surface area). Theagreement of these rates supports the assumption of a stea-dy state regolith profile. Thus, the total denudation rate iscommonly equated with the rate of bedrock transformationto regolith for the Rio Icacos watershed and other siteswhere weathering is assumed to be occurring at steady state(e.g., Murphy et al., 1998; White et al., 1998; Riebe et al.,2003; Turner et al., 2003; Buss et al., 2005; Fletcher et al.,2006).

Although steady state assumptions are frequently usedin models the coupling between erosion and weathering ad-vance rates is difficult to understand. Fletcher et al. (2006)proposed that the rate at which saprolite is created fromthe Rio Icacos bedrock is controlled by the chemistry ofpore fluids at depth and that this chemistry is in turn con-trolled by the thickness of the regolith profile. Pore fluidO2 may diffuse into the bedrock, oxidizing Fe(II)-bearingprimary minerals, creating elastic strain that spheroidallyfractures the Rio Icacos bedrock (Fletcher et al., 2006). Inmany locations such as Rio Icacos, bedrock disaggregatesto regolith at spheroidally weathering interfaces. Geochem-ical reactions involving Fe-silicate minerals have been cred-ited with causing the fracturing that initiates regolithformation at such interfaces (e.g., Larsen, 1948; Simpson,1964; Eggler et al., 1969; Isherwood and Street, 1976;

Chatterjee and Raymahashay, 1998; Scarciglia et al.,2005; Fletcher et al., 2006).

Fletcher et al. (2006) proposed that oxidation of horn-blende (where DVreaction > 0) drives fracturing and controlsthe rate of weathering advance. Given that dissolved O2 insoil pore waters decreases with depth, they pointed out thatdeeper soils would have slower weathering advance ratesthan shallower soils. Their model is consistent with porewater dissolved O2 as the parameter that couples erosionto the weathering advance rate. Herein we present a de-tailed geochemical investigation of the spheroidally weath-ering bedrock–saprolite interface in the Rio Icacoswatershed. Specifically we examine the occurrence and ratesof individual mineral weathering reactions within the sphe-roidally fractured zone, with particular emphasis on the Fe-silicate minerals, in order to identify the fracture-initiatingreaction, to understand the relationships between specificreactions and porosity development, and to quantify theflux of Fe(II) from weathering bedrock.

1.1. Spheroidal weathering

Spheroidal weathering refers to a well known but poorlyunderstood type of exfoliation in which corestones are sur-rounded by concentric layers of weathering rock (Ollier,1971). Corestones are blocks of residual bedrock that areweathered in place along pre-existing joint planes, roundingoff corners to create boulders (Linton, 1955; Ollier, 1971).The concentric layers surrounding corestones are variouslycalled sheets, layers, shells, plates, scales, or rindlets(Bisdom, 1967; Ollier, 1971; Fritz and Ragland, 1980).Envisioning an entire sequence of layers as a weatheringrind, here we adopt the term rindlets to describe the individ-ual layers consistent with terminology used by Fritz andRagland (1980), Turner et al. (2003), Buss et al. (2005),and Fletcher et al. (2006). We also refer to a complete setof rindlets as a rindlet zone (Figs. 1 and 2).

Spheroidal weathering is most common in homoge-neous, jointed, competent rocks, primarily granite and ba-salt although it has also been observed on gneisses,schists, andesites, sandstones, and greywackes (e.g., Ollier,1967; Heald et al., 1979; Fritz and Mohr, 1984; Chatterjeeand Raymahashay, 1998). Granitic corestones are foundin a variety of climatic regions on all continents except Ant-arctica and thus represent a widespread phenomenonimpacting regolith formation worldwide.

Large-scale exfoliation structures such as tors or domesin the Sierra Nevada have long been thought to be createdby the dilation induced by the relief of stress duringerosional unloading (e.g., Gilbert, 1904; Farmin, 1937;Schattner, 1961; Thomas, 1974), a mechanism thought tobe independent of—although aggravated by—surficial pro-cesses such as hydration and freeze–thaw (e.g., Farmin,1937). Spheroidal weathering has often been recognized asa different, or special, case of exfoliation and has beenattributed to surficial processes, most commonly chemicalweathering (Farmin, 1937; Ollier, 1971; Fletcher et al.,2006). It is unclear why some rocks weather spheroidallyand others do not. Some investigators have proposed thatthe percentage of biotite (Le Pera and Sorriso-Valvo,

Weathering of the Rio Blanco quartz diorite, Puerto Rico 4489


2000) or feldspar (Ferry, 1984) controls susceptibility tospheroidal weathering and others have argued for(Isherwood and Street, 1976; Sequeira Braga et al., 2002)or against (Ollier, 1988; Romani and Twidale, 1998;Mignon and Thomas, 2002) climatic controls. A numberof researchers have suggested that dilation during mineralweathering could induce spheroidal fracturing or fracture-induced disintegration (e.g., Larsen, 1948; Simpson, 1964;Eggler et al., 1969; Isherwood and Street, 1976; Begle, 1978;Bustin and Mathews, 1979; Chatterjee and Raymahashay,1998; Fletcher et al., 2006). Ollier (1967) argues that a netvolume change is unlikely because the rocks retain jointand mineral structure after spheroidal cracking. Indeed,

spheroidal weathering often produces saprolite, which is de-fined as an isovolumetric weathering product. Fletcher et al.(2006) demonstrated quantitatively that only a very smallnet volume change is required to generate a build-up ofelastic strain energy sufficient to produce the fractures thatdefine individual rindlets.

Fletcher et al. (2006) postulated that a positive DVreaction

(DVreaction =P

Vproducts �P

Vreactants) for the precipitationof ferric oxides after the oxidation of ferrous mineralsbuilds up elastic strain energy in the rock, which causesthe fracturing that demarcates rindlets. In this scenario,the rate of spheroidal weathering, and hence the rate of ad-vance of the bedrock–saprolite interface (the rindlet zone),is a function of the concentrations of ferrous minerals andoxygen, the reaction rate, the rate of transport, and themechanical properties of the rock. Fletcher et al. (2006)proposed that the dissolution of hornblende in a graniticrock coupled with a precipitation of ferrihydrite could yielda positive DVreaction. However, there are several possiblereactions involving ferrous minerals common to graniticrocks that could produce a ferric product: dissolution ofhornblende, biotite, or augite followed by precipitation ofa ferric (hydr)oxide mineral, or solid-phase oxidation ofFe(II) within hornblende, biotite, or augite. In the presentpaper we seek to identify the first weathering reaction thatoccurs in the corestones in order to refine the model pre-sented in Fletcher et al. (2006) and to identify controls onthe thickness of the rindlet zone and its transformation tosaprolite.

In a spheroidally weathering system, a rindlet sequencedefines the bedrock–saprolite interface, where bedrockchemically weathers and disaggregates to form saprolite(Fig. 2). For an assumed steady state system such as theRio Icacos weathering profile (Brown et al., 1995; Riebeet al., 2003; Fletcher et al., 2006), the rate of formation ofthe saprolite is equal to the movement of the boundary be-tween the saprolite and the bedrock and to the total denu-dation rate. This boundary moves downward into thebedrock as rindlets form during spheroidal weathering.

Fig. 1. (a) Photograph of the primary sampling site along aroadcut on Route 191 (see map, Fig. 3). Here the �1 m diametercorestone is surrounded by �50 cm zone of rindlets and �2 m ofsaprolite. (b) From Buss et al. (2004), close up photograph of thetop of the corestone pictured in (a). The sub-horizontal rindletsequence pictured was sampled in the present study.

Fig. 2. Simplified diagram of the weathering profile in the RioIcacos watershed. The quartz diorite bedrock is rounded intocorestones, which fracture concentrically forming a �0.2–2 m thickrindlet zone, overlain by 2–8 m of saprolite and 0.5–1 m of soil. Thearea immediately surrounding the rindlet–saprolite interface is herecalled the ‘‘protosaprolite zone”.

4490 H.L. Buss et al. / Geochimica et Cosmochimica Acta 72 (2008) 4488–4507


Turner et al. (2003) conceptualized the Rio Icacos weather-ing profile as multiple parallel weathering fronts that ad-vance inward towards the center of the corestones as wellas inward from inter-rindlet boundaries towards the rindletcenters. In this model, the net rate of advance of the multi-ple rindlet-scale fronts equals the net rate of advance of thecorestone-scale front. Likewise, the net rate of advance ofthe corestone-scale front equals the total denudation rate.If the rindlet-bounding fractures are a series of parallelweathering fronts (as per Turner et al., 2003), then chemicaland mineralogical trends should be measurable across a sin-gle rindlet. That is, weathering intensity should decreasemarkedly from a rindlet’s edges towards the rindlet’s inte-rior. In this case, the weathering advance rate would bethe sum of the rates of the individual parallel rindlet fronts.Alternatively, the entire rindlet zone could be a singleweathering front with only two controlling interfaces: thebedrock–rindlet interface and the rindlet–saprolite interface(Fig. 2).

1.2. Field site and sample collection

The Rio Icacos watershed in Puerto Rico’s LuquilloExperimental Forest is located on Rio Blanco quartz dio-rite bedrock mantled by 200 ka regolith, which was datedusing 10Be (Brown et al., 1995). The regolith is composedof 2–8 m of saprolite and 0.5–1 m of soil. These Picacho-Ciales complex soils, previously classified as ultisols(Boccheciamp et al., 1977), are now considered inceptisolsdue to weak B horizon development, despite being highlyweathered (USDA NCRS, 2002). The quartz dioriteweathers spheroidally forming corestones that average2 m in diameter, surrounded by 0.2–2 m rindlet zonescontaining individual rindlets of 2.6 cm average thickness(Turner et al., 2003; Buss et al., 2004, 2005; Fletcher etal., 2006).

The Rio Blanco quartz diorite bedrock contains zonedplagioclase, quartz, hornblende, partially chloritized bio-tite, and minor primary iron oxides and accessory miner-als (Murphy et al., 1998; White et al., 1998). Theoverlying saprolite contains kaolinite, goethite, quartz,and biotite. Biotite weathering occurs in two stages viatwo mechanisms. First, an ‘‘altered biotite” phase formsthat has a larger d (001) spacing, and higher Al/Si andFe(III)/Fe(II) ratios, and less Fe, Mg, and K than freshbiotite (Dong et al., 1998; Murphy et al., 1998). Thisphase was detected in the saprolite, but not in the freshrock (Dong et al., 1998). Next, altered biotite weathersto kaolinite in the saprolite either via epitaxial overgrowthof two layers of kaolinite onto one layer of altered biotiteor in a 1:1 ratio with the formation of an intermediate hal-loysite layer (Dong et al., 1998; Murphy et al., 1998).Within the rindlet zone, plagioclase, hornblende, and chlo-rite must weather completely because these minerals arenot found in the saprolite (Murphy et al., 1998; Whiteet al., 1998; Turner et al., 2003). In the present paperwe explore the weathering reactions occurring within therindlet zone and at all of the interfaces associated withthe rindlet zone: corestone–rindlet, rindlet–saprolite, andthe multiple parallel rindlet–rindlet interfaces.

In June 2003 and July 2004 we collected a suite of solidsamples from the Rio Icacos watershed including saprolite,corestone, and rindlets (Figs. 1 and 2). Rindlets were sam-pled from around a corestone exposed at a roadcut onRoute 191 (Fig. 3). The samples include the edge of a core-stone and a continuous 49 cm section of rindlets above thecorestone, spanning the entire rindlet sequence betweencorestone and saprolite. The corestone samples were col-lected using a sledgehammer. The rindlet samples werecarefully removed with a hammer and chisel, wrapped inplastic and packing tape, and labeled according to positionand orientation. At the rindlet–saprolite interface, the rind-lets are more friable, thinner, and lighter in color than themajority of the rindlet zone. Likewise, the saprolite justabove the interface is slightly harder and darker in colorthan the bulk of the overlying saprolite. This 7 cm thickzone, consisting of four �1 cm thick, highly weatheredrindlets and roughly 3 cm of adjacent saprolite, we termthe protosaprolite zone (Fig. 2). The rindlet–saprolite inter-face runs through the protosaprolite zone. In this outcrop,we refer to the rindlet zone as the �46 cm region containingall of the hard, cohesive rindlets between the corestone andthe protosaprolite zone. Rindlet-bounding fractures anasta-mose and outer rindlets (further from the corestone) tend tosplit, complicating precise definition of a single rindlet.Therefore, samples bounded by parallel fractures along achosen transect were defined as individual rindlets. Wherethe transect crossed anastamosing fractures or partiallysplit rindlets, the sample was collected as a single pieceand considered a single rindlet. Delicate saprolite and pro-tosaprolite samples were coated with melted wax on at leastone side before removal from the outcrop to preserve thestructure and orientation.

2. ANALYTICAL METHODS

2.1. Sample preparation and analysis

Bulk densities of the rindlet samples were determined bycoating samples with thin layers of rubber cement andimmersing them in water to measure volume displacement.Bulk elemental analysis was performed on pulverized andsieved (150 lm) corestone, rindlet, and saprolite samples.These analyses included major and minor elements byinductively coupled plasma atomic emission spectrometry(ICP-AES) after lithium metaborate fusion digestion andFeO by titration after a multi-acid digest (SGS MineralLaboratories, Ontario, Canada). Subsamples from withinindividual rindlets were taken by careful grinding with aDremel rotary tool equipped with diamond points.

Minerals within thin sections prepared from each rindletwere identified using optical microscopy and electron probemicroanalysis (EPMA, Cameca SX-50). EPMA was alsoused to measure the composition of crystals as a functionof distance from the corestone and from rindlet edges.Backscattered electron (BSE) images, mosaic maps ofBSE images, and X-ray elemental maps of the thin sectionswere made on an FEI Quanta 400 SEM equipped with anenergy dispersive X-ray spectrometer (EDS). Modal analy-sis was performed on backscattered images using XT-Docu



software (v. 3.2, Soft-Imaging System GmbH, Munster,Germany) to point count phases identified by the userbased on the backscatter grayscale value and crystal mor-phology. This technique has been used to measure porosityin weathering rinds (Dorn, 1995; Dixon et al., 2006). Tosupplement microscopic phase identification in the highlyweathered rindlets and in the protosaprolite zone, X-raydiffraction (XRD) was performed on randomly oriented,powdered samples on a Scintag X2 theta–theta goniometerwith a copper target. A step scan was used with a step sizeof 0.02� and 2 s per step.

2.2. X-ray microprobe imaging

To look for incipient oxidation reactions within Fe(II)-containing phases in the corestone that may be too smallto identify with the above mentioned techniques, synchro-tron micro-fluorescence maps of mineral grains in thin sec-tions were made at the Stanford Synchrotron RadiationLaboratory (SSRL) on Beam Line 2.3. This technique per-mits identification of microscale spatial variations in oxida-tion state and stoichiometry within single mineral grains.Beam size on the sample was approximately 2 � 2 lm at

Fig. 3. Maps of the field area. A star indicates location of the sampled road cut.



full width half maximum using Kirkpatrick-Baez (K-B)focusing optics (X-Radia). X-ray fluorescence data were re-corded using a single element Si Vortex detector (SSI).Monochromatic X-rays were selected using a water cooledSi(220) phi = 0 double crystal monochromator. The X-rayenergy was calibrated to the first inflection point of an Femetal foil, E = 7112 eV. Maps were collected at several inci-dent energies (7121, 7125, 7130, and 7142 eV) in continuousraster scanning mode in order to collect the fluorescence atseveral distinguishing points within the Fe edge to deter-mine Fe oxidation states in the sample. Fluorescence mapswere analyzed using the Microanalysis Toolkit (Webb,2006). Fe and K fluorescence counts were normalized tothe measured intensity of the incident X-ray beam (I0).The proportion of Fe(II) and Fe(III) at each point in themap was determined by linear least squares analysis basedon the normalized fluorescence spectra of Fe(II) (0.2 MFe(NH4)2(SO4)2 solution) and Fe(III) (Fe2O3) standards.X-ray absorption near edge spectra (XANES) were col-lected on the same beam line on spots of interest and micro-probe data were fitted to XANES spectra (Webb, 2006) toconfirm phase identification.

2.3. Mass transfer calculations

In a weathering profile, the mass of an immobile elementis, by definition, the same in both the parent rock and theweathered material for some volume of material. However,the mass fraction of the immobile element (e.g., g Ti g�1

saprolite) in the weathered material will differ from thatin the parent if weathering processes have resulted in achange in density. By comparing the mass fractions of mo-bile elements that of an immobile element the mass transfer(loss or gain of an element relative to the parent rock com-position) can be documented (e.g., Brimhall and Dietrich,1987; Anderson et al., 2002).

Mass transfer of individual elements relative to Ti wascalculated for each rindlet and saprolite sample. Open-sys-tem mass transport is quantifiable in the present samplesbecause the parent material is homogenous (Rio Blancoquartz diorite), of uniform age (Tertiary), and contains arelatively inert component (Ti) that is present in both theparent and product material (Chadwick et al., 1990). Be-cause saprolite is isovolumetrically weathered rock, volu-metric strain (change in volume as a result of stress),calculated on the basis of an assumed immobile element,should be near-zero. White et al. (1998) identified Ti as rel-atively immobile in the Puerto Rican saprolite relative tothe bedrock. The relative immobility of Ti in the rindletzone was confirmed by our calculations of very low volu-metric strain with respect to Ti:

eTi;w ¼qpCTi;p

qwCTi;w

� 1 ð1Þ

where eTi,w is the volumetric strain in the weathered (w)sample with respect to Ti, qw is the bulk density of theweathered material (rindlets), qp is bulk density of the par-ent (p) rock, CTi,w is the mass fraction of Ti in the weath-ered material, and CTi,p is the mass fraction of Ti in theparent rock. Ti occurs in the Rio Blanco quartz diorite par-

ent rock as ilmenite and as a structural component of bio-tite that is not lost from the biotite lattice during weathering(Murphy et al., 1998). Thus, we assume Ti to be immobilein order to estimate mass loss during weathering.

Mass transfer (gain or loss) of an element j relative to theparent rock, represented by the mass transfer coefficient,sTi,j, can be calculated as:

sTi;j ¼Cj;wCTi;p

Cj;pCTi;w

� �� 1 ð2Þ

Here, Cj,w is the mass fraction of element j in the weatheredmaterial, and Cj,p is the mass fraction of element j in theparent material. If sTi,j = �1, element j is entirely lost. IfsTi,j > 0, a net gain of element j relative to the parent rockis indicated. The average of five analyses from a single core-stone was used in the present calculations as the parent rockcomposition.

3. RESULTS

3.1. Porosity development

Three types of cracks were observed in the system andwill be referred to as macro-cracks, rindlet micro-cracks,and grain micro-cracks. The macro-cracks demarcate indi-vidual rindlets, run subparallel to the corestone–rindletinterface, and are easily observed in the outcrop (Fig. 1).Rindlet micro-cracks, found within every rindlet we exam-ined, are contained within individual rindlets, are not pref-erentially aligned, are not associated with any givenmineral, and cut across multiple crystals of all phases indis-criminately. Density, length, and width of rindlet micro-cracks increase with distance away from the corestone,towards the saprolite (Fig. 4). In the rindlets closest tothe corestone, apertures of rindlet micro-cracks are<10 lm wide and short (rarely extending across an entire6 mm2 SEM image). From about 20 cm above of the core-stone (about halfway between the corestone and the sapro-lite) and beyond, rindlet micro-cracks that are 25–90 lmwide are common, with some as large as 250 lm wide(Fig. 4c). By this point it becomes difficult to image a6 mm2 area that does not contain rindlet micro-cracks thatspan the image. Rindlet micro-cracks were not observedwithin the corestone.

Grain micro-cracks are associated with individual min-eral grains and are typically observed along cleavageplanes or along grain boundaries. Grain micro-cracksare found throughout all rindlets and are most noticeablewithin �5 mm of the macro-cracks (along the edges ofevery rindlet). Within the outer 9 or 10 mm of the core-stone, occasional grain micro-cracks are identifiable. Nomicro-cracks are recognized more than 10 mm into thecorestone.

The bulk density decreases steadily across the rindletzone, from 2.7 g cm�3 for the corestone to 1.8 g cm�3 atthe rindlet–saprolite interface (Fig. 5a), and mirrors the in-crease in porosity determined by point counting (Fig. 5b).Porosity and low molecular weight pore-filling precipitateswere readily distinguished from the crystalline phases usingEDS and SEM. To limit variability in the point counts due



Fig. 4. Backscattered electron (BSE) images of thin sections. (a) Large zoned plagioclase crystals dominate the un-fractured, apparently un-weathered corestone. (b) In a rindlet �8 cm above the corestone–rindlet interface, occasional zoned plagioclase crystals partially dissolve andform secondary precipitates within the interior of the crystals (shown with the thick black arrow), quartz crystals are pristine, and partiallychloritized biotite grains show no signs of dissolution or alteration in BSE images. (c) A rindlet sampled from �20 cm above the corestone istypical of most rindlets: all rindlets have rindlet micro-cracks (an example is shown with a white arrow) that cut across multiple crystals whilemost hornblende crystals remain visibly un-altered throughout the majority of the rindlet zone. (d) In the last intact rindlet below theprotosaprolite zone, �47 cm above the corestone–rindlet interface, chlorite is highly weathered, and hornblende, although cracked, shows noobvious dissolution features. The arrow points to a round hole within the hornblende where an apatite inclusion is believed to have dissolvedwithout affecting the adjoining hornblende (sub-spheroidal apatite inclusions are commonly observed in the unaltered corestone). Note thedifference in scale of this image. (e) In the protosaprolite zone, �2 cm below the rindlet–saprolite interface, plagioclase crystals are highlyweathered and hornblende crystals show the first signs of dissolution. Quartz remains visibly un-altered throughout the rindlet andprotosaprolite zones. (f) In the protosaprolite zone, �1 cm below the rindlet–saprolite interface, the plagioclase is almost entirely lost, much ofthe pore space created by dissolution of primary minerals is filled with gibbsite, and to a lesser extent, kaolinite. Here hornblende crystals aresignificantly smaller than they are in the corestone.



to large rindlet micro-cracks and image resolution, porosityand mineral phases in the majority of the rindlet zone wereonly counted on same-size BSE images (�6 mm2 areas) thatdid not contain rindlet micro-cracks that spanned the entireimage. However, the extensive network of micro-cracksmade this impossible in the outermost. For these samples,care was taken to avoid micro-cracks larger than �60 lmwide.

Porosity in the corestone is extremely low (0.03%) atapproximately 36 mm below the corestone–rindlet inter-face. Throughout the rindlet zone, non-crack porosity isdominated by intra-crystal weathering of the hydrother-mally altered calcic cores of zoned plagioclase crystals(Fig. 4b). Additional porosity is created by apatite crystals,which weather inward from the edges, forming apatite-shaped pores (Fig. 4d). Apatite crystals are mostly presentas inclusions within hornblende and biotite and less com-monly within primary iron oxides. They are also foundadjacent to other primary minerals. With increasing

distance from the corestone, apatite content decreases,while the abundance of apatite-shaped pores increase.

3.2. Chemical mobility

Calculations of the mass transfer coefficient, sTi,j (Eq.(2)), indicate the average order of cation mobility in the en-tire rindlet zone relative to the corestone:P > Ca � Na > Fe(II) > K > Mn > Si �Mg > Fe (tota-l) > Al, which differs from what has been calculated forthe saprolite: Na � Ca > Mg > Si > K > Al > Fe (total)(White et al., 1998). The most notable difference is the lowermobility of Mg in the rindlet zone relative to the saprolite.These results suggest rapid weathering of plagioclase andapatite and slower weathering of Fe-silicates (all of whichcontain Mg) in the rindlet zone, relative to the saprolitewhere Mg is lost during the weathering of biotite (Murphyet al., 1998; White et al., 1998).

Mass transfer coefficients also reveal most of the rindletsto be only slightly depleted (or enriched) in most elementsrelative to the corestone. The largest bulk chemical differ-ences relative to the corestone are found in the 7 cm thickprotosaprolite zone (Table 1) indicating that the majorityof the chemical weathering occurs over a very narrow zone.These chemical differences correlate with the relative degreeof weathering observable in BSE images (Fig. 4). Theseimages reveal little difference between the rindlet samplesand the corestone, with the exceptions of micro-cracks andplagioclase weathering in the rindlet samples. Within theprotosaprolite zone, however, intense weathering is evident.

3.3. Mineralogy

The average mineralogical composition of the corestonedetermined by digital point counting is given in Table 2 andis consistent with previous work on the Rio Blanco quartzdiorite bedrock (Seiders, 1971; Murphy et al., 1998; Whiteet al., 1998; Turner et al., 2003). Large (�200–1000 lm) pla-gioclase crystals with altered calcic cores are the most abun-dant mineral, comprising �50 vol %, followed by quartz at�20 vol %. The altered plagioclase cores are sericitized, typ-ical of hydrothermal alteration. Fe-silicate phases includehornblende; biotite, which is partially replaced by chlorite;and some other chlorite that does not have a biotite mor-phology. The chlorite is typical of low-grade metamorphicor hydrothermal alteration. Hornblende is the most abun-dant Fe-silicate phase, followed by biotite. No pyroxenewas observed and chlorite was relatively minor. Accessoryminerals include primary Fe oxides with ilmenite exsolutionlamellae, sphene, apatite, and zircon. Mineral formulas forplagioclase and the major Fe-silicate phases were deter-mined by EPMA (Table 3).

With the exception of a slight increase in grain micro-cracking along rindlet edges, no trends in mineralogy,chemistry, or porosity were recognized across individualrindlets (e.g., Fig. 7), contradictory to the Turner et al.(2003) model of a series of multiple parallel weatheringfronts made up of individual rindlets. Also, with the nota-ble exception of apatite and, to a lesser extent, plagioclase,no distinct trend was noted in the distribution of mineral

0

10

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30

40

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1.8 2.0 2.2 2.4 2.6 2.8

RindletsSaprolite

Density (g/cm3)

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)

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a

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0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15

Corestone

Volume Percent

Dis

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(cm

)

Open Porosity Open + Filled Porosity

Saprolite

Rindlets

b

Fig. 5. (a) Individual rindlets show a steady decrease in bulkdensity as a function of distance from the corestone. (b) Volumepercent open porosity and total porosity (including low molecularweight secondary precipitates such as gibbsite filling pore spaces)determined by point counting using XT Docu software onbackscattered electron images.



phases throughout the intact rindlet zone. Throughout themost pristine half (�25 cm) of the rindlet zone, precipitatedweathering products are rare and mostly confined to plagio-clase cores and along some micro-cracks. Hornblende ap-pears pristine throughout most of the rindlet zone.Specifically, no reddening due to oxidation, no dissolutionfeatures, and no alteration to secondary phases was ob-served. In fact, within the majority of the rindlet zone, pla-gioclase and apatite inclusions commonly weather outleaving the surrounding hornblende essentially unchangedand rindlet micro-cracks commonly cross-cut hornblendecrystals without evidence of chemical alteration or dissolu-tion along the edges (Fig. 4d).

In contrast, hornblende alteration is occasionally visibleabout 6 cm below the rindlet–saprolite interface (in the lasttwo intact rindlets below the protosaprolite zone). Herehornblende begins to look more reddish and some zones

Table 1Mass transfer (sTi,j) of elements across the rindlet and protosaprolite zones

Distancea Al Ca Feb K Mg Mn Na P Si Fe(II)

Rindlet zone

0.00 0.03 0.02 �0.02 0.09 �0.02 �0.04 0.03 �0.05 0.01 �0.010.70 �0.01 0.05 �0.05 �0.38 0.05 0.02 �0.05 �0.17 0.08 �0.135.50 0.79 0.58 0.02 0.68 0.17 0.04 0.88 0.05 0.59 �0.099.00 0.04 �0.01 0.02 �0.02 0.06 0.07 �0.02 �0.01 0.06 �0.03

14.50 �0.07 �0.23 �0.03 �0.19 �0.11 �0.05 �0.14 �0.15 �0.05 �0.2321.50 �0.03 �0.23 0.02 �0.13 �0.05 0.02 �0.15 �0.21 �0.03 �0.2130.00 �0.36 �0.55 �0.05 �0.20 �0.10 �0.12 �0.53 �0.64 �0.29 �0.3039.00 �0.36 �0.58 �0.11 �0.25 �0.12 �0.16 �0.57 �0.44 �0.38 �0.3345.00 �0.30 �0.57 �0.16 �0.16 �0.10 �0.18 �0.55 �0.37 �0.37 �0.3647.35 �0.08 �0.08 0.02 �0.24 0.24 0.18 �0.15 �0.35 �0.01 �0.0247.95 0.10 �0.15 �0.06 �0.13 0.02 0.01 �0.05 �0.39 0.04 �0.1948.00 �0.11 �0.39 �0.03 �0.17 0.04 0.02 �0.38 �0.37 �0.20 �0.24

Protosaprolite

48.45 0.44 �0.19 �0.02 �0.21 0.07 0.00 �0.05 �0.41 0.11 �0.1448.75 �0.09 �0.74 �0.12 �0.29 �0.09 �0.20 �0.75 �0.82 �0.44 �0.3949.40 �0.48 �0.81 �0.07 �0.48 �0.20 �0.29 �0.87 �0.90 �0.47 �0.4150.20 �0.40 �0.88 �0.16 �0.60 �0.38 �0.45 �0.94 �0.95 �0.50 �0.5751.00 �0.23 �0.90 �0.16 �0.70 �0.49 �0.52 �0.94 �0.95 �0.42 �0.6552.00 2.13 �0.98 0.33 �0.85 �0.75 �0.36 �0.96 �1.00 0.35 �0.89

a Distance in cm from the corestone–rindlet interface.b Fe is total iron.

Table 2Average mineralogical compositiona of the corestone

Phase Volume %

Porosity 0.03Filled porosityb 0.88Quartz 19.9Plagioclase 49.3Fe-silicates 24.0Fe, Ti oxides 2.5Apatite 0.62

a Volume % composition determined by point counting back-scattered electron images of thin sections using image analysissoftware.

b Filled porosity is very fine-grained low molecular weightmaterial found associated with cracks or pore space. In thecorestone this material is dominantly pre-existing hydrothermalalteration products rather than weathering products.

Table 3Mineral formulas measured by electron microprobe

Mineral Formula Sample

Hornblende ðCa1:73Na0:29K0:06ÞðMg2:64Fe2þ1:95Mn0:09Ti0:12Al0:37ÞðSi7:24Al0:76ÞO22ðOHÞ2 Corestone

Hornblende ðCa1:57Na0:26K0:06ÞðMg2:66Fe2þ1:91Mn0:10Ti0:10Al0:47ÞðSi7:39Al0:61ÞO22ðOHÞ2 Middle rindleta

Hornblende ðCa1:76Na0:29K0:07ÞðMg2:66Fe2þ1:99Mn0:09Ti0:11Al0:33ÞðSi7:21Al0:79ÞO22ðOHÞ2 Protosaprolite

Biotite K0:89ðFe2þ1:23Fe3þ

0:05Mg1:19Mn0:02Ti0:18Al0:16ÞðSi2:88Al1:12ÞO10ðOHÞ2 Corestoneb

Biotite K0:73ðFe2þ1:01Fe3þ

0:15Mg1:16Mn0:02Ti0:18Al0:28ÞðSi2:88Al1:12ÞO10ðOHÞ2 Middle rindletBiotite K0:33ðFe2þ

0:87Fe3þ0:37Mg1:06Mn0:014Ti0:18Al0:35ÞðSi2:87Al1:13ÞO10ðOHÞ2 Protosaprolite

Chloritec ðMg3:32Fe2þ0:29Fe3þ

1:00Al1:39ÞðSi1:61Al2:39ÞO10ðOHÞ8 CorestonePlagioclase Na0.50Ca0.48K0.01Al1.43 Si2.56O8 Corestone

a The middle rindlet is located �24 cm above the corestone and �24 cm below the protosaprolite.b Biotite formulas were calculated assuming that of the total Fe, 3.7% is Fe(III) for the corestone (Murphy et al., 1998), 13% for the middle

rindlet, and 30% for the protosaprolite.c After Turner et al. (2003).



of Fe enrichment can be detected although dissolution fea-tures (e.g., etching) are not seen until the protosaprolitezone. In contrast to the hornblende crystals in the corestoneand in the rindlet zone, hornblende crystals in the proto-saprolite zone are significantly smaller and dramaticallyetched (Fig. 6). Similar saw-toothed etching along weath-ered hornblende cleavage planes has been reported by oth-ers (e.g., Berner and Schott, 1982; Anand and Gilkes, 1984;Velbel, 1989). Hornblende grains are extremely difficult tofind under SEM above the rindlet–saprolite interface,although XRD patterns indicate that hornblende persistsin the first 2 cm of the saprolite, but not beyond. EPMAanalyses (2–5 lm spot size) of >300 hornblende grains re-vealed no compositional changes as a function of distancefrom the corestone (Fig. 7).

As mentioned before, the rindlet–saprolite interface runsthrough the middle of the protosaprolite zone (Fig. 2). Inthe rindlet portion of the protosaprolite zone, primary min-eral grains are surrounded by abundant precipitated gibb-site with minor kaolinite identified by elemental ratiosmeasured by EDS and EPMA. Kaolinite becomes moreabundant above the saprolite–rindlet boundary. Boxworkreplacement of primary iron oxide crystals by iron(oxy)hydroxides (e.g., Nahon, 1986; Velbel, 1989; Delvigne,1998) was observed only within the protosaprolite zone.Quartz remains relatively pristine throughout the rindletand the protosaprolite zones (Fig. 4). Specifically, no disso-lution channels, no etch pits, and few grain micro-crackswere observed in quartz grains.

Point counting of mineral phases in BSE images usingimage analysis software was difficult for individual Fe-sili-cate phases due to overlapping grayscale values, whichare a function of the mean atomic number of the sample.Specifically, it was not possible to quantify the Fe-silicateminerals individually due to the partial hydrothermal, pseu-domorphic replacement of biotite by chlorite and the simi-lar chemistries (and hence similar average atomic numbersand grayscale values) of the chlorite, biotite, and horn-blende in these samples. The volume percentages of theseFe-silicates (as a group) indicates that they are relativelyconstant over the rindlet zone, but decrease in abundance

within the protosaprolite zone (Fig. 8). Apatite, whichwas readily point-counted (�0.6 vol % in the corestone),is entirely lost within the rindlet zone.

Throughout the rindlet zone, the weathering of plagio-clase from the calcic cores outward to the sodic rims (Fig.4b and c) and the oxidation of biotite are readily observedin thin sections under a petrographic microscope. Biotitegrains appear more reddish in all of the rindlets than inthe corestone, indicating oxidation. In the corestone, no evi-dence was found to indicate plagioclase weathering (distinctfrom pre-existing hydrothermal alteration). Conversely,zones of oxidized iron within corestone biotite crystals weredocumented in X-ray microprobe images (Fig. 9a) of biotitecrystals 2.7 cm from the corestone–rindlet interface. Linearleast square fitting of the Fe fluorescence counts at each en-ergy measured at each point in the image map was per-formed to determine the Fe(II) and Fe(III) content of thesample (Fig. 9). Fe(III)–K correlation plots show that areasof higher amounts of Fe(III) have lower K counts (Fig. 10).

Fig. 6. (a) Backscattered electron images of the corestone reveal pristine hornblende crystals. (b) Hornblende crystals in the protosaprolitezone contain dissolution features along cleavage planes and clumps of gibbsite and some kaolinite surround the crystals.

-40

-20

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60

80

100

120

140

160

0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0

Fe/Si Mg/Si

Dis

tanc

e fr

om C

ores

tone

-Rin

dlet

Inte

rfac

e (m

m)

x/Si Ratio of Hornblende Crystals

Fig. 7. Fe/Si and Mg/Si ratios of hornblende crystal grainsmeasured by EPMA as a function of distance from the corestonereveal no change in composition across the rindlet zone or acrossindividual rindlets. The negative values on the y axis refer tosamples within the corestone.



Masking of these individual areas show that the center of thebiotite grains have approximately constant ratios ofFe(III):K, while the altered regions are localized to the lowerK/higher Fe(III) regions (Fig. 11). XANES spectra of thesezones were consistent with oxidized biotite and were notconsistent with Fe(III)-(hydr)oxide phases nor hornblende(Fig. 12). Oxidized biotite could not be definitively distin-guished from chlorite on the basis of XANES spectra alone,but the two phases were easily differentiated by their relativeratios of K to total Fe, which are relatively high in biotite,lower in the oxidized zones, and close to zero in chlorite.Thus, the first weathering reaction observed in this systemis the oxidation of biotite within the corestone.

4. DISCUSSION

Based on our microscopic, spectroscopic, chemical, anddiffraction data, the following sequence of weathering reac-tions is proposed. First, within the corestone, diffusion ofoxygen into the fresh rock leads to oxidation of Fe(II) to

RindletsDis

tanc

e fro

m C

ores

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(cm

)

Fe-silicates Plagioclase

Saprolite

Corestone

A

RindletsDis

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(cm

)

Fe-silicates Plagioclase

Saprolite

Corestone0

10

20

30

40

50

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

Rindlets

Fe-silicates

Volume percent

Plagioclase

Saprolite

Corestone

A

Fig. 8. Volume percent composition of Fe-silicates (horn-blende + biotite + chlorite) and of plagioclase with distance fromthe corestone. Modal analysis was performed by point countingusing XT Docu software on backscattered electron images.

Fig. 9. X-ray microprobe fluorescence maps showing (a) Fe(III) and (b) Fe(II) content in a biotite crystal within the corestone, 2.7 cm fromthe corestone–rindlet interface. The color scale runs from blue to red indicating least to most Fe(III) or Fe(II). (For interpretation of thereferences to color in this figure legend, the reader is referred to the web version of this paper.)



Fe(III) within the biotite lattice, which loses K+ ions fromthe interlayer to maintain charge balance:

K0:89(Fe2þ1:23Fe3þ

0:05Mg1:19Mn0:02Ti0:18Al0:16)(Si2:88Al1:12)

O10ðOHÞ� þ 0:36Hþ þ 0:09O2 ! 0:18H2O

þ 0:36Kþ þ K0:53ðFe2þ0:87Fe3þ

0:41Mg1:19Mn0:02

Ti0:18Al0:16ÞðSi2:88Al1:12ÞO10ðOHÞ�2 ð3Þ

This reaction creates a form of biotite referred to as ‘‘al-tered biotite” by Dong et al. (1998), which is characterizedby an expansion of the (001) d-spacing from 10 to 10.5 A aswell as by the oxidation of lattice-bound Fe(II). In the lowporosity corestone, this expansion builds up elastic strainenergy that leads to fracturing, forming a macro-crackdemarcating an individual rindlet. The macro-cracks actas conduits for fluids that penetrate into the rindlets alongthe macro-crack boundaries to form grain micro-cracks,which tend to congregate along rindlet boundaries and atthe corestone–rindlet interface. Rindlet micro-cracks, whichfracture multiple mineral grains of multiple phases without

Fig. 11. X-ray microprobe fluorescence maps showing Fe(III) and K content of the biotite grain and altered biotite zones indicated in Fig. 10.(a) Fe(III) and (b) K content in the bulk biotite crystal. (c) Fe(III) and (d) K content of the altered biotite zones. The color scale runs fromblue to red indicating least to most Fe(III) or K. (For interpretation of the references to color in this figure legend, the reader is referred to theweb version of this paper.)

0 500 1000 1500 2000 25000

50

100

150

200

250

300

Fe(III) Counts

K C

ount

s

Altered Biotite

Biotite grain

Biotite edges

Fig. 10. A plot of K versus Fe(III) fluorescence counts for the biotitegrain shown in Fig. 9 shows that areas of higher Fe(III) have lowerK. The linear trend of points from the origin to the cluster of pointsindicating the biotite grain represents a constant ratio in Fe(III):Kthat is due to thinning along the edges of the biotite crystal. Note thatthe Fe(III):K ratios in the altered biotite region deviate significantlyfrom the trend in Fe(III):K ratios along the crystal edges and thuscannot be attributed simply to variations in thickness.

7100 7120 7140

0.0

0.5

1.0

1.5

Biot Fe2 peak Biotite standard Goethite standard Hematite standard Biotite sample, bulk Biotite sample, Fe3 zone

μ(E

)

Biot Fe3 peak

Energy (eV)

Fig. 12. X-ray absorption near edge spectra (XANES) of fresh andoxidized zones of a single biotite grain within a corestone thinsection compared to XANES of mineral standards.



following grain boundaries, also form after macro-cracking(i.e., they are not observed in the corestone), although theexact cause of these micro-cracks remains unknown.Regardless, rindlet micro-cracks permit penetration of flu-ids into the rindlet interiors, facilitating further biotite oxi-dation as well as plagioclase dissolution:

Na0:50Ca0:48K0:01Al1:43 Si2:56O8 + 5.76 H2CO3 + 2.24 H2O

! þ 0:50! Naþ þ 0:48Ca2þ þ 0:01Kþ þ 1:43Al3þ

þ 2:56H4SiO4 þ 5:76HCO3�� ð4Þ

As these reactions progress across the rindlet zone, somesaprolitization reactions occur: complete loss of chloriteand oxidation of biotite to its saprolite composition. Withinthe protosaprolite zone, the final transformation of rela-tively hard, intact rindlets into soft, disaggregated saprolitetakes place. This transformation is accomplished as theremaining plagioclase is lost, all of the hornblende dissolvesto completion:

(Ca1:73Na0:29K0:06)(Mg2:64 Fe2þ1:95Mn0:09Ti0:12Al0:37)

ðSi7:24Al0:76ÞO22ðOHÞ2 þ 16:56H2CO3 þ 5:2H2O

! 7:24H4SiO4 þ 1:13Al3þ þ 0:06Kþ þ 0:29Naþ

þ 1:73Ca2þ þ 2:64Mg2þ þ 1:95Fe2þ

þ 0:09Mn2þ þ 0:12TiO2 þ 16:56HCO3�

þ 0:12O2 ð5Þ

and goethite begins to form. Although hornblende dissolu-tion is insignificant within the rindlet zone, the absence ofhornblende in the saprolite indicates that complete dissolu-tion must occur over the �7 cm protosaprolite zone (Fig.6b). Although hornblende is also oxidized within the proto-saprolite zone, this signal is eclipsed by the complete dissolu-tion of the mineral. Throughout the saprolite, altered biotiteweathers to kaolinite (Murphy et al., 1998) and quartz dis-solves, developing etch pits (Schulz and White, 1999).

4.1. Weathering gradients and reaction stoichiometry

In a steady state weathering profile, gradients in bulkchemical composition relative to depth are generated by theweathering reactions taking place. Negative gradients (onplots of depth versus chemical composition) indicate a lossof a given element or mineral, while positive slopes indicatean increase. Here we present solid state elemental weatheringgradients (bs in m kg mol�1) following White (2002) fromwhich we extract information about the various weatheringreactions occurring over the spheroidally weathering profile.

The mass fraction (e.g., g element g�1 rock) of a mobileelement that is not incorporated into a secondary mineralafter release from a weathering primary mineral, can bemodeled as a linear decrease from an initial mass fractionC0 at depth z1 to Cw at a shallower depth z0 (Fig. 13). HereC0 corresponds to the mass fraction of the element withinthe protolith (the corestone) and Cw is the mass fractionof the element in a weathered sample at some depth z. Thevalue Cw may differ from the measured value Cj,w (Eq. (2))due to changes in density during weathering. These changesare estimated by normalizing the mass relative to an elementthat is relatively inert to chemical weathering, here Ti:

Cw ¼ Cj;wCTi;p

CTi;w

� �ð6Þ

where CTi,p (mol kg�1) is the mass fraction of the assumedinert element (Ti) in the protolith and CTi,w is the massfraction of Ti in the weathered material. This normalizationis based on the same principles as the calculation of themass transfer coefficient (Eq. (2)) and can be easilyconverted:

sTi;j ¼Cw

Cj;p� 1 ð7Þ

Values of Cw with depth for several elements are shown inFig. 14 with linear fits. Within the profile studied here,two discrete weathering gradients can be discerned for sev-eral elements (Fe, K, Mg, and Mn): the first across therindlet zone, the second through the protosaprolite zone(Fig. 2). Assuming a steady state profile, differences in gra-dient for a single element in the protosaprolite versus inthe rindlet zone indicate a change in the rate of releaseof that element, which can signify a change in mineralweathering rates and/or a change in the reactions takingplace. For example, the weathering gradient in Fe(II) issignificantly shallower across the protosaprolite zone(bs = 0.088 m kg mol�1) than in the rindlet zone(bs = 2.02 m kg mol�1, Fig. 14a) indicating a more rapidloss of Fe(II) in the protosaprolite and potentially a differ-ent reaction. Because the weathering gradients in Fe(III)across the rindlet and protosaprolite zones (Fig. 14b) arereversed with respect to the gradients in Fe(II) (Fig.14a), it is clear that Fe(II) is lost via oxidation to Fe(III),rather than by removal from the system.

In a multi-mineralic system, multiple weathering reac-tions may occur simultaneously, each contributing to theelemental weathering gradients in different proportions.The loss of Mg (Fig. 14c) in the rindlet and protosaprolitezones could result from the dissolution of biotite, horn-

Fig. 13. Schematic diagram (after White, 2002) depicting thedistributions of a mobile element in a weathering profile.



blende, or chlorite, or the oxidation of hornblende. The lossof K (Fig. 14d) could be caused by biotite oxidation or dis-solution of biotite or hornblende; and the loss of Mn (Fig.14f) could indicate dissolution of hornblende or biotite. Theloss of Fe(II) over the rindlet zone could be caused by anyof the following reactions: dissolution of hornblende, oxi-

dation of Fe(II) in hornblende, dissolution of biotite, oxida-tion of Fe(II) in biotite, dissolution of chlorite, or oxidationof Fe(II) in chlorite. These reactions involve Fe2+ as well asK+ and/or Mg2+. We can represent these reactions by theirstoichiometries with respect to these ions. The stoichiome-try of biotite oxidation is 1 mol Fe2+ oxidized per mol

0.7

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Saprolite"Protosaprolite"

Rindlets

Corestone

Na

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Mg


Rindlets

Corestone

Concentration Cw (mol kg-1)

Prof

ile D

epth

(m)

0.7

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Saprolite

"Protosaprolite"

Rindlets

Corestone

Fe(II)

Prof

ile D

epth

(m)

0.7

0.6

0.5

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0.3

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RindletsCorestone

K

0.7

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0.2

0.1

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0.00 0.01 0.02 0.03 0.04 0.05 0.06

Saprolite

"Protosaprolite"

Rindlets

Corestone

Mn


A

B

C F

0.7

0.6

0.5

0.4

0.3

0.2

0.1

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0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0

E

D


Rindlets

Corestone

Fe(III)

Prof

ile D

epth

(m)

0.7

0.6

0.5

0.4

0.3

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0.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6 1.8 2.0


Rindlets

Corestone

Na

0.7

0.6

0.5

0.4

0.3

0.2

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0.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4

Mg


Rindlets

Corestone


Prof

ile D

epth

(m)

0.7

0.6

0.5

0.4

0.3

0.2

0.1

0.0

0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9

Saprolite

"Protosaprolite"

Rindlets

Corestone

Fe(II)

Prof

ile D

epth

(m)

0.7

0.6

0.5

0.4

0.3

0.2

0.1

0.0

0.00 0.05 0.10 0.15 0.20 0.25 0.30 0.35 0.40


RindletsCorestone

K

0.7

0.6

0.5

0.4

0.3

0.2

0.1

0.0

0.00 0.01 0.02 0.03 0.04 0.05 0.06

Saprolite

"Protosaprolite"

Rindlets

Corestone

Mn


A

B

C F

0.7

0.6

0.5

0.4

0.3

0.2

0.1

0.0

0.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6 1.8 2.0

E

D


Rindlets

Corestone

Na

0.7

0.6

0.5

0.4

0.3

0.2

0.1

0.0

0.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4

Mg


Rindlets

Corestone


Prof

ile D

epth

(m)

0.7

0.6

0.5

0.4

0.3

0.2

0.1

0.0

0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9

Saprolite

"Protosaprolite"

Rindlets

Corestone

0.7

0.6

0.5

0.4

0.3

0.2

0.1

0.0

0.00 0.05 0.10 0.15 0.20 0.25 0.30 0.35 0.40


RindletsCorestone

Fe(II) K

Prof

ile D

epth

(m)

0.7

0.6

0.5

0.4

0.3

0.2

0.1

0.0

0.00 0.01 0.02 0.03 0.04 0.05 0.06

Saprolite

"Protosaprolite"

Rindlets

Corestone

Mn


Aa Dd

B

C F

0.7

0.6

0.5

0.4

0.3

0.2

0.1

0.0

0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0


Rindlets

Corestone

Fe(III)

Prof

ile D

epth

(m)

Eb e

c f

Fig. 14. Weathering gradients (bs) for (a) Fe(II), (b) Fe(III), (c) Mg, (d) K, (e) Na, and (f) Mn, reveal two distinct trends for most elements:one through the rindlet zone, and one through the protosaprolite zone. Enrichments relative to the parent values seen as positive excursions atthe protosaprolite zone were not included in the linear fits. These excursions are thought to be related to the horizontality of the particularrindlet sequence studied combined with the difference in permeability between the saprolite and the rindlet zone. This possibility should befurther examined by analyzing sub-vertical rindlet sets from the sides of the corestones. For zones that were subsampled more intensively(protosaprolite, the corestone, and the first 3 rindlets), subsamples from single rindlets were averaged together.



K+ lost and the stoichiometry of hornblende oxidation is 2mol Fe2+ oxidized per mol Mg2+ lost.

To estimate the relative importance of the different Fe-silicate weathering reactions occurring within the rindletzone, we set up and solve a system of equations describingthe weathering gradients for several elements (Fe(II), Mg,K, and Mn) as functions of the stoichiometries of the pos-sible reactions:

1

bFeðIIÞ¼ mFeðIIÞ

b;o Mb;o þ mFeðIIÞh;o Mh;o þ mFeðIIÞ

b;d Mb;d þ mFeðIIÞh;d Mh;d

þ mFeðIIÞc;d M c;d ð8Þ

1

bK

¼ mKb;oMb;o þ mK

b;dMb;d þ mKh;dMh;d ð9Þ

1

bMg

¼ mMgh;o Mh;o þ mMg

b;d Mb;d þ mMgh;d Mh;d þ mMg

c;d M c;d ð10Þ

1

bMn

¼ mMnb;d Mb;d þ mMn

h;d Mh;d ð11Þ

Here bs is a weathering gradient (m kg mol�1) wheres = Fe(II), K, Mg, or Mn and mi;j

x represent the stoichiom-etric coefficients for x = Fe(II), K, Mg, and Mn, and sub-scripts i = b, h, or c indicating biotite, hornblende,chlorite, respectively, and j = o or d indicating oxidationor dissolution, respectively. For example, mFeðIIÞ

b;o indicatesthe moles of Fe(II) ‘‘lost” per mol of biotite oxidized. Thesevalues are listed in Table 4. M indicates the quantity(mol m�1 kg�1) of mineral (b, h, or c) undergoing eitheroxidation or dissolution (o or d) in the rindlet zone. We

reduce the unknowns from 5 to 4 by assuming that all chlo-rite is dissolved over the rindlet zone. This is reasonable be-cause chlorite—the least abundant Fe-silicate phase in thebedrock—is not observed in the protosaprolite or saprolite.This assumption gives Mc,d = 0.185 mol m�1 kg�1 of chlo-rite dissolved. Solving the system of equations yieldsMb,d = �1.0 mol m�1 kg�1, Mh,d = 0.39 mol m�1 kg�1,Mb,o = 2.8 mol m�1 kg�1, Mh,o = �0.01 mol m�1 kg�1.The negative values for biotite dissolution (Mb,d) andhornblende oxidation (Mh,o) likely reflects the sequestrationof Mg-within secondary phases or loss of Mg during biotiteoxidation, which was not considered here. Vermiculitelayers interstratified with altered biotite in the protosapro-lite zone are one likely candidate for a Mg sink. Thisintermediate phase, although detected by XRD and TEMin the saprolite by Murphy et al. (1998) and Dong et al.(1998), were not modeled here because they vary in compo-sition and are a very small component of the total mineral-ogy. Despite this simplification, these results clearly showthat Fe(II) loss across the rindlet zone is dominated by bio-tite oxidation. Optically, hornblende oxidation does not ap-pear to be significant within the rindlet zone. Thisconclusion is supported by these calculations and by EPMAresults, which do not indicate any loss of cations fromhornblende as a function of distance across the rindlet zone(Fig. 7).

The stoichiometric coefficients for Fe and K during bio-tite oxidation, mFeðIIÞ

b;o and mKb;o, represent an increase in ferric

iron in the mineral formula of 26.3%: the observed

Table 4List of variables for Eqs. (8)–(11)

mFeðIIÞb;o 0.36 mol mol�1 mol Fe(II) lost during biotite oxidation

mFeðIIÞh;o 1.95 mol mol�1 mol Fe(II) lost during hornblende oxidation

mFeðIIÞb;d 1.23 mol mol�1 mol Fe(II) lost during biotite dissolution

mFeðIIÞh;d 1.95 mol mol�1 mol Fe(II) lost during hornblende dissolution

mFeðIIÞc;d 0.29a mol mol�1 mol Fe(II) lost during chlorite dissolution

mKb;o 0.36 mol mol�1 mol K lost during biotite oxidation

mKb;d 0.88 mol mol�1 mol K lost during biotite dissolution

mKh;d 0.06 mol mol�1 mol K lost during hornblende dissolution

mMgh;o 0.973 mol mol�1 mol Mg lost during hornblende oxidation

mMgb;d 1.19 mol mol�1 mol Mg lost during biotite dissolution

mMgh;d 2.637 mol mol�1 mol Mg lost during hornblende dissolution

mMgc;d 3.32a mol mol�1 mol Mg lost during chlorite dissolution

mMnb;d 0.02 mol mol�1 mol Mn lost during biotite dissolution

mMnh;d 0.09 mol mol�1 mol Mn lost during hornblende dissolution

Mb,o 2.81 mol m�1 kg�1 mol biotite oxidized per unit rockMh,o �0.01 mol m�1 kg�1 mol hornblende oxidized per unit rockMb,d �1.05 mol m�1 kg�1 mol biotite dissolved per unit rockMh,d 0.39 mol m�1 kg�1 mol hornblende dissolved per unit rockMc,d 0.185 mol m�1 kg�1 mol chlorite dissolved per unit rock

bFe(II) 2.02 m kg mol�1 Fe(II) weathering gradientbK 8.92 m kg mol�1 K weathering gradientbMg 2.71 m kg mol�1 Mg weathering gradientbMn 72.3 m kg mol�1 Mn weathering gradient

a Chlorite variables determined from formula of Turner et al. (2003).



difference in oxidation between bedrock biotite and biotitein the deepest saprolite (Murphy et al., 1998). If we assume3.7% of the biotite iron in the corestone is ferric followingMurphy et al. (1998), then convert an additional 26.3% ofthe ferrous iron to ferric, removing the molar equivalentin K, the resulting composition is consistent the measuredbiotite compositions (EPMA) in the protosaprolite. Theseresults are consistent with a model in which iron in biotiteis oxidizing across the rindlet zone and ejecting K+ ions tocompensate for the loss of electrons. These results also indi-cate that biotite achieves the oxidation state of the saprolitebiotite (Murphy et al., 1998) within the rindlet zone.

When a chemical gradient can be assigned to a mineralweathering reaction, that gradient can be thought of asthe weathering front for that particular reaction. Thus,the Na gradient in the rindlet zone defines the plagioclaseweathering front and the Fe(II) gradient through the rindletand protosaprolite zones define the biotite oxidation andhornblende dissolution fronts, respectively. In the saprolite,the K gradient defines the biotite dissolution front (Buss,2006). Therefore, although the bedrock–saprolite interface(i.e., the rindlet zone) can be thought of as the bedrockweathering front, the system contains individual mineralweathering fronts that occur in different portions of theprofile and at different rates (Fig. 15).

4.2. Quantification of mineral weathering rates

Elemental weathering gradients as discussed above canalso be used to calculate long-term, average mineral reac-tion rates. Using the methodology of White (2002) we usethe solid state weathering gradients (Figs. 13 and 14) to cal-culate weathering rates across the rindlet zone for horn-blende and plagioclase dissolution and biotite oxidation.

The solid-state reaction rate R for a given mineral in aweathering profile is calculated from the elemental distribu-tion in the profile using the following expression fromWhite (2002):

R ¼ 10�3 1

ubsxbs

ð12Þ

where / is the mass fraction of the mineral in the weather-ing material (g g�1), b is the stoichiometric coefficient of theelement in the mineral (mol mol�1), s is the specific surfacearea of the mineral (m2 g�1), bs is the weathering gradient(m kg mol�1), and x is the weathering advance rate(m s�1). Here the weathering advance rate equals the totaldenudation rate.

4.3. Biotite weathering

In powder XRD analyses and TEM images, Murphy etal. (1998) and Dong et al. (1998) both noted the presence ofan altered biotite phase in the Rio Icacos saprolite having ad (00 1) spacing of 10.5 A (compared to the 10.0 A typicalof biotite); higher Al/Si and Fe(III)/Fe(II) ratios; and lowertotal Fe, Mg, and K than fresh biotite. XRD patterns ofthese altered biotite grains contain a (060) peak at 1.55 Aindicating a trioctahedral structure. A peak was not ob-served at 14 A, characteristic of vermiculite or chlorite(Murphy et al., 1998). Similarly, Mg2+ saturation did notyield a 24 A peak and the 10.5 A peak did not shift afterMg-saturation, Mg2+ saturation/ethylene glycol solvation,K+ saturation, or K+ saturation with heating to 110 �C(Murphy et al., 1998), all traditional indicators of interstr-atified vermiculite (e.g., Malla and Douglas, 1987;Sawhney, 1989). However, Murphy et al. (1998) identifiedseveral 14 A layers in biotite grains using TEM and inter-preted the 10.5 A d (001) peak and a co-existing 3.37 Apeak as evidence for a randomly interstratified biotite/ver-miculite phase (sometimes called hydrobiotite) containing�12% vermiculite. This interpretation was made based onMerings principles (summarized by Moore and Reynolds,1997), which state that randomly interstratified phases willproduce reflections between the (001) peaks of the end-members (e.g., 10.0 A biotite and 14 A vermiculite) thatare shifted relative to the proportion of the two endmem-bers in the mixed structure. When Dong et al. (1998) mea-sured altered biotite d (001) spacings with TEM excludingthe 14 A layers, the 10.5 A spacings were still observed.This result was taken as evidence that the altered biotiteis a separate phase, distinct from both fresh biotite and int-erstratified vermiculite/biotite. Possible causes of the expan-sion from 10.0 to 10.5 A were not discussed by Dong et al.(1998).

The altered biotite phase was found in the saprolite butnot in the bedrock, and was therefore presumed to form at

0.5

0.0

-0.5

-1.0

-1.5

-2.0

-2.5

-3.0

-3.5

-4.0

-1.0 -0.8 -0.6 -0.4 -0.2 0.0

Saprolite

Rindlets

τTi,j

Dis

tanc

e (m

)

PlagioclaseDissolution

Biotite Oxidation

HornblendeDissolution

Biotite Dissolution

Fig. 15. Semi-schematic diagram indicating the location andgradient of the individual mineral weathering fronts in a general-ized profile for the Rio Icacos watershed. sTi,j is the mass transfercoefficient (Eq. (2)) where j = Na (plagioclase dissolution), Fe(II)(biotite oxidation or hornblende dissolution), or K (biotite disso-lution). Dotted portions of the lines are extensions of the calculatedvalues. Distance is set to zero at the rindlet–saprolite interface withpositive values into the rindlet zone and negative values into thesaprolite. The thin dashed line indicates the corestone–rindletboundary. The biotite dissolution front was calculated from K datafrom Buss (2006) for a ridgetop saprolite profile near the roadcutstudied here. However, this front should be considered approxi-mate because Ti is a structural component of the biotite as well asthe normalizing element. The plagioclase, hornblende, and biotiteoxidation fronts were calculated from data collected on the roadcutsamples in the present study.



the bedrock–saprolite interface (Dong et al., 1998), in otherwords, within the rindlet zone. Murphy et al. (1998) notedthat biotite in the saprolite is oxidized relative to biotite inthe bedrock. Gradients in Fe(II) and K across the rindletzone (Fig. 14a and d) indicate that biotite oxidation occursthere, consistent with the formation of an oxidized ‘‘alteredbiotite” phase. However, we also documented zones of bio-tite oxidation coupled with K-depletion in the bedrockcorestone (Fig. 11), apparently preceding dissolution reac-tions and porosity development.

The expansion of biotite layers from 10.0 to 10.5 A dur-ing oxidation is consistent with the model proposed byFletcher et al. (2006). The weathering gradient for Fe(II)through the rindlet zone is consistent with the reaction pro-file calculated by Fletcher et al. (2006) (Fig. 16). This result,coupled with evidence that biotite oxidation is the firstweathering reaction that occurs in the corestone, suggeststhat biotite oxidation to form altered biotite is most likelythe reaction that initiates spheroidal fracturing. Dilationof biotite has been proposed by several researchers as amechanism for fracturing granite (e.g., Eggler et al., 1969;Isherwood and Street, 1976; Bustin and Mathews, 1979;Dixon and Young, 1981). However, the expansion is typi-cally explained as caused by the formation of vermiculitelayers or other clays during biotite weathering (e.g., Eggler

et al., 1969; Isherwood and Street, 1976). The relativelysmaller expansion that occurs during the oxidation of bio-tite may be a novel explanation for reaction-induced frac-turing in granitic rock.

A weathering rate for the oxidation of biotite in the rind-let zone can be determined using Eq. (12) and the weatheringgradient of Fe(II) across the rindlet zone (bs = 2.02 m kgmol�1) and the average weathering advance rate, x, of43 m Ma�1 (Brown et al., 1995). Other parameters usedare / = 0.095 g g�1 (White et al., 1998); specific BET surfacearea for fresh biotite, s = 0.1 m2 g�1 (Acker and Bricker,1992); and b = 0.36 mol Fe(II) mol�1, which is the stoichi-ometric coefficient of Fe(II) oxidation (mFeðIIÞ

b;o , Table 4).The resulting rate of biotite oxidation within the rindlet se-quence is 8.2 � 10�14 mol m�2 s�1, or log R = �13.1. Bycomparison, the rate of biotite weathering to kaolinite with-in the saprolite is log R = �15 (Murphy et al., 1998).

4.4. Plagioclase weathering

Plagioclase weathering occurs steadily across the rindletand protosaprolite zones, as evidenced by the Na gradient(Fig. 14e), and produces the porosity (void space withinremnant plagioclase rims observed in thin sections) thateventually causes the rindlets to disintegrate in the proto-saprolite zone. Eq. (12) can again be used to calculate a pla-gioclase weathering rate from the Na gradient (bs = 0.45m kg mol�1) and mineralogical parameters / = 0.564g g�1 (White et al., 1998), b = 0.5 mol Na mol�1 (Table3), and s = 0.1 m2 g�1. This specific BET surface area isan average value for fresh oligoclase or andesine of 300–600 lm grain size (Holdren and Speyer, 1987), which iswithin the range of plagioclase grain sizes in the corestoneand rindlets (�200–1000 lm). The resulting rate of plagio-clase weathering is 1.1 � 10�13 mol m�2 s�1, or log-R = �13.0. This rate is faster than that reported byTurner et al. (2003) for rindlet system plagioclase weather-ing (logR = �14.3 to �14.6). Although Turner et al. (2003)used a range of specific surface areas (0.1–1.0 m2 g�1), ourrates are still faster even using this range (logR = �13.0to �14.0). This apparent discrepancy may be a matter ofscale. Turner et al. (2003) calculated average plagioclaseweathering rates using sections of rindlet sets exhibiting dif-ferent degrees of weathering from different rindlet se-quences within the watershed. One would expect rindletsets of different orientations, overlying saprolite thickness-es, elevations, and proximity to surface waters to haveslightly different weathering rates. Consequently, a rate thatincorporates data from several different sample sets can beconsidered to be an average rate over a larger area. The ratepresented here considers a single sub-horizontal rindletzone atop a corestone and is therefore averaged over asmaller, more densely sampled scale.

4.5. Hornblende weathering

Previous researchers who have studied the Rio Icacosweathering system have noted the absence of hornblendein the saprolite (White et al., 1998; Turner et al., 2003).Hornblende is the most abundant Fe-silicate in the bed-

1.0

0.8

0.6

0.4

0.2

0.0

1.0 0.8 0.6 0.4 0.2 0.0

Nor

mal

ized

Dis

tanc

e fr

om S

apro

lite

Modelτ

Fe(II) measured

Fraction of Solid Reactant Consumed

Fig. 16. Fletcher et al. (2006) modeled the oxidation of Fe(II) (asFeO for simplicity) in the corestone hornblende to produce anFe(III) precipitate, which occurs with an increase in DV, causingelastic strain energy to build-up until the corestone fractures toproduce a rindlet. The fraction of total Fe(II) that has reactedcalculated by Fletcher et al. (2006) is shown here as a continuousline with steps representing the formation of rindlet-definingmacro-cracks. Here we compare the profile calculated by thenumerical model to the biotite oxidation gradient (the Fe(II)gradient across the rindlet zone, see Fig. 14a), represented assTi,Fe(II) normalized to the 0–1 scale, where 0 = the saprolite–rindletinterface and 1 = the corestone–rindlet interface. Similar to thenumerical model, biotite oxidation begins near the corestone with alarge loss of Fe(II), followed by a leveling off near the saprolite.



rock, but the mechanism and location of hornblende weath-ering has not been previously identified. Pseudomorphicreplacement of hornblende by clay minerals such as chlo-rite, chlorite–saponite, or saponite has been documentedin soil and saprolite (Wilson and Farmer, 1970; Anandand Gilkes, 1984), but no widespread evidence for suchphase changes was observed in the present samples. Simi-larly, dissolution of hornblende and re-precipitation of iron(oxy)hydroxides, gibbsite, or kaolinite can occur in soil,saprolite, and weathering rinds (e.g., Velbel, 1989), butwas not observed in the rindlet samples: i.e., we observedno etching of the hornblende grains, no boxwork textureor precipitates associated with the rindlet hornblende crys-tals. In fact, plagioclase and apatite inclusions within horn-blende crystals are commonly seen to weather leaving thesurrounding hornblende visibly un-weathered (Fig. 4d).EPMA and EDS analyses confirm that the surroundinghornblende is not measurably altered (Fig. 7). These resultssuggest that hornblende dissolution and subsequent precip-itation of Fe(III) (oxy)hydroxides is not the reaction thatproduces the positive DVreaction that initiates spheroidalfracturing as suggested by Fletcher et al. (2006).

In some systems, hornblende weathering can also be de-tected optically by color changes from dark green to palegreen due to loss of iron, or to brownish green, due to oxi-dation of iron (Deer et al., 1962; Wilson and Farmer, 1970).Although hornblende crystals in the outermost rindlets doappear to be slightly oxidized, overall, evidence for horn-blende weathering within the rindlet zone is scarce. There-fore, hornblende oxidation is also an unlikely candidatefor the reaction that causes the spheroidal fracturing.Although macro- and micro-cracks permit fluid flow intothe rindlet zone, total permeability remains low (Turneret al., 2003). Low permeability coupled with high mineral/fluid ratios and increased solute (Al, Si, Na, Ca, Mg) con-centrations due to plagioclase and, to a lesser extent, chlo-rite dissolution are inferred to keep the pore fluid within therindlet zone saturated with respect to hornblende, prevent-ing dissolution.

To estimate the dissolution rate of hornblende in theprotosaprolite using Eq. (12), we consider the gradient inFe(II) across the protosaprolite zone (bs = 0.088 m kgmol�1, Table 4) and mineralogical parameters /= 0.063 g g�1 (White et al., 1998), b = 1.95 mol Fe(II)mol�1 (Table 4), and s = 0.2 m2 g�1 (Brantley and Mellott,2000). The resulting rate of hornblende weathering is6.3 � 10�13 mol m�2 s�1 (logR = �12.2).

In a steady state profile, the 7 cm thickness of the proto-saprolite zone represents a very short time interval relativeto the time required to develop the entire profile (about 3 mtotal thickness at the sampled outcrop). Thus, the completeloss of hornblende over the 7 cm protosaprolite zone im-plies extremely rapid weathering, which is reflected by thecalculated hornblende weathering rate. This rate is fasterthan other published rates for hornblende weathering inthe field, but is within range of laboratory weathering rates(see compilations in White and Brantley, 1995; Brantley,2004). This similarity to laboratory rates is not so surprisingwhen we consider that the protosaprolite zone has high ex-posed mineral surface area, high porosity, and high fluid

flux. Typically mineral weathering rates estimated fromfield data are calculated from watershed fluxes and are aver-aged over the time period of profile development. In thepresent study, by densely sampling the weathering profileand calculating the hornblende dissolution rate for onlythe �7 cm thick zone where the reaction occurs, we obtainan in-situ rate that is not ‘‘diluted” over the entire profile orwatershed.

Microorganisms living at depth near the saprolite–bed-rock interface may benefit from an increase in availabilityof inorganic nutrients, Fe(II) in particular, which are re-leased from the bedrock during mineral weathering. Theferrous iron released from hornblende in the protosaprolitezone provides substrate for iron-oxidizing bacteria (Buss etal., 2005). Because most of the Fe(II) within biotite is oxi-dized in-situ rather than released, biotite does not contrib-ute significantly to the flux of Fe(II) available to deepsaprolite microorganisms. Iron-oxidizing bacteria fix CO2,producing organic carbon to support heterotrophic organ-isms in the ecosystem. Growth of lithoautotrophs such asiron-oxidizing bacteria in the Rio Icacos saprolite are fa-vored by substrate fluxes and are likely dominant at depth(Buss et al., 2005), contributing to the formation of Fe(III)-(hydr)oxides in the saprolite. The flux of Fe(II) from horn-blende dissolution in the protosaprolite could support amaximum growth rate of 4.5 � 10�5 mol C m�3 h�1 foriron-oxidizing bacteria, based on the model of Buss et al.(2005). Because these bacteria alter the concentrations ofFe(II) and O2, they may also affect the rate of spheroidalweathering.

5. CONCLUSIONS

The quartz diorite bedrock in the Rio Icacos watershedweathers spheroidally forming corestones surrounded byzones of concentric, partially weathered rindlets overlainby saprolite. This process produces two distinct weatheringinterfaces: (1) the bedrock–rindlet interface where the disag-gregation is initiated via spheroidal macro-cracking to formrindlets, and (2) the rindlet–saprolite interface where therindlets disintegrate into saprolite due to extensive micro-cracking and the completion of plagioclase and hornblendeweathering. Within the �7 cm thick protosaprolite zone,hornblende rapidly dissolves to completion at a rate of6.3 � 10�13 mol m�2 s�1, providing the dominant flux ofFe(II) to the saprolite and the saprolite biota. However,within the �47 cm rindlet zone, hornblende weathering isinsignificant, plagioclase dissolves at a rate of1.1 � 10�13 mol m�2 s�1, and biotite is oxidized at a rateof 8.2 � 10�14 mol m�2 s�1. The oxidation of biotite formsa K-deficient ‘‘altered biotite” with an expansion of the d

(00 1) spacing from 10.0 to 10.5 A. Oxidation of biotitewas also identified within individual biotite grains in thebedrock corestone by X-ray microprobe imaging andXANES as the earliest weathering reaction documentedin the bedrock. Expansion of biotite during oxidation with-in the corestone is consistent with the reaction-driven sphe-roidal fracturing model of Fletcher et al. (2006) and is likelythe reaction responsible for initiating the spheroidalfracturing and thus the disaggregation of intact bedrock.



Oxidation of biotite begins when O2 diffuses into the core-stone, and thus the weathering advance rate is dependentupon the concentration of O2 in the pore water. This con-centration is expected to vary with regolith thickness, pro-viding a positive feedback between denudation andweathering advance rates. The ability to map oxidationstates within individual crystals using synchrotron-basedtechniques has proven invaluable for the observation ofincipient weathering reactions within ‘‘pristine” bedrock,facilitating identification of rate limiting reactions in cou-pled processes such as spheroidal weathering.

ACKNOWLEDGMENTS

We thank A.F. White and R.C. Fletcher for helpful discus-sions and field support; D. Eggler and E. Merino for assistancewith optical microscopy; M. Angelone and J. Cantolina foranalytical assistance; M. Rosario-Torres, J. Troester, andG. Hernandez for field support; and S. Anderson and three anon-ymous reviewers for comments that helped improve themanuscript. Funding provided by DOE grant no. DE-FG02-05ER15675, the Penn State Biogeochemical Research Initiativefor Education (BRIE) supported by NSF-IGERT Grant No.DGE-9972759, and the Penn State Center for EnvironmentalChemistry and Geochemistry. H.L. Buss acknowledges fellowshipsupport of the NSF Graduate Research Fellowship Program andpostdoctoral support from the National Academy of Sciences Re-search Associateship Program. S.L Brantley acknowledges sup-port from the Center for Environmental Kinetics Analysissupported by NSF Grant No. CHE-0431328. Portions of this re-search were carried out at the Stanford Synchrotron RadiationLaboratory, a national user facility operated by Stanford Univer-sity on behalf of the U.S. Department of Energy, Office of BasicEnergy Sciences. The SSRL Structural Molecular Biology Pro-gram is supported by the Department of Energy, Office of Bio-logical and Environmental Research, and by the NationalInstitutes of Health, National Center for Research Resources,Biomedical Technology Program.

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