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Geochimica et Cosmochimica Acta 73 (2009) 2034–2060 www.elsevier.com/locate/gca

Biogeochemistry of mineral–organic associations across a long-term mineralogical soil gradient (0.3–4100 kyr), Hawaiian Islands

Robert Mikutta a,g,*, Gabriele E. Schaumann b, Daniela Gildemeister b, Steeve Bonneville c, Marc G. Kramer d, Jon Chorover e, Oliver A. Chadwick f, Georg Guggenberger a,g

a Soil Sciences, Martin Luther University Halle-Wittenberg, Weidenplan 14, 06108 Halle (Saale), Germany b Environmental and Soil Chemistry, University of Koblenz-Landau, Germany c School of Earth and Environment, University of Leeds, UK d Earth and Planetary Sciences, University of California, USA e Department of Soil, Water and Environmental Science, University of Arizona, USA f Department of Geography, University of California, USA g Institute of Soil Science, Leibniz University Hannover, Germany

Received 14 August 2008; accepted in revised form 16 December 2008; available online 22 January 2009

Abstract

Organic matter (OM) in mineral–organic associations (MOAs) represents a large fraction of carbon in terrestrial eco- systems which is considered stable against biodegradation. To assess the role of MOAs in carbon cycling, there is a need to better understand (i) the time-dependent biogeochemical evolution of MOAs in soil, (ii) the effect of the mineral com- position on the physico-chemical properties of attached OM, and (iii) the resulting consequences for the stabilization of OM. We studied the development of MOAs across a mineralogical soil gradient (0.3–4100 kyr) at the Hawaiian Islands that derived from basaltic tephra under comparable climatic and hydrological regimes. Mineral–organic associations were characterized using biomarker analyses of OM with chemolytic methods (lignin phenols, non-cellulosic carbohydrates) and wet chemical extractions, surface area/porosity measurements (N2 at 77 K and CO2 at 273 K), X-ray photoelectron spectroscopy (XPS), transmission electron microscopy (TEM), and differential scanning calorimetry (DSC). The results show that in the initial weathering stage (0.3 kyr), MOAs are mainly composed of primary, low-surface area minerals (olivine, pyroxene, feldspar) with small amounts of attached OM and lignin phenols but a large contribution of micro- bial-derived carbohydrates. As high-surface area, poorly crystalline (PC) minerals increase in abundance during the sec- ond weathering stage (20–400 kyr), the content of mineral-associated OM increased sharply, up to 290 mg C/g MOA, with lignin phenols being favored over carbohydrates in the association with minerals. In the third and final weathering stage (1400–4100 kyr), metastable PC phases transformed into well crystalline secondary Fe and Al (hydr)oxides and kaolin minerals that were associated with less OM overall, and depleted in both lignin and carbohydrate as a fraction of total OM. XPS, the N2 pore volume data and OM–mineral volumetric ratios suggest that, in contrast to the endmem- ber sites where OM accumulated at the surfaces of larger mineral grains, topsoil MOAs of the 20–400-kyr sites are com- posed of a homogeneous admixture of small-sized PC minerals and OM, which originated from both adsorption and precipitation processes. The chemical composition of OM in surface-horizon MOAs, however, was largely controlled by the uniform source vegetation irrespective of the substrate age whereas in subsoil horizons, aromatic and carboxylic C correlated positively with oxalate-extractable Al and Si and CuCl2-extractable Al concentrations representing PC

* Corresponding author. Address: Soil Sciences, Martin Luther University Halle-Wittenberg, Weidenplan 14, Halle, 06108 Saale, Germany. E-mail address: [email protected] (R. Mikutta).

0016-7037/$ - see front matter Ó 2009 Elsevier Ltd. All rights reserved. doi:10.1016/j.gca.2008.12.028 Evolution of mineral–organic associations upon substrate aging 2035 aluminosilicates and Al-organic complexes (r2 > 0.85). Additionally, XPS depth profiles suggest a zonal structure of sorbed OM with aromatic carbons being enriched in the proximity of mineral surfaces and amide carbons (peptides/pro- teins) being located in outer regions of MOAs. Albeit the mineralogical and compositional changes of OM, the rigidity of mineral-associated OM as analyzed by DSC changed little over time. A significantly reduced side chain mobility of sorbed OM was, however, observed in subsoil MOAs, which likely arose from stronger mineral–organic bindings. In conclusion, our study shows that the properties of soil MOAs change substantially over time with different mineral assemblages favoring the association of different types of OM, which is further accentuated by a vertical gradient of OM composition on mineral surfaces. Factors supporting the stabilization of sorbed OM were (i) the surface area and reactivity of minerals (primary or secondary crystalline minerals versus PC secondary minerals), (ii) the association of OM with micropores of PC minerals (via ‘sterically’ enhanced adsorption), (iii) the effective embedding of OM in ‘well mixed’ arrays with PC minerals and monomeric/polymeric metal species, (iv) the inherent stability of acidic aromatic OM components, and (iv) an impaired segmental mobility of sorbed OM, which might increase its stability against desorption and microbial utilization. Ó 2009 Elsevier Ltd. All rights reserved.

1. INTRODUCTION due to chemical and molecular size fractionation (Aufdenk- ampe et al., 2001; Chorover and Amistadi, 2001; Feng Up to >90% of a soil’s carbon inventory resides in inti- et al., 2005). In soils, Fe oxides exhibit a strong affinity mate association with minerals (Baisden et al., 2002; Basile- for highly oxidized aromatic, lignin-derived OC (Kaiser Doelsch et al., 2007; Crow et al., 2007). Organic C (OC) et al., 1997; Chorover and Amistadi, 2001; Mikutta et al., bound to minerals is considered relatively stable (Torn 2007), whereas river sediments and kaolinite were reported et al., 1997, 2005; Keil et al., 1998; Kalbitz et al., 2005; Mi- to retain N-containing compounds selectively (Aufdenk- kutta et al., 2006) and controls the long-term sequestration ampe et al., 2001). A mineralogical impact on OM compo- of organic matter (OM) in terresrial and marine ecosystems. sition has also been tentatively suggested by analysis of OM The ability of ecosystems to sequester atmospheric carbon composition (lignin, cutin and suberin acids) of mineralog- thus depends on the partitioning of soil and sediment car- ically diverse density fractions of an Andic Dystrudept A bon into more stable fractions (IPCC, 2001) such as min- horizon (Sollins et al., 2006). Straightforward evaluation eral–organic associations (MOAs).1 Organic matter of a mineralogical control can, however, be compromised stabilization by minerals has been envisioned as a combined by the fact that MOM is subjected to in-situ microbial effect of the chemical properties of mineral sorbents (sur- transformations, which may mask the mineral-induced face area and charge) and the structural properties of the chemical imprint (Keil et al., 1998). Most studies related OM involved, with aromatic structures being most resistant to sorptive fractionation and differential stabilization of towards biodegradation (Kalbitz et al., 2005; Mikutta et al., OM at mineral surfaces have been conducted in sediments 2007). (Keil et al., 1998; Arnarson and Keil, 2001; Dickens Most primary minerals in soils are in disequilibrium et al., 2006; Arnarson and Keil, 2007), thus limiting our with extrinsic environmental factors like the temperature understanding of long-term compositional changes of and hydrochemical regime. Thus, weathering triggers the MOM in soil environments. formation of thermodynamically more stable mineral Whereas the chemical composition of soil OM has been phases such as Fe and Al (hydr)oxides and aluminosilicate subject to numerous studies, little is known about the time- clay minerals. As a result, the inventory of clay and second- dependent alteration of the physical properties of mineral- ary metal (hydr)oxides in soils can increase substantially attached OM in natural environments. Soil OM undergoes over time spans of thousands of years (e.g., Harden, 1987; structural changes upon aging (Schaumann, 2006; Hurrass Masiello et al., 2004). Such mineralogical shifts affect the and Schaumann, 2007). Differential scanning calorimetry amount of soil OC stored (Torn et al., 1997; Brenner experiments indicate a decrease in segmental mobility of et al., 2001; Basile-Doelsch et al., 2007), but little is known OM over time, likely as a result of water- or metal-mediated about compositional variations of mineral-associated OM cross-linkings between OM chains (Schaumann, 2005). The (MOM) in different mineralogical settings. Batch experi- flexibility of OM bound directly to mineral surfaces is ex- ments reveal that OM often sorbs to minerals selectively pected to decrease as more ligands are involved in mineral attachment, and as such, flexibility should depend on the type of minerals present and on the amount and composi- 1 Abbreviations used: MOA, mineral–organic association; A- tion of OM accumulated. Steric rearrangements of OM at MOA and B-MOA, mineral–organic associations from A and B mineral surfaces leading to the neoformation of organic–or- horizons, respectively; LSAG, long substrate age gradient; MOC ganic (Collins et al., 1995) or organic–mineral bonds (Kai- and MN, mineral-associated organic carbon and total nitrogen; ser et al., 2007) are thus supposed to confer a greater MOM, mineral-associated organic matter; OC, organic carbon; stability to sorbed OM. In consequence, mineralogical PC, poorly crystalline; SSA, specific surface area; TEM–EDS, changes over time affect mineral–OM interactions in at least Transmission electron microscopy–energy dispersive spectroscopy; three aspects: (i) by the development of reactive surfaces XPS, X-ray photoelectron spectroscopy. 2036 R. Mikutta et al. / Geochimica et Cosmochimica Acta 73 (2009) 2034–2060 that will enable OM to interact with minerals, (ii) by creat- MOAs was examined by XPS and molecular biomarker ing conditions that allow for sorptive fractionation pro- analyses (lignin phenols, non-cellulosic carbohydrates), cesses of OM, and (iii) by changing the physical and (iii) the alteration of the physical state of OM in MOAs properties of sorbed OM. over time was investigated by differential scanning calorim- This study investigates the changing nature of mineral– etry (DSC). organic associations during progressive basalt weathering (Pliocene to Holocene) in the Hawaiian archipelago. We 2. MATERIALS AND METHODS complement previous research on mineral–OM relation- ships in soils of this weathering gradient known as ‘long 2.1. Study site and sample collection substrate age gradient’ (LSAG) where the transformation of mineral matter over 0.3–4100 kyr causes a differential Detailed descriptions of the geological, climatic and mor- storage of OM (Torn et al., 1997, 2005; Chorover et al., phological settings of the LSAG sites are given in Vitousek 2004; Vitousek, 2004). The main objective of this study is (2004) and Chorover et al. (2004). All soils were formed from to examine the time-dependent formation and characteris- basaltic materials composed primarily of tephra overlying tics of MOAs and the consequences for the chemical and lava or a lava-ash mixture. Parent material age, measured physical properties of MOM, and thus to contribute to using 14C and K/Ar radiometric techniques, ranges from the understanding of OM stabilization in soils. We hypoth- 0.3 to 4100 kyr (Fig. 1 and Table 1). The sites are located at esize that the time-dependent transformation of primary sil- 1130–1500 m elevation above sea level with a mean annual icates dominating in juvenile basaltic lava into more mature temperature of 289 K and a mean annual precipitation of secondary minerals is accompanied by distinct changes in 2500 mm. At each site the montane rainforest is dominated the physico-chemical properties of MOM. The LSAG is by Metrosideros polymorpha (>75%) while the subcanopies most suitable to trace the impact of minerals because each are more diverse across the sites, including different native site has a similar geological setting, climate, and plant spe- tree ferns in the genus Cibotium. Soils of the six sites have cies composition, but different soil mineral composition. been classified as (1) Thaptic Udivitrand (Thurston), (2–4) The similar plant species assemblage at each site is critical Aquic Hydrudand (Laupahoehoe, Kohala, Pololu), (5) to the study because it minimizes the influence of primary Aquic Hapludand (Kolekole), and (6) Plinthic Kandiudox organic resources on the chemical variability of MOM. (Kokee) (Chorover et al., 2004). Source vegetation, organic Mineral–organic associations were isolated densiometrical- layers, and mineral A and B horizons (0–20; 10–100 cm; Ta- ly and examined along the LSAG according to three as- ble 1) were sampled in June 2007 analogous to the sampling pects: (i) changes in mineralogical composition, surface scheme described in Chorover et al. (2004). Samples were area and porosity were elucidated by selective extractions, carefully mixed in the field, double packed in Ziploc bags, X-ray photoelectron spectroscopy (XPS), gas adsorption and stored in the laboratory coolers on blue ice. High prior- (N2 at 77 K and CO2 at 273 K), and transmission electron ity-shipping to Germany was accomplished at a temperature microscopy coupled with energy dispersive spectroscopy <278 K within four days. In the laboratory, soil samples were (TEM–EDS); (ii) the chemical transformation of OM in immediately pushed through a 2-mm sieve, roots, visible

Fig. 1. The Hawaiian archipelago and study sites (ESRI data base). Table 1 Site age, origin, sampling depth and pH of soils, and basic properties of isolated mineral–organic associations (MOAs) from A and B horizons of the long substrate age gradient. a a b d d d e d Site Island Depth Soil pH Absolute MOC MN C/N MOCNaOH XAD- Fedith Feox Feox / Alox AlCu Siox Clay mineral c f age (H2O) density 8 Fedith composition (MOA) vlto fmnrlogncascain pnsbtaeaging substrate upon associations mineral–organic of Evolution (kyr) cm (g/cm3) (mg/g) (mg/g) (%) (%) (mg/g) (mg/g) (mg/g) (mg/g) (mg/g) A horizons Thurston 0.3 Hawaii 2–9 5.5 2.9 18.7 (0.8) 2.3 (0.3) 8 70 56 5.0 4.3 0.9 2.6 1.5 0.8 P, F, a Laupahoehoe 20 Hawaii 0–10 4.1 2.4 177.5 (0.2) 9.9 (0.1) 18 47 66 153.6 84.2 0.5 3.3 0.9 0.3 F, A, m, q Kohala 150 Hawaii 0–7 3.9 2.1 187.6 (0.1) 13.2 (0.1) 14 63 68 46.9 37.4 0.8 18.5 11.1 0.8 A, F, v, hiv, q Pololu 400 Hawaii 0–10 3.9 1.8 290.4 (0.0) 19.4 (0.1 15 66 81 43.7 40.0 0.9 7.6 6.1 0.1 A, F, v, hiv, k, q Kolekole 1400 Molokai 0–8 4.2 2.3 131.1 (0.1) 9.7 (0.1) 13 60 65 40.1 15.4 0.4 4.2 1.9 0.1 K, Gi, F, he, q Kokee 4100 Kauai 0–16 4.8 3.2 40.0 (0.1) 2.8 (0.2) 14 74 49 241.3 4.4 0.0 0.4 0.2 0.0 K, Gi, Go, he B horizons Thurston 0.3 — 15–23 6.1 2.8 23.2 (0.1) 2.5 (0.1) 9 46 51 11.5 10.7 0.9 8.1 1.5 4.1 P, F, a Laupahoehoe 20 — 45–67 5.3 2.3 113.6 (0.1) 5.6 (0.2) 20 86 46 100.8 77.9 0.8 113.5 7.6 29.5 F, A, m, q Kohala 150 — 41–60 4.0 2.1 120.6 (0.1) 4.4 (0.1) 27 92 37 15.4 16.9 1.1 137.8 15.9 44.7 A, F, hiv, q Pololu 400 — 44–64 4.9 2.5 81.3 (0.0) 3.6 (0.2) 23 88 46 124.1 56.6 0.5 50.5 9.2 4.9 A, F, v, hiv, k, q Kolekole 1400 — 30–49 4.9 2.6 19.9 (0.1) 2.0 (0.1) 10 100 29 53.3 14.4 0.3 14.0 2.4 2.0 K, Gi, f, he, q, m Kokee 4100 — 50–100 5.2 3.1 12.1 (0.1) 1.6 (0.2) 7 75 21 143.9 2.3 0.0 2.1 0.4 0.1 K, Gi, Go, he, m a MOC and MN = mineral-associated organic C and N. b Fraction of the mineral-associated OC extractable in 0.1 M NaOH. c Fraction of the NaOH-extractable OC that adsorbs to XAD-8 resin after acidification to pH 2. d Fedith = dithionite-extractable Fe; Feox ,Alox,Siox = oxalate-extractable Fe, Al, Si. e AlCu = CuCl2-extractable Al. f Data compiled from Chorover et al. (2004): upper case letters indicate major constituents, lower case letters indicate minor constituents; P, plagioclase; A, short-range-ordered Al gels or aluminosilicates (e.g., allophane); F, ferrihydrite; Q, quartz; V, vermiculite; HIV, hydroxyl-interlayered vermiculite; K, kaolin minerals (kaolinite and/or halloysite); Gi, gibbsite; He, hematite; Go, goethite; M, magnetite. 2037 2038 R. Mikutta et al. / Geochimica et Cosmochimica Acta 73 (2009) 2034–2060 plant and animal remains were removed, and the material OC from MOAs was tested by treatment with a diluted was stored field-moist in the dark at 277 K until use. NaOH solution in N2 atmosphere. Briefly, 0.5 g of pre- dried MOAs were extracted three times with 20 mL of 2.2. Density fractionation and chemical analyses 0.1 M NaOH (1/40 wt./vol.). After shaking for 24 h in the dark (250 rpm; 298 K), samples were centrifuged (20 min; Density fractionation: Mineral–organic associations with 10,000g), filtered through polyvinylidene fluoride syringe q > 1.6 g/cm3 were obtained by suspending the field-moist filters (<0.45 lm), and the amount of extracted OC was (20–74 wt.% H2O), low-aggregated samples in sodium poly- analyzed as non-purgeable OC by Pt-catalyzed oxidation tungstate (SPT) solution. Contaminant OC and N concen- at 993 K (Shimadzu TOC-V, Japan). The replicate variabil- trations in the original SPT salt were at the detection limit ity of OC concentration was <2%. The contribution of aro- (0.001% OC and 0.001% N). Briefly, between 31 and matic components in extracted OM, defined as 98 g field-moist soil (25 g dry weight after freeze-drying at ‘hydrophobic’ OC, was estimated by XAD-8 fractionation 0.3 mbar) were mixed with 125 mL SPT giving a final den- (Aiken and Leenheer, 1993). sity of 1.6 ± 0.01 g/cm3. Following vigorous agitation on a horizontal shaker (300 rpm) for 30 min, the suspensions 2.4. X-ray photoelectron spectroscopy (XPS) were left to stand for 30 min and then centrifuged at 4500g for 15 min. If dispersion was incomplete, the shaking Mineral–organic associations sieved to a size of period was extended. A 1.6-g/cm3 density cutoff was chosen <200 lm were analyzed in triplicate by XPS using a PHI because MOAs, e.g., those derived from coprecipitation of 5700 ESCA-System instrument (Physical Electronics, OM with Al, can exhibit very low absolute densities (1.7 g/ Chanhassen, USA). Pre-dried samples were pressure- cm3)(Kaiser and Guggenberger, 2007a). Ultrasonic disper- mounted onto a conducting indium foil and inspected visu- sion was avoided because of the risk to transfer MOAs con- ally using the instrument’s microscope to identify suitable taining poorly crystalline (PC) minerals into the density areas without large mineral grains. This procedure was pri- fraction <1.6 g/cm3 (Kaiser and Guggenberger, 2007a). marily important for the youngest sample (Thurston), The floating particulate OM was transferred on 0.7-lm which still contains larger primary mineral particles. Then glass microfibre filter (GF/F Whatman) and vacuum fil- samples were excited with non-monochromated Al Ka radi- tered. The procedure was repeated until no floating partic- ation (Eexc = 1486.6 eV) at an incident angle of 45° and an ulate OM was observed. The particulate OM and MOA electron beam spot size of 800 lm(500 lm2). Surface fractions were washed with distilled water to an electric composition analyses of MOAs up to a depth of 3–5 nm conductivity of <50 dS/m. In some cases, colloids started included a survey element scan and a high-resolution scan to disperse at 100 dS/m. Then further washing was at the C1s edge. The bulk element composition of MOAs avoided. To prevent slow-freezing artefacts on surface was revealed by scanning from 0 to 1100 eV using a pass en- properties of MOAs, samples were shock-frozen in liquid ergy of 93.90 eV, a channel width of 0.4 eV/step, and a mea- N2 (77 K) and subsequently freeze-dried at 0.3 mbar. sure time of 20 ms/step (exposition time = 10 min); in the Chemical analyses: Selective extractions of MOAs were high-resolution mode, MOAs were scanned at the C1s edge carried out in triplicate at 298 K in the dark using pre-dried using a pass energy of 11.75 eV, a channel width of 0.1 eV/ samples (0.3 mbar for 24 h) by (i) dithionite–citrate (Blake- step and measure time of 100 ms/step (exposition more et al., 1987), (ii) oxalic oxalate (Blakemore et al., time = 30 min). Significant X-ray-induced alteration of 1987), and (iii) copper(II) chloride (Juo and Kamprath, MOM that could cause false structural C assignments is ex- 1979). Element concentrations in extraction solutions were pected only for exposure times >30 min (Dengis et al., 1995; measured by inductively coupled plasma atomic–emission Zubavichus et al., 2004). Vacuum during measurements spectroscopy (ICP–AES, Jobin Yvon JY 70 plus, France) was 3 109 mbar. Contamination rate by adventitious after calibration with element standards dissolved in the C was estimated to be 0.1 atom%/h. Sample charging dur- respective extraction matrices. Dithionite–citrate extract- ing analysis lead to peak shifts of 63 eV, which was cor- able Fe represents non-silicate Fe, i.e., crystalline and PC rected based on the maximum principal C1s sub-peak Fe (hydr)oxides; oxalate-extractable Fe, Al and Si derive centered at 285 eV. Data analysis was accomplished with from PC aluminosilicates, ferrihydrite/nanocrystalline goe- the instrument software package (Multipak V6.1A; Physi- thite, and Al and Fe in organic complexes, whereas the dif- cal Electronics) using manufacturer-based sensitivity fac- ference between dithionite–citrate and oxalate-extractable tors whereas graphics were generated using Unifit for Fe (Fedith–ox) is taken as a measure of the content of crys- Windows, Version 2008 (Hesse et al., 2003). Identification talline Fe oxides. Copper(II) chloride-extractable Al mainly of binding energies was done according to Moulder et al. represents organically complexed and exchangeable Al (Juo (1992). The chemical composition of MOM was revealed and Kamprath, 1979). by deconvoluting the C1s peak into sub-peaks by fitting Gaussian–Lorentzian functions (85%/15%) that are as- 2.3. Organic carbon, total nitrogen, and extractability signed to different C environments (NIST XPS database, Monteil-Rivera et al., 2000). The full-width-at-half-maxi- The OC and total N content of finely ground and pre- mum (FWHM) of fitted sub-peaks was allowed to vary be- dried samples were measured in triplicate by dry combus- tween 0 and 1.5. Additionally, we analyzed the chemical tion at 1423 K using a Vario EL CNS analyzer (Elementar composition of MOAs from A horizons along depth pro- Analysensysteme, Hanau, Germany). The extractability of files in one representative sample area (500 lm2) immedi- Evolution of mineral–organic associations upon substrate aging 2039

+ ately after sputtering with Ar ions for 15 min (t15) and He pycnometer (Micromeritics ACCUPyc 1330, Norcross, 14 2 30 min (t30), respectively [ion fluence = 1.6 10 /(cm s)]. USA). The experimental error was 61%. All surface area Thirty minutes were set as time limit for Ar+ gassing to and porosity analyses were carried out at least in duplicate. avoid the reduction of C1s signals below that acceptable for peak processing. Etching velocity was 1.4 nm/min 2.6. Transmission electron microscopy (TEM) for pure SiO2. While the sputter efficiency for OM and min- erals differs and hence the actual sputter depth remains un- Before TEM analysis, MOA samples from A horizons known, this procedure is appropriate to reveal shifts in the were suspended in methanol and sonicated for 5 min. One vertical composition of OM at particle surfaces (Amelung drop of the suspensions was placed onto a porous carbon et al., 2002). film (Agar Scientific Ltd., Stansted, UK) on a Cu TEM grid and left to dry at room temperature for 5 min. Particle 2.5. Surface area and porosity analyses aggregates covering holes in the support film were imaged on a FEI CM200 instrument (FEI Company, Eindhoven, Gas adsorption and desorption was measured with a The Netherlands) equipped with a field-emission gun oper- Nova 4200 analyzer (Quantachrome Corp., Boynton ated at an acceleration voltage of 200 kV. Energy dispersive Beach, USA) using N2 and CO2 as adsorbates. For N2 X-ray spectroscopy (EDS) analyses were performed on adsorption, about 100–1000 mg of solid was weighed into MOA particles for areas of various size (20–200 nm) using the sample cell to ensure sufficient surface area for analy- a UTW ISIS EDX analyzer with an ultrathin window sis. Adsorbed water was removed by outgassing the sam- allowing detection of light elements (Oxford Instruments, ple at 313 K and 103 mbar for 48 h. After re-weighing, United Kingdom). N2 adsorption–desorption isotherms were recorded at 3 77 K in the partial pressure region 2.5 10 –0.99 P/P0 using 40 adsorption and 40 desorption points. P0 was 2.7. Differential scanning calorimetry (DSC) remeasured after every 20th adsorption/desorption point to account for short-term alterations of atmospheric pres- To characterize the degree of rigidity of MOM we used sure. Thermal equilibration was 90 s, equilibration time DSC and determined the glass transition-like step transi- was set minimal 240 s and maximal 400 s, and pressure tion temperature (T*) of OM as a measure for the matrix tolerance was 0.1 mm Hg. The specific surface area rigidity (Hurrass and Schaumann, 2005; Schaumann and (SSA) was determined from the relative pressure range LeBoeuf, 2005). Three to six replicates of each sample of 0.05–0.3 by the BET equation (Brunauer et al., 1938). (2–8 mg MOA) that were equilibrated for at least two In the BET model, the C constant expresses the energy weeks at 293 K and 31% humidity were weighed into stan- of adsorption to a surface and relates to the net enthalpy dard Al pans. The pans were sealed hermetically prior to of N2 adsorption (DHxs) as follows (Rouquerol et al., the DSC experiment. Differential scanning calorimetry 1998; Mayer, 1999) experiments were performed using a DSC Q1000 calorim- eter (TA Instruments, Alzenau, Germany) with TZero CBET exp½ðDH ads DH condÞ=RT expðDH xs=RT Þð1Þ Ò technology and a refrigerated cooling system, and N2 where DHads is the adsorption enthalpy of the gas directly as a purge gas. The samples were abruptly cooled in the on the surface, DHcond is the enthalpy of gas condensation, DSC instrument to 223 K and then heated at a rate of R is the universal gas constant, and T is the temperature of 10 K/min from 223 K to 383 K and 403 K for samples gas adsorption. The mineral SSA of MOAs still accessible from A and B horizons, respectively. This first heating cy- corr to N2 (SSA ) was estimated by correction of gas adsorp- cle was followed by a second abrupt cooling and a subse- tion data for OM contents according to: quent heating cycle. The DSC data were analyzed using corr Universal Analysis software Version 4.1 (TA Instruments). SSA ¼ SSA=ð1 MOM=1000Þð2Þ Baseline correction was performed by subtracting the ther- where SSA denotes measured specific surface area values mogram of the second heating cycle from that of the first (m2/g), and MOM refers to the organic matter content one to separate the non-reversing from the reversing ef- (2 MOC; mg/g). The volume of mesopores was deter- fects as discussed by Hurrass and Schaumann (2005). mined by the Barrett–Joyner–Halenda model (Barrett The glass transition-like step transition is indicated by et al., 1951) using the adsorption leg of the isotherm; the to- an inflection point in the thermogram (Hurrass and tal pore volume was estimated at P/P0 = 0.99. Schaumann, 2005; Schaumann and LeBoeuf, 2005). Oper- Carbon dioxide adsorption was performed at 273 K in ationally, three tangent lines were applied and T* was de- 6 an ice bath from 2 10 to 0.03 P/P0 using 38 adsorption fined as the temperature at the curve inflection. In order to points. Outgassing and measuring conditions were equal to verify that transitions are solely due to OM, we also ana- the N2 adsorption experiment. Before determination of the lyzed the thermal behavior of a suite of pure minerals at weight of the water-free sample and to avoid contamination 31% humidity including 2-line ferrihydrite, allophane, goe- with CO2, the sample was backfilled with N2. For both thite, PC Al(OH)3, gibbsite, vermiculite, and kaolinite adsorbates, micropore volumes were determined by the (clay minerals <2 lm, Na+-saturated). We found no sig- Dubinin–Radushkevich equation (Gregg and Sing, 1982). nificant transition in the temperature range examined. The absolute density of MOAs (Table 1), as required by Only ferrihydrite revealed a small transition at 363– the instrument software, was measured in triplicate with a 373 K but neither the transition temperature nor the tran- 2040 R. Mikutta et al. / Geochimica et Cosmochimica Acta 73 (2009) 2034–2060 sition intensity in the MOA samples correlated with the standard) plus 200 lL derivatization solution (20 mg/mL ferrihydrite content (oxalate-extractable Fe), suggesting o-methylhydroxylamine in pyridine) into the vials and leav- that minerals do not significantly contribute to the ob- ing them for 30 min at 348 K. Then, after cooling for served step transitions. 15 min, 400 lL N,O-bis(trimethylsilyl)trifluoroacetamide were added and samples were again heated to 348 K for 2.8. Chemical analysis of mineral-associated organic matter 5 min. After derivatization, samples were measured by cap- illary gas chromatography (GC-2010, Shimadzu Corp., To- Lignin-derived phenols: Lignin-derived phenols were kyo, Japan) equipped with a fused silica SPB-5 column quantified in duplicate after CuO oxidation. Briefly, 25– (30 m, 0.25 mm i.d., 0.25-lm film, Supelco). Helium served 500 mg of organic layer and MOA samples were mixed with as carrier gas at a constant pressure of 1 bar. The injector 15 mL 2 M NaOH, 50–100 mg Fe(NH4)2(SO4)2 6H2O, 250– was set to 523 K and the FID to 573 K. The temperature 500 mg CuO and 50 mg glucose and oxidized for 2 h at program comprised the following steps: heating to 433 K 443 K (Amelung et al., 1999). The lignin-derived phenols for 4 min, heating to 458 K (rate: 8 K/min), 1.5 min at were extracted from the acidified solution (pH 1.8–2.2) 458 K, heating to 523 K (rate: 4 K/min), 0.5 min at using C18 solid-phase columns (Mallinckrodt Baker Corp., 458 K, heating to 573 K (rate: 50 K/min) and 5 min at Phillipsburg, NJ), eluted with ethyl acetate and derivatized 573 K. The recovery of myo-inositol from MOA samples with a 1:1 mixture of pyridine and N,O-bis(trimethyl- averaged 102 ± 15%, those of litter samples averaged silyl)trifluoroacetamide. The derivatives were separated 115 ± 8%, and those of O samples averaged 116 ± 4%. and quantified using a gas chromatograph mass spectrom- eter (GCMS-QP 2010, Shimadzu Corp., Tokyo, Japan) and 2.9. Terminology and statistics a SPB-5 fused silica column (30 m, 0.25 i.d., 0.25-lm film; Supelco, Bellefonte, PA). Injector temperature was 523 K In the following, mineral–organic associations from A and the temperature programme included 373 K for horizons are termed ‘A-MOAs’, those from B horizons 1 min, 15 K/min to 523 K, 523 K for 10 min, 30 K/min to ‘B-MOAs’, respectively. The Pearson correlation coefficient 573 K, 573 K for 7 min. Helium was used as carrier gas was calculated based on z-transformed variables using Stat- with a constant pressure of 1.1 bar. As internal recovery istica 5.1 (StatSoft Inc., Tulsa, USA). standards, ethylvanillin was added prior to CuO oxidation and phenylacetic acid prior to derivatization. The recovery 3. RESULTS AND DISCUSSION of ethylvanillin from MOA samples averaged 92 ± 2%, those of litter samples averaged 100 ± 2%, and those of O 3.1. Time-dependent transformation of mineral–organic samples averaged 115 ± 2%. Cupric oxide treatment pro- associations duces a suite of phenolic monomers derived from intact lig- nin polymers (Benner et al., 1990) and its degradation and 3.1.1. Mineralogical composition at different weathering transformation products (Guggenberger and Zech, 1994; stages Kaiser et al., 2004). The sum of the eight oxidation prod- The mineral composition of LSAG soils varies substan- ucts was taken to indicate lignin-derived OM (Ko¨gel, tially over 4100 kyr of basalt weathering causing significant 1986): vanillyl (V) and syringyl (S) units were calculated transformation of mineral matter (Torn et al., 1997; Chor- from the concentrations of their respectiveP aldehydes, car- over et al., 2004; Ziegler et al., 2005). Despite rock weath- boxylic acids, and ketones:P V = (vanillin + vanillic ering being a continuous process, the most significant acid + acetovanillone), S = (syringaldehyde + syringic transformations of minerals in MOAs with time occur with- acid + acetosyringone). Cinnamyl units (C) were derived in three operationally defined weathering stages (I–III). from the summed concentration of ferulic acid and p-cou- Stage I corresponds to the 0.3-kyr site (Thurston) and is maric acid. Thus, the total content of lignin-derived phenols characterized by rapid weathering of basaltic tephra, thus + (K8)P was determined by the sum of structural units: causing a substantial loss of non-hydrolyzing cations (K , K8= (V + S + C). Na+,Ca2+,Mg2+) and the creation of Al and H+ acidity Non-cellulosic carbohydrates: Sugars were determined as (Chorover et al., 2004). Primary minerals like olivine, their corresponding oximes by gas chromatography after pyroxene, plagioclase (andesine and anorthite) present at acid hydrolysis of MOA samples (Amelung et al., 1996). the 0.3-kyr site (Table 1) are weathered almost completely Briefly, freeze-dried MOAs were weighed into 25-mL before 20 kyr giving rise to metastable minerals that domi- hydrolysis flasks (corresponding to 5–10 mg OC), 10 mL nate during stage II (20–400 kyr). Ferrihydrite/nanocrystal- of 4 M trifluoroacetic acid and 100 lL myo-inositol (inter- line goethite, Al gels, and allophane constitute the nal standard) were added. Samples were reacted for 4 h at dominant mineral phases during phase II of the <2-lm 378 K and then filtered through glass microfibre filter fraction (Chorover et al., 1999) whereas quartz and Ti-bear- (GF6 Whatman). The filtrate was dried on a rotary evapo- ing minerals (anatase, ilmenite) comprise the larger grain rator at 313 K, taken up in 10 mL distilled H2O, and subse- sizes (not shown). A significant portion of Al dissolved quently purified using preconditioned XAD-7 and Dowex from minerals during weathering stage II is held in com- 50Wx8 resins. The freezed-dried eluate was taken up in plexes with OM as shown by the large CuCl2-extractable total 2 mL distilled H2O and transferred into a 3-mL reac- Al concentrations (Table 1). Moreover, at the youngest tivial, and freeze-dried again. Derivatization was accom- and intermediate-aged sites (0.3–400 kyr; stage I and II), a plished by adding 200 lL methyl glucose (derivatization predominant portion of the secondary Fe exists in the form Evolution of mineral–organic associations upon substrate aging 2041 of PC Fe (hydr)oxides as indicated by high Fe reactivity ra- charge minerals. The MOC versus MN correlation coefficient 2 tios (Feox/Fedith: 0.8). In stage III (1400 and 4100 kyr), was close to unity (r = 0.92; p < 0.001; n = 12), suggesting a the metastable mineral phases are progressively converted predominant organic origin of N. Both MOC and MN into more crystalline Fe oxides in case of ferrihydrite (Ko- concentrations increased during weathering stage I and II kee: Feox/Fedith = 0), and into gibbsite, kaolinite and hal- concurrent with the formation of PC mineral phases (Torn loysite in case of Al gels and allophane (Chorover et al., et al., 1997). Assuming that at least in A horizons, the 2004; Vitousek, 2004; Ziegler et al., 2005). Fig. 2 displays MOC contents reflect differences in the capacity of minerals three representative XP scans of mineral–organic associa- and monomeric/polymeric metals to bind OM, the sorption tions, showing that across all weathering stages the suite capacity increased from the youngest site towards a maxi- of detected elements at particle surfaces was comparable. mum at 400 kyr by a factor of 16 and then subsequently de- Aluminum, Si, Fe, and Ti were the major inorganic ele- clined again by 86% towards the oldest site. Hence, the ments, whereas basic cations (Ca, Mg) were only detected MOC accumulation proceeds faster than its decline into later at the 0.3-kyr site (Table EA1). More Fe was present at weathering stages (1400–4100 kyr), which is linked to the ra- the particle surfaces of the 20-kyr and 4100-kyr MOAs pid formation of metastable minerals and metal–organic and a weak but significant correlation between the atom complexes during weathering stage II and the subsequent percentages of Fe and Ti (r2 = 0.36; p < 0.05; n = 12) sug- slow transformation of PC phases into more stable minerals. gests a partial coexistence of Fe and Ti minerals. Across On average 63 ± 9% OC was extracted by 0.1 M NaOH the LSAG, Alox concentrations did not covary with Feox from A-MOAs whereas more OC was liberated from B- 2 but covaried strongly with Siox concentrations (r = 0.94; MOAs (82 ± 19%; Table 1). The larger MOC extractability p < 0.001; n = 12), reflecting the contribution of allophane in subsoil MOAs indicates a larger extent of direct mineral– during stage II. The abundance of PC minerals (allophane, organic bindings compared to the A-MOAs. For A-MOAs, (hydr)oxides) as early products of basalt weathering is, the majority of the extracted MOC was adsorbed to XAD-8 however, not restricted to wet tropical conditions but has resin (Table 1), suggesting a large contribution of aromatic also been observed for a dry soil chronosequence at the substances. In contrast, the released OC from B-MOAs was Canary Islands (Lanzarote) covering substrate ages of slightly less hydrophobic (Table 1). Notably, the fraction of <40 kyr and a mean annual precipitation of only XAD-8 adsorbable MOC was smallest in weathering stage 140 mm (Jahn et al., 1992). III (1400 and 4100 kyr), suggesting that the crystalline min- erals present at the oldest sites are associated with more 3.1.2. Yields and extractability of mineral-associated organic hydrophillic compounds than the minerals at the other sites matter (Table 1). Shifts in mineral composition were closely related to the concentration of MOC and mineral-associated N (MN). 3.1.3. Specific surface area and mesoporosity MOAs contained between 18 and 290 mg MOC/g corre- The specific surface area (SSA) and porosity of MOAs sponding to an average of 79 ± 4% of total OC, and 2– are displayed in Table 2. The A-MOA at the youngest site 19 mg MN/g, averaging 88 ± 4% of total N (Table 1). This (0.3 kyr; weathering stage I) exhibits the lowest SSA and means that the main proportion of OM in LSAG soils is asso- porosity due to the prevalence of primary silicates, consis- ciated with minerals; a consequence of the acidic pH (Table 1) tent with SSA values reported for primary minerals such that favors sorptive interactions between OM and variable as olivine, anorthite, or diopside that not exceed 2m2/g

Fig. 2. Representative X-ray photoelectron spectra of mineral–organic associations from A horizons of the long substrate age gradient. Spectra are stacked by a constant value. 2042

Table 2

Specific surface area (SSA), CBET constant, net enthalpy of N2 adsorption (DHxs), pore volume data, surface enrichment of C and N, and the volumetric ratio of OM and minerals in mineral– 2034–2060 (2009) 73 Acta Cosmochimica et Geochimica / al. et Mikutta R. organic associations from A and B horizons of the long substrate age gradient. corr, a b c d e f Site age SSA CBET DHxs SSA MIV MEV TPV Surface enrichment Volumetric ratio VOM /Vmineral 3 3 N2 CO2 2–10 nm 10–50 nm <200 nm C N qOM = 1.4 g/cm qOM = 1.7 g/cm (kyr) (m2/g) KJ/mol (m2/g) (mm3/g) (mm3/g) A horizons Thurston 0.3 2.3 151 3.3 2.4 1.0 2.5 1.0 4.9 14.8 16.8 8.4 0.08 0.07 Laupahoehoe 20 54.4 141 3.2 84.4 23.5 25.3 32.5 44.7 146.9 1.6 1.3 1.54 0.99 Kohala 150 28.8 123 3.1 46.0 13.0 22.0 13.8 63.2 140.7 1.4 1.4 1.27 0.85 Pololu 400 19.0 72 2.7 45.3 7.5 21.4 20.4 63.7 150.9 1.3 1.2 2.87 1.57 Kolekole 1400 20.2 86 2.9 27.4 8.5 11.0 10.3 52.8 115.6 1.7 1.6 0.76 0.55 Kokee 4100 47.6 195 3.4 51.8 9.5 10.9 21.4 20.0 64.5 5.2 3.9 0.23 0.18 B horizons Thurston 0.3 9.3 206 3.4 9.8 4.0 6.8 4.3 8.8 19.7 9.0 5.5 0.10 0.08 Laupahoehoe 20 62.9 153 3.2 81.4 27.0 52.7 30.0 99.1 255.0 1.9 1.6 0.58 0.43 Kohala 150 55.0 153 3.2 72.4 23.0 55.5 36.0 98.8 194.7 1.7 1.8 0.56 0.42 Pololu 400 65.2 183 3.3 77.9 28.0 29.5 28.5 80.0 221.6 2.3 2.3 0.40 0.31 Kolekole 1400 64.8 179 3.3 67.4 7.5 15.3 33.5 84.4 219.8 5.7 2.3 0.08 0.06 Kokee 4100 70.9 145 3.2 72.7 3.5 7.9 39.4 141.7 265.6 7.2 1.3 0.06 0.05 a SSA corrected for OM content (see Section 2.5). b Micropore volume (<2 nm). c Mesopore volume. d Total pore volume. e Surface enrichment = C or N (XPS)/C or N (elemental analysis). f 3 Assumptions: mOM =2 mOC , qOM = 1.4 or 1.7 g/cm , VOM =mOM /qOM , Vmineral = VMOA VOM , where VMOA = 1000 g/qMOA , qMOA data are listed in Table 1. Evolution of mineral–organic associations upon substrate aging 2043 over a broad range of grain sizes (Brantley and Mellott, despite of the large OM contents, also agrees well with 2000). During 20 kyr of basalt weathering, SSA increased the high points of zero proton charge particular of stage- 23-fold as a result of the neosynthesis of secondary mineral II soils (6 in 1 mM LiCl electrolyte) (Chorover et al., phases. Despite of its large OM content, significant mineral 2004). In principle, mineral surfaces might remain bare of surface area (SSAcorr = 84 m2/g) and micropore volume OM either by being located in microsites inaccessible to 3 (24 mm /g) of the 20-kyr A-MOA is still accessible to N2, OM, e.g., in small aggregates, by electrostatic shielding by confirming that, in the freeze-dried state, OM only incom- pre-existing sorbed OM, or by being chemically unreactive pletely coats the mineral surfaces (Arnarson and Keil, towards OM. 2001). After this initial SSA increase, the SSA of A-MOAs dropped during stage II (20–400 kyr) due to the sorption of 3.1.4. Importance of mineral microporosity OM, which is reflected by declining CBET constants (Table In soils, mineral micropores with diameters <2 nm can 2). In fact, the MOC concentrations correlated negatively arise from crystal defects, the stacking edges of expandable 2 with the CBET constants (r = 0.58; p < 0.01; n = 12), con- 2:1 clay minerals, dead-ends of pores at domain boundaries firming that natural OM when sorbed to mineral surfaces of Fe (hydr)oxides or from interparticle spaces of PC miner- lowers the enthalpy of N2 adsorption (Kaiser and Guggen- als (Cabrera et al., 1981; Cornell and Schwertmann, 1996; berger, 2003; Mo¨dl et al., 2007; Wagai et al., 2009). The lar- Fischer et al., 1996; Aringhieri, 2004). The close proximity corr ger CBET constants and SSA values of B-MOAs (Table of pore walls render micropores favorable sorption sites for 2) reflect their smaller OM input and hence smaller OM OM because of the high density of hydroxyl groups and the contents in comparison to the A-MOAs. Finally in weath- ability of OM to sorb to mineral surfaces through multiple ering stage III (1400–4100 kyr), the SSA and the CBET con- bonds (Kaiser and Guggenberger, 2003). The N2- and CO2- stants of A-MOAs increased again because of the detectable micropore volumes of MOAs are displayed in Ta- prevalence of minerals with smaller amounts of bound ble 2. The CO2-accessible micropore volume exceeded the OM. Freeze-dried A-MOAs from stage-II sites contained N2-derived micropore volume in all MOAs by a factor of significantly more volume in <200-nm pores than MOAs 1.1–2.9 (Table 2). This can be explained by the molecular from endmember soils, indicating a particularly small pore sieving of N2 at 77 K, whereas CO2 at 273 K is capable of dif- geometry of associations between OM and PC mineral fusing through the flexible organic network into mineral phases. In contrast to A-MOAs, in B horizons the SSA, micropores (Kwon and Pignatello, 2005; Mikutta and Mi- CBET constants, and pore volume of MOAs show less var- kutta, 2006; Kaiser et al., 2007; Mikutta et al., 2008). The iation (Table 2), but the overall pattern still reflects the CO2-micropore volumeP of MOAs correlated positively with cumulative effect of mineral alteration and OM sorption. oxalate-extractable (Fe, Al, Si) concentrations (r2 = 0.96; The large CBET constants of all MOAs (Table 2) provide p < 0.001; n = 12) whereas MOC failed to predict the CO2- 2 evidence that significant mineral surface area is bare of OM micropore volume (r = 0.14). Consequently, CO2 particu- (Wagai et al., 2009). More extensive OM coatings/sur- larly probed micropores of PC mineral phases such as Fe roundings as indicated by smaller CBET values (72, 86) ex- and Al (hydr)oxides and allophane rather than those of isted in A-MOAs of the 400-kyr and 1400-kyr site. OM. Similarly, Kaiser et al. (2007) showed that the CO2 sorp- Compared to other soils (Wagai et al., 2009), these CBET tion at 273 K to goethite and ferrihydrite is largely indepen- constants were still high owing to the presence of uncov- dent from the amount of sorbed OC even at extremely high ered, microporous minerals. In accordance with the large OC concentrations comparable to the maximal OC concen- CBET constants, also the atom% contribution of C to the trations in the MOAs studied (300 mg OC/g). Also for syn- bulk XP spectra of MOAs (hydrogen ignored) was <50% thetic coprecipitates, the reduction of the mineral micropore (14–49%; not shown) and also revealed that less C is asso- volume was much less intense for CO2 than for N2 if the OM ciated with particle surfaces in B-MOAs compared to content was increased (Mikutta et al., 2008). Therefore, the A-MOAs (Table EA1). Additionally, we adopted the difference between the N2- and CO2-micropore volumes in empirical approach of Mayer (1999, 2005) which relates the MOAs, particularly during weathering stage II, suggests the difference in the net N2 adsorption enthalpy between that OM bound to PC minerals partly cloggs the micropore bare and organically coated mineral surfaces [D(DHxs)] to entrances (Mikutta and Mikutta, 2006). Hence, at the vicin- the OM fractional coverage according to D(DHxs)=DHbare ity of micropores, OM is ‘sterically’ stabilized by multiple 0.535 DHcoated = 1.08 (fractional coverage) , where DHbare bonds to the mineral surface (Kaiser and Guggenberger, denotes the N2 adsorption enthalpy for the bare mineral 2007b), which would reduce OM desorption and thus biodeg- surfaces and DHcoated for the organically coated ones. radation (Kaiser and Guggenberger, 2003, 2007b; Mikutta Due to the dominating oxide-type mineral composition, et al., 2007). Consequently, the association of OM with 2-line ferrihydrite served as energetic reference for a blank microporous mineral phases particularly in stage II is an mineral surface (CBET = 239 and DHbare = RT ln CBET = additional factor that favors the accumulation and stabiliza- 3.5 KJ/mol). The assumption inherent to this approach is tion of OM in the LSAG soils. that all mineral surfaces are energetically homogeneous, which strictly is not valid for natural microporous samples. 3.1.5. Enrichment of organic matter at mineral surfaces and However, when using this approach, one would end up with volumetric ratios no more than 50% of the N2-accessible mineral surface area Table EA1 compiles the XPS-based element composi- coated by OM – in line with the high CBET constants and tion of MOA surfaces. Despite the low to moderate OM XPS data. The presence of significant inorganic surfaces, coverages, the surfaces of MOAs appear enriched with 2044 R. Mikutta et al. / Geochimica et Cosmochimica Acta 73 (2009) 2034–2060

OM as evidenced by the large OC concentrations (88– on particle surfaces, i.e., Vmineral >> VOM. In contrast, in 372 mg OC/g MOA) and MN concentrations (2–23 mg A-MOAs from intermediate-aged sites (stage II), VOM MN/g MOA; Table EA1). The XPS-based surface MOC was either close to or even larger than Vmineral (Table 2). concentrations are comparable with those observed for goe- The large volume fraction of sorbed OM means that adsorp- thite reacted with dissolved OM in the sorption plateau re- tion processes alone are insufficient to explain the OM accu- gion (276 mg OC/g; Kaiser and Guggenberger, 2007b) and mulation in A-MOAs during weathering stage II. More with the maximum concentrations observed in different likely, the data imply that between 20 and 400 kyr, where soils including Podzols, Andisols, Cambisols, and a Luvisol the pH was lowest and the CuCl2-exchangeable Al was larg- (Yuan et al., 1998; Gerin et al., 2003). The surface C and N est (Table 1), the precipitation of OM and the subsequent enrichment was calculated as the ratio between XPS-de- accretion of mineral particles or the precipitation of OM with rived MOC and MN concentrations and those derived from multivalent metals (Al, Fe) or hydroxides play a significant elemental analysis. Values >1 correspond to an enrichment, role in the formation of MOAs (Chorover et al., 2004). Note- values <1 to a depletion of MOC and MN at particle sur- worthy, the VOM/Vmineral ratios of MOAs were larger at low- faces relative to the bulk concentration. Surface enrichment er soil pH (Fig. 3), possibly because charge neutralization by values of MOC and MN ranged between 1.3–16.8 and 1.2– H+ or low-molecular Al or Fe species favors the precipitation 8.4, respectively (Table 2). MOC and MN were always en- of natural OM (Scheel et al., 2008). It is important to realize riched at mineral surfaces with the strongest enrichment at that the OM in stage-II A-MOAs is effectively stabilized de- the youngest (stage I) and oldest site (stage III) (Table 2). spite of seemingly less direct mineral–organic bondings in- The surface C and N enrichment in LSAG endmember volved. This may, in addition to the stabilization derived MOAs reflect a larger volume ratio of minerals to OM, from adsorption of OM to microporous PC minerals, also re- or, in other words, that larger mineral grains are coated sult from the fact that OM in precipitated flocs (i) is more re- with OM. In contrast, in surface horizons of intermedi- calcitrant due to its high aromatic C and low N content ate-aged sites, the smaller surface C and N enrichment as (Scheel et al., 2007) and (ii) is partly inaccessible for microor- well as the larger volume contained in nanometer-sized ganisms due to the flocs’ small pore sizes and the high degree pores (Table 2) imply that MOAs consist of a homogeneous of metal-C cross-linkings (Scheel et al., 2008). In the respec- array of PC minerals and OM. Volumetric ratios between tive B-MOAs, the volumetric ratios were generally smaller sorbed OM (VOM) and the mineral phase (Vmineral) were than in A-MOAs, with Vmineral always clearly exceeding calculated to illustrate the potential mode of OM accumu- VOM, thus implying that adsorptive processes are more rele- lation. Here, we utilized the pycnometric q data of MOAs vant in subsoil than in OM-rich surface soil horizons. (Table 1) and assumed qOM to vary between 1.4 and 1.7 g/cm3, thus giving also credit to the potentially higher 3.1.6. Transmission electron microscopy density of sorbed OM (Kaiser and Guggenberger, 2007b). The six A-MOA samples were imaged and analyzed Table 2 depicts the VOM/Vmineral ratios, which range be- by TEM and EDS to reveal potential differences in tween 0.05 and 2.87 depending on the qOM assumed. In MOA structures across the LSAG. For each MOA sam- endmember A-MOAs (0.3 and 4100 kyr), volumetric ratios ple, a reasonable number of particles (10 < n < 15) of var- were small (0.07–0.23) and consistent with the C surface ious dimensions were analyzed in order to investigate a enrichment data, thus favoring the ‘coating model’ of OM representative fraction of MOAs for each site. At the

Fig. 3. Plot of the volumetric ratio of sorbed OM (VOM) versus mineral matter (Vmineral) in mineral–organic associations from A (A-MOA) and B horizons (B-MOA) of the long substrate age gradient. Volumetric ratios were calculated based on the following assumptions: 3 mOM =2 mOC, qOM = 1.4 g/cm , VOM =mOM/qOM, Vmineral = VMOA VOM, where VMOA = 1000 g/qMOA. The corresponding qMOA values are given in Table 1. Evolution of mineral–organic associations upon substrate aging 2045

0.3-kyr site, we found massive aggregates (10–20 lm) tens of microns). Additionally, film-like amorphous struc- consisting of primary mineral particles that were partly tures rich in Fe, partly in Ti and clay minerals (Si, Al) coated with amorphous-like inorganic phases (Fig. 4A) were observed (Fig. 4D). Again, almost no carbon was enriched in Fe, Al and Si. Only small amounts of C were detected in these PC structures, confirming the relatively detected by EDS which confirms that OM is not present small fractional OM coverage despite of the large OM as uniform coating (not shown). Consistent with the larg- content. However, as this MOA fraction likely contains est Feox concentration, at Laupahoehoe (20 kyr) we ob- precipitated OM, we cannot rule out that OM was partly served strong signals of Fe, also to a lesser extent of lossed during the ultrasonic pretreatment. At the 1400- Ti, associated with PC Fe (hydr)oxides, likely ferrihy- kyr site, we observed more heterogeneous aggregates, drite, that covered larger primary particles and aggregates consisting of mixtures of crystalline aluminosilicates and (Fig. 4B). Interestingly, almost no carbon coexisted with oxides (Fig. 4E), as well as Si-depleted film-like struc- these phases. At the 150-kyr site (Kohala), the imaged tures. At the most weathered 4100-kyr site, again massive aggregates generally had a smaller size (1–2 lm), were aggregates were present, whose surfaces appeared to be more homogeneous, and contained more Al and Si rela- covered by aggregates of crystalline Fe (hydr)oxides tive to the 20-kyr site (Fig. 4C). At the 400-kyr site, (Fig. 4F). In summary, during weathering stage II, clearly where the concentration of MOC and the OM coverage PC phases dominate at the nanoscale. At weathering was maximal, massive aggregates were observed (several stage III, secondary clay minerals and crystalline Fe

Fig. 4. Representative transmission electron micrographs and X-ray electron dispersive spectra of mineral–organic associations from A horizons of the long substrate age gradient. Primary mineral particles at the 0.3-kyr site are partly covered with phases of poor crystallinity (A) while such poorly crystalline phases dominate the 20-kyr (B), 150-kyr (C) and 400-kyr site (D). Note, the presence of film-like structures at the 400-kyr site (black arrow). At the 1400-kyr (E) and 4100-kyr (F) site, aggregates are composed to a much larger extent by secondary crystalline aluminosilicates and Fe (hydr)oxides, the latter particularly covering larger mineral particles at the 4100-kyr site. 2046 R. Mikutta et al. / Geochimica et Cosmochimica Acta 73 (2009) 2034–2060

(hydr)oxides were abundant, the latter particularly cover- ylic C; the two latter being comparably abundant (3.2–6.3% ing larger aggregates. versus 4.7–7.1%). The average speciation of MOC in A horizons followed the sequence: aromatic C > ether/alcohol 3.2. Time-dependent chemical alteration of sorbed organic C>a-C > aliphatic C > ketonic/aldehyde C > amide C = matter carboxylic C >> p–p*, suggesting the prevalence of pheno- lic substances (e.g., lignin, polyphenols, tannins) and 3.2.1. X-ray photoelectron spectroscopy carbohydrates. This structural analogy was interrupted at XPS was used as a non-destructive surface-probing tool the 20-kyr site, where aromatic C was significantly depleted that provides in-situ information about the chemical state in favor of aliphatic C and a-C (Table 3). The only of OM at particle surfaces. The chemical composition of moderate impact of minerals on the chemical variability MOM was revealed by studying the effect of O, C, and N of A-horizon MOM derives from the large input of fresh on XPS C1s binding energies. Peak deconvolution was per- C from a similar source (M. polymorpha). In contrast, the formed similar to Monteil-Rivera (2000) because this fitting chemical composition of OM at particle surfaces in B hori- routine allows distinction between (i) amide C and carboxyl zons was more variable than in A horizons, in particular C (not possible by 13C NMR) and (ii) aromatic and ali- with respect to aliphatic and aromatic C moieties but fol- phatic C, and (iii) yielded comparable quantities of car- lowed no clear trend through weathering stages I–III (Table boxyl groups in humic acids as the wet chemical analysis. 3). The order of abundance was: aromatic C > a-C = ether/ Spectral shifts in core level C1s binding energy were as- alcohol C > aliphatic C > carboxylic C > ketonic/aldehyde signed to different chemical environments of C: (1) unsub- C > amide C >> p–p*. Whereas the average fraction of stituted aromatic carbon (C@C; 284.4 eV), (2) aliphatic amide C in A- and B-MOAs was similar, the B-MOAs were carbon (CAC/CAH; 285 eV), (3) a-carbon (CAC(O)O; slightly enriched in carboxylic C. This suggests a larger de- 285.6 eV), (4) ether or alcohol carbon (CAO; 286.4 eV), gree of oxidative alteration of subsoil MOM and/or a (5) ketonic or aldehyde carbon (C@O; 287.7 eV), (6) amide stronger retention of acidic OM components. carbon (C(O)N; 288.5 eV), (7) carboxylic carbon (C(O)O; To test for a mineralogical effect on MOM’s chemical 289.3 eV) and (8) p–p*-shake-up satellite from the aromatic composition, we related the spectral abundance of each car- structure (290.8 eV). Examples of high-resolution XPS C1s bon type (in atom%) to the XPS-based inorganic surface spectra across the LSAG MOAs are shown in Fig. 5. De- composition of MOAs (n = 12; normalized based on the spite the major mineral transformations during weathering sum of Fe, Al, Si and Ti). We found that aliphatic C corre- stages I–III there were only moderate differences in the lated positively (r2 = 0.72; p < 0.001) with surficial Fe con- average chemical composition of MOM in A horizons (Ta- centrations whereas carboxyl C correlated positively with ble 3). All MOAs were rich in aromatic C (29–34%) and surficial Al concentrations (r2 = 0.78; p < 0.001). Alcoholic contained relatively small amounts of amide C and carbox- C, mainly originating from carbohydrates, correlated

Fig. 5. Examples of high-resolution XPS C1s spectra of mineral–organic associations from A horizons of the 0.3-kyr, 150-kyr, and 4100-kyr site of the long substrate age gradient. Numbers of deconvoluted sub-peaks refer to (1) unsubstituted aromatic carbon (ArACAC/ArACAH), (2) aliphatic carbon (CAC/CAH), (3) a-carbon (CAC(O)O), (4) ether or alcohol carbon (CAO), (5) ketonic or aldehyde carbon (C@O), (6) amide carbon (C(O)N), (7) carboxylic carbon (C(O)O), and (8) p–p*-shake-up satellite from the aromatic structure. Evolution of mineral–organic associations upon substrate aging 2047

Table 3 X-ray photoelectron spectroscopy-based speciation of carbon binding environments in mineral–organic associations from A and B horizons of the long substrate age gradient (in atom%): unsubstituted aromatic carbon (ArACAC/ArACAH), aliphatic carbon (CAC/CAH), a-carbon (CAC(O)O), ether or alcohol carbon (CAO), ketonic or aldehyde carbon (C@O), amide carbon (C(O)N), carboxylic carbon (C(O)O), and p– p*-shake-up satellite from the aromatic structure. Values in parenthesis refer to the standard error of the mean of triplicate measurements; values following the slash (/) refer to the average full-width-at-half-maximum (FWHM) of the deconvoluted peak. CV connotes the coefficient of variation. Site ArACAC CACCAHCAC(O)O CAOC@O C(O)N C(O)O p–p* age ArACAH (kyr) 284.4 eVa 285.0 eV 285.6 eV 286.4 eV 287.7 eV 288.5 eV 289.3 eV 290.8 eV A horizons Thurston 0.3 33.3 (0.8)/1.5 7.3 (0.5)/0.7 13.1 (1.9)/1.0 26.0 (2.7)/1.5 9.6 (1.3)/1.3 3.4 (0.9)/1.1 7.1 (0.1)/1.5 0.3 (0.1)/0.6 Laupahoehoe 20 26.0 (1.6)/1.5 20.3 (0.3)/1.2 21.8 (1.3)/1.5 13.8 (0.7)/1.5 6.3 (1.3)/1.3 6.3 (0.6)/1.2 5.5 (0.5)/1.4 0.1 (0.0)/0.5 Kohala 150 32.6 (1.1)/1.5 8.5 (1.1)/0.9 16.7 (1.2)/1.1 22.8 (0.8)/1.5 9.1 (1.2)/1.4 4.3 (1.3)/1.2 5.9 (0.9)/1.5 0.1 (0.1)/0.6 Pololu 400 33.6 (1.5)/1.5 8.5 (0.8)/1.5 15.3 (0.9)/1.5 24.1 (0.7)/1.5 9.3 (0.3)/1.5 4.5 (1.0)/1.5 4.7 (0.3)/1.5 0.1 (0.0)/1.5 Kolekole 1400 32.7 (5.1)/1.5 7.5 (1.8)/0.8 18.7 (2.2)/1.1 23.0 (2.6)/1.5 9.5 (2.2)/1.4 3.2 (1.3)/1.0 5.4 (0.1)/1.5 0.0 (0.0)/0.6 Kokee 4100 28.6 (2.1)/1.5 13.5 (2.8)/1.0 13.1 (0.1)/1.0 25.4 (1.0)/1.5 7.5 (0.9)/1.2 5.8 (1.6)/1.2 5.8 (0.6)/1.5 (0.1)/0.6 Mean 31.1 10.9 16.4 22.5 8.5 4.6 5.7 0.2 CV (%) 10.0 46.9 20.7 19.9 15.7 27.2 13.7 80.2 B horizons Thurston 0.3 34.1 (1.7)/1.5 9.5 (1.3)/0.9 22.1 (2.0)/1.3 17.3 (1.5)/1.5 5.8 (1.6)/1.2 4.2 (0.8)/1.0 6.5 (0.5)/1.5 0.5 (0.1)/1.2 Laupahoehoe 20 27.0 (1.9)/1.5 14.1 (0.8)/1.1 22.5 (1.1)/1.5 14.4 (0.6)/1.5 5.5 (0.5)/1.4 7.3 (0.4)/1.4 8.8 (0.3)/1.5 0.4 (0.2)/0.9 Kohala 150 39.8 (4.0)/1.5 5.0 (1.8)/0.6 15.6 (1.1)/1.0 19.9 (1.3)/1.5 6.0 (0.5)/1.1 2.7 (0.5)/0.8 10.0 (0.7)/1.5 1.0 (0.4)/1.0 Pololu 400 32.5 (2.7)/1.5 11.0 (3.3)/0.9 16.7 (1.8)/1.2 17.9 (0.9)/1.5 5.4 (0.3)/1.2 8.2 (0.6)/1.3 8.3 (0.4)/1.5 0.0 (0.0)/0.5 Kolekole 1400 39.7 (4.3)/1.5 6.2 (3.4)/0.6 17.1 (0.8)/1.0 18.8 (1.1)/1.5 5.6 (0.1)/1.0 3.1 (0.8)/0.8 8.8 (0.5)/1.5 0.8 (0.2)/1.1 Kokee 4100 26.2 (2.7)/1.5 18.9 (2.6)/1.1 15.9 (0.6)/1.0 16.6 (0.8)/1.3 8.4 (0.7)/1.4 3.5 (0.3)/1.1 9.4 (0.6)/1.5 1.2 (0.5)/1.3 Mean 33.2 10.8 18.3 17.5 6.1 4.8 8.6 0.6 CV (%) 17.8 48.0 17.1 10.9 18.3 48.1 13.9 64.7 a Binding energy. positively with Si concentrations (r2 = 0.69; p < 0.001). that acidic, UV-absorbing aromatic components of natural These correlations suggest an impact of the bulk inorganic OM were preferentially adsorbed to PC Al(OH)3 or precip- composition of MOAs on the chemical composition of itated with Al and, once adsorbed/precipitated, possessed a MOM but, due to the source inspecificity of XPS, they do remarkable resistance against biodegradation (Scheel et al., not indicate causation. Therefore, we transformed the 2007). This stability arose from two factors: the strong atom% concentrations of C types into their absolute con- chemical binding of OM to the inorganic phase (thus limit- centrations (mg C type/g MOA) and used the more specific ing desorption) and the a priori recalcitrant structure of selective chemical extraction data for regressions, being aromatic molecules. Analogously, we rationalize that beside aware that the elemental composition of the bulk sample the association of OM with microporous, high-SSA miner- and their surfaces can differ. When evaluating A- and B- als or with monomeric/polymeric metal species also the MOAs independently, we observed that in both A and B inherent stability of such acidic aromatic compounds assists horizons aliphatic and amide C correlated positively with in the stabilization of OM. 2 Feox concentrations (A horizons: r = 0.94; 0.67; B hori- zons: r2 = 0.90; 0.94; n =6;Fig. 6). This surprising correla- 3.2.2. XPS ‘depth profiles’ tion probably points to a preferential association of PC Fe To examine the composition of deeper MOA layers and phases with microbial remnants and/or to the biogenetic the binding state of carbons, a representative area origin of PC Fe phases in these moist LSAG soils but this (500 lm2) of A-MOAs was analyzed after Ar+ sputtering. hypothesis still needs to be ascertained. Due to the only Note that the XPS elemental composition of powdered moderate differences in OM composition in A-MOAs no MOA samples following Ar+ sputtering represents an inte- further relationship between the concentration of other C gral over all exposed, rough particle surfaces, which inevita- types and mineralogical properties was observed. In con- bly includes variable erosion effects for different kinds of trast, in the B-MOAs, aromatic C was more abundant in materials. Fig. 7 shows that signals from Si and Al, pre- samples rich in PC aluminosilicates and Al-organic com- dominantly originating from aluminosilicates, emerged plexes as revealed by positive relationships with Alox,AlCu, most markedly during sputtering in almost all A-MOAs. 2 and Siox (r = 0.92; 0.96; 0.83; Fig. 6). The same holds true In contrast, at the 20-kyr and 4100-kyr site, Fe signals from for carboxylic C (r2 = 0.98; 0.88; 0.85; Fig. 6). Although Fe phases increased most substantially with depth, consis- only revealed for MOAs from the B horizons, the data sug- tent with the larger Fe (hydr)oxide content of these MOAs gest that, on a mass-basis, particularly PC aluminosilicates (Table 1). At t30, surficial OC concentrations were still >13 and Al phases favor the retention of acidic aromatic com- atom% (P64 mg OC/g MOA), suggesting the presence of pounds. Guan et al. (2006) and Scheel et al. (2007) observed thick organic coatings/agglomerations. At the 400-kyr site 2048 R. Mikutta et al. / Geochimica et Cosmochimica Acta 73 (2009) 2034–2060

Fig. 6. (A–C) Regression plots of XPS-based carbon types with oxalate-extractable Fe, Al, and Si concentrations in mineral–organic associations from A (A-MOA) and B horizons (B-MOA). Note that the oxalate-extractable metal concentrations refer to bulk MOA samples whereas XPS data reveal the average composition of OM at particle surfaces.

Fig. 7. Change of the inorganic surface composition (Al, Si, Ti and Fe) and of C and N concentrations in mineral–organic associations from A horizons upon Ar+ sputtering. The data for C and N represent the percentage loss of C and N concentrations after 30 min of etching. Evolution of mineral–organic associations upon substrate aging 2049 where the MOC concentration and the volumetric OM– structure (Zaporojtchenko et al., 2007). We therefore exam- mineral ratio were largest, the XPS C1s signal decline was ined OM extracted from the 0.3-kyr MOA by means of much less intense, again revealing that MOM exists in a NaOH–NaF in order to test for the effect of Ar+ irradiation particularly homogeneous mineral–organic microfabric. In on the chemical composition of LSAG-specific natural OM. contrast, at the two oldest sites (stage III), thinner organic In line with previous work, we found that Ar+ sputtering coatings/agglomerations are indicated by a more pro- for 5 min reduced the atom% abundance of aromatic C nounced decline of the C concentration (Fig. 7). by 4.5% while increasing the concentration of aliphatic C While the sample number is limited, the XPS data reveal by 6.2% (Fig. 9). A general decrease of aromatic C in consistent depth trends in all examined MOAs. These Ar+-treated polymers and natural OM, however, was not changes can be followed in Fig. 8, showing the alteration observed in our MOA samples, but rather an increase (Ta- of C functionalities in the 20-kyr A-MOA upon sputtering. ble 4). This discrepancy, therefore, points to vertical differ- Ar+ etching generally caused (i) a roughly parallel decrease ences in OM composition at mineral particles rather than to of surface C and N concentrations (Fig. 7), (ii) a relative in- Ar+-induced chemical modifications. The increasing aro- crease of the signal at 284.4 eV reflecting aromatic C, and matic C content with depth agrees well with the idea of aro- (iii) a concomitant relative decrease of signals in the binding matic compounds possessing a high affinity to oxide-type energy range 287.7–289.3 eV, connoting mainly ketonic/ minerals (Kaiser and Guggenberger, 2000; Chorover and aldehyde, amide, and carboxylic C, respectively (Table 4). Amistadi, 2001; Mikutta et al., 2007; Scheel et al., 2008), Also alcoholic C, mainly comprising carbohydrates, tended hence, with their larger propensity to accumulate more di- to decrease with sputter duration (except from the 1400-kyr rectly at the mineral–organic interface. Of note, carboxylic sample). Since abrasion is not likely to destroy the microag- C declined slightly with depth (Table 4), possibly indicating gregate structure, the observed shifts in C functional groups their preferential removal upon Ar+ treatment (Zap- with sputter time reflects differences in the molecular struc- orojtchenko et al., 2007). ture of MOM. It is known that Ar+ gassing under ultra- In line with the strong MOC–MN correlation in MOAs, high vacuum conditions can induce chemical modifications the parallel decrease in C and N concentration confirms of synthetic polymers such as the selective removal of phe- that OM bound to minerals is the major source of organic nyl rings (Zaporojtchenko et al., 2005, 2007). Also, poly- N(Yuan et al., 1998). Moreover this result demonstrates mers exposed to Ar plasma revealed a decrease in that in LSAG A-MOAs, organic N does not become en- aromatic C in favor of aliphatic C (France and Short, riched towards the mineral–organic interface relative to 1998). Beside polymer scission and the formation of OC. Hence, this finding does not support the recently pro- cross-linkings, ion treatment can also cause the removal posed conceptual model of mineral–organic associations of oxygen-functional groups depending on the polymer’s which states that amphiphilic nitrogenous compounds are

Fig. 8. Depth-dependent change of the chemical composition of mineral-associated OM (20-kyr site; A horizon) upon Ar+ etching for 15 and 30 min. Numbers of deconvoluted sub-peaks refer to (1) unsubstituted aromatic carbon (ArACAC/ArACAH), (2) aliphatic carbon (CAC/ CAH), (3) a-carbon (CAC(O)O), (4) ether or alcohol carbon (CAO), (5) ketonic or aldehyde carbon (C@O), (6) amide carbon (C(O)N), (7) carboxylic carbon (C(O)O), and (8) p–p*-shake-up satellite from the aromatic structure. 2050 R. Mikutta et al. / Geochimica et Cosmochimica Acta 73 (2009) 2034–2060

Table 4 Changes of the C binding environments in mineral–organic associations from A horizons in the assigned spectral regions (see Table 3) upon + 2 Ar sputtering for 15 min (t15) and 30 min (t30). The analyzed area corresponded to 500 lm . The D notation corresponds to percent changes between t0 and t30. (Atom%) Thurston Laupahoehoe Kohala

t0 t15 t30 D t0 t15 t30 D t0 t15 t30 D ArACAC, ArACAH 35 44 42 +19 22 42 46 +107 33 45 38 +14 CAC, CAH 2 4 3 +66 14 7 7 51 6 2 5 20 CAC(O)O 13 18 21 +62 21 23 21 2 16 21 22 +36 CAO28201931 20 14 15 25 24 21 24 0 C@O116828 7 6 4 40 11 5 6 51 C(O)N 5 1 1 88 8 2 2 68 2 1 2 30 C(O)O 6 6 4 26 7 5 4 45 8 5 4 48 p–p* 0 1 2 +>100 0 0 1 +69 0 1 1 +>100

(Atom%) Pololu Kolekole Kokee

t0 t15 t30 D t0 t15 t30 D t0 t15 t30 D ArACAC,ArACAH 35 46 49 +41 34 47 51 +50 25 57 42 +66 CAC, CAH63353 8 3 2 72 19 1 7 62 CAC(O)O 18 18 15 16 20 21 14 29 13 12 17 +34 CAO23212111 20 14 22 +12 24 18 21 10 C@O104549 10 12 4 53 9 5 5 41 C(O)N 3 3 3 12 5 0 1 73 4 2 1 66 C(O)O 5 3 4 21 4 4 4 264521 p–p* 0 1 1 +>100 0 0 0 0 0 1 1 +>100

Fig. 9. Effect of Ar+ sputtering (5 min) on the chemical composition of freeze-dried OM extracted from the 0.3-kyr MOA (A horizon) by means of NaOH–NaF. Numbers of deconvoluted sub-peaks refer to (1) unsubstituted aromatic carbon (ArACAC/ArACAH), (2) aliphatic carbon (CAC/CAH), (3) a-carbon (CAC(O)O), (4) ether or alcohol carbon (CAO), (5) ketonic or aldehyde carbon (C@O), (6) amide carbon (C(O)N), (7) carboxylic carbon (C(O)O), and (8) p–p*-shake-up satellite from the aromatic structure.

preferentially sorbed to mineral surfaces to initiate a verti- 3.2.3. Absolute concentrations of lignin-derived phenols cal differentiation (multilayering) of organic coatings (Kle- Beside cellulose and hemicelluloses, lignin is the most ber et al., 2007). important biopolymer added to soil by plants and the ma- Evolution of mineral–organic associations upon substrate aging 2051 jor source of phenolic compounds in soil OM. The contri- ceeded the increase in MOC concentrations (Table 5). In bution of lignin-derived phenols to organic soil horizons contrast, K8 contents in B-MOAs were about one order and MOAs across the LSAG was analyzed by alkaline of magnitude smaller than in A-MOAs and remained at a CuO oxidation (Table EA2). Even though the organic lay- low level through all weathering stages (Fig. 10A). Beside ers (L and O) at the six sites contained similar amounts of by adsorption of lignin-rich OM to minerals, the pro- lignin-derived phenols (K8), the concentration of K8in nounced increase of K8 towards the 400-kyr site may mirror MOAs was highly variable and closely matched the accu- to some extent the precipitation of forest floor leachates mulation pattern of MOC as shown by a positive correla- rich in aromatic C moieties. This would accord with the 2 tion between K8 and MOC (r = 0.81; p < 0.001; n = 12). high VOM/Vmineral ratios of stage-II A-MOAs (Fig. 3)as No relationship with mineral phase variables was observed, well as with the finding that carboxyl-rich, UV-absorbing indicating no source-specific retention of lignin residues. In aromatic compounds of dissolved OM have a high propen- A-MOAs, K8 increased steadily during weathering stage I sity to precipitate with Al (Scheel et al., 2007, 2008). and II, reaching a maximum at Pololu (400 kyr), and decreasing thereafter into stage III (Fig. 10A). The more 3.2.4. Lignin-derived phenols in mineral-associated organic favorable accumulation of lignin-derived phenols in A- matter and the degree of alteration MOAs during stage I and II is revealed by the increase in The contribution of lignin-derived phenols to MOC in A K8 concentrations towards the next older site, which ex- horizons also increased steadily towards the 1400-kyr site

Fig. 10. (A–F) Time-dependent change of lignin-derived phenol concentrations (K8) in organic layers and mineral–organic associations from A (A-MOA) and B horizons (B-MOA) (A); acid-to-aldehyde ratio of vanillyl (B) and syringyl units (C). The grey bands in B and C connote the acid-to-aldehyde ratios in collected plant source materials. Non-cellulosic carbohydrate concentrations (D); the GM/AX ratio (E) and RF/AX ratio (F). The GM/AX ratio refers to the mass quotient of (galactose + mannose)/(arabinose + xylose) and that of the RF/AX ratio to (rhamnose + fucose)/(arabinose + xylose). Roman numbers indicate the different weathering stages (see text). 2052 R. Mikutta et al. / Geochimica et Cosmochimica Acta 73 (2009) 2034–2060

Table 5 of side chain oxidation as evidenced by markedly increased Factorial change of organic carbon (MOC), lignin phenols (K8), (ac/al)V and (ac/al)S ratios (Fig. 10B and C). In addition to and non-cellulosic carbohydrate concentrations (on mass-basis) in microbial alteration, the high acid-to-aldehyde ratios in mineral–organic associations towards the next older site. MOAs can derive from the selective retention of acidic lig- MOC Lignin phenols (K8) Carbohydrates nin compounds by charged mineral surfaces (Gru¨newald A horizons et al., 2006; Hernes et al., 2007) or from the preferential pre- 0.3 ? 20 kyr 9.5 12.8 8.6 cipitation of acidic aromatic substances (Scheel et al., 2007, 20 ? 150 kyr 1.1 1.4 0.9 2008). The latter aspects can explain the larger absolute lig- 150 ? 400 kyr 1.5 1.8 1.5 nin-derived phenol concentrations as well as the larger con- 400 ? 1400 kyr 0.5 0.5 0.7 tribution of lignin phenols to MOM during weathering 1400 ? 4100 kyr 0.3 0.2 0.1 stage II. However, the generally high degree of lignin side B horizons chain oxidation persisting through time suggests that in 0.3 ? 20 kyr 4.9 1.3 1.1 particular acidic lignin components are stabilized by inter- 20 ? 150 kyr 1.1 0.9 0.2 action with soil minerals and hydrolyzable metals. Hence, 150 ? 400 kyr 0.7 1.0 1.1 the long-term mineralogical shift primarily affects the abso- 400 ? 1400 kyr 0.2 0.2 0.7 lute concentration of mineral-associated lignin phenols 1400 ? 4100 kyr 0.6 0.5 0.1 rather than their ‘diagenetic’ state.

3.2.5. Contribution of hydroxybenzenecarboxylic acids and then dropped off at the oldest site (Fig. 11A). This Beside the lignin phenol monomers, a suite of hydroxy- means that during weathering stage I and II not only did benzenecarboxylic acids (BA) was detected as CuO oxida- the total concentration of lignin-derived phenols increase tion products (Table EA2; Fig. 11A and B). These over time but also the relative contribution of lignin-derived compounds comprise between 23–39% and 50–255% of phenols to MOM increased. This finding corroborates that MOC-normalized K8 in A- and B-MOAs, respectively. in weathering stage II (20–400 kyr) lignin-derived phenols However, the origin of most of these compounds in soils are preferentially stabilized in associations with PC miner- is not specific. Dihydroxybenzoic acid (3,5di-H BA) may als or monomeric/polymeric metal species. The minor con- derive from polyhydroxyaromatic tannins or other flavo- tribution of lignin phenols to total MOC in A-MOAs at the noids (Gon˜i and Hedges, 1995; Louchouarn et al., 1999; 0.3-kyr site can be explained by the larger abundance of Dickens et al., 2007) but is also considered as a degradation low-SSA primary minerals (olivine, pyroxene and feldspar), product of lignin (Prahl et al., 1994). All detected BAs were which do not favor the selective retention of lignin-derived already liberated from source vegetation leaves and are phenols in comparison to other OM components like car- present in both A- and B-MOAs (Table EA2). The BAs ex- bohydrates (see Section 3.2.7). Except from the 0.3-kyr site, tracted from live leaf tissues of M. polymorpha upon CuO the MOC-normalized K8 yields of B-MOAs were 87–92% oxidation likely originate from its polyphenol/tannin com- smaller than in A-MOAs, comprising only between 0.9 ponents (Ha¨ttenschwiler et al., 2003). The BA/K8 ratio, and 1.6 mg/g of MOC (Table EA2). We do not think that including 3,5di-H BA, increased exponentially from source the smaller subsoil K8 yields are caused by a less efficient vegetation leaves, O and L layers to the MOAs from A and extractability in NaOH, because subsoil MOC was gener- B horizons (Fig. 12B). The facts that BAs (i) are released ally more extractable under alkaline conditions (Table 1). from source vegetation leaves, organic soil layers, and Given the large XPS-based concentration of aromatic C MOAs and (ii) become progressively enriched in MOAs rel- in A- and B-MOAs (30%) and the small amounts of ative to lignin-derived phenols suggests a high stability and/ CuO-digestible lignin monomers particularly in B-MOAs, or a strong retention of respective source compounds (likely this implies either (i) an underestimation of lignin phenols polyphenols/tannins). However, we cannot exclude that due to the CuO technique, (ii) major lignin transformations some BAs are lignin degradation products. (thus escaping the analytical window), or (iii) the presence of other, non-lignin-derived phenolic compounds. The occurrence of non-lignin-derived phenols appears realistic 3.2.6. Absolute concentrations of non-cellulosic since, besides coming from lignin, soluble polyphenols carbohydrates (92 ± 3 mg/g) and tannin (65 ± 4 mg/g) are major aromatic Table EA3 provides the MOA- and OC-based concentra- components of Metrosideros litter (Rothstein et al., 2004). tions of hydrolyzable neutral and acidic (glucuronic acid, Fig. 12A plots the S/V versus C/V ratios of lignin in galacturonic acid) monosaccharides across the LSAG. On source leaf materials, organic layers, and MOAs. The litter a mol%-basis, glucose, mannose, and galactose were the layer reflects the predominant angiosperm (M. polymorpha) dominant hexoses across the LSAG, accounting on average origin (S/V 0.7–1.1), which becomes progressively altered in for 62 ± 2% and 59 ± 2% of the total carbohydrate pool in the O layers as shown by decreasing S/V and increasing A- and B-MOAs, respectively (not shown). Plant-derived acid-to-aldehyde ratios of vanillyl (ac/al)V and syringyl pentoses (xylose and arabinose) were less abundant, com- (ac/al)S units (Fig. 10B and C; Table EA2). This trend re- prising on average only 25 mol% in both top- and subsoil sults from the preferential removal of syringyl moieties MOAs. Uronic acid concentrations were generally (Miltner and Zech, 1998; Mo¨ller et al., 2002; Otto and <5 mol% (2.7 ± 0.3 mol%; mean ± standard error; n = 12). Simpson, 2006). Compared with the organic layers, lignin As with lignin phenol monomers, the concentration of neu- phenols in A- and B-MOAs exhibit a much higher degree tral and acidic carbohydrates paralleled the MOC accumula- Evolution of mineral–organic associations upon substrate aging 2053

Fig. 11. (A–D) Time-dependent change of phenol concentrations (A = A horizon; B = B horizon) and non-cellulosic carbohydrate concentrations (C = A horizon; D = B horizon) normalized to mineral-associated OC (MOC). Abbreviations: K8, lignin phenols; BA, benzoic acid; 2H BA, 2-hydroxybenzoic acid; 3H BA, 3-hydroxybenzoic acid; 4H BA, 4-hydroxybenzoic acid; 3,5 diH BA, 3,5-dihydroxybenzoic acid; Xyl, xylose; Ara, arabinose; Rib, ribose; Rha, rhamnose; Fuc, fucose; Man, mannose; Gal, galactose; Glu, glucose; GluA, glucuronic acid; GalA, galacturonic acid.

tion, culminating during weathering stage II and declining terms of quantity) as bulk MOC. Although absolute carbo- thereafter (Fig. 10D). Thus, neutral and acidic carbohydrate hydrate concentrations in A-MOAs (mg/g MOA) increase concentrations correlated positively and significantly with towards the 400-kyr site, the respective concentrations in MOC concentrations (r2 = 0.76 and r2 = 0.78, respectively; B-MOAs decrease continuously from weathering stage I to p < 0.001). In contrast to the lignin phenols, the fractional III (Fig. 10D). This inverse trend in A- and B-MOAs sug- change of carbohydrate concentrations in A-MOAs towards gests an input control for carbohydrates into deeper soil the next older site was roughly congruent to that of MOC horizons, i.e., carbohydrates appear already effectively re- concentrations (Table 5). This reveals that carbohydrate tained in A horizons. The large absolute carbohydrate con- components respond similarly to mineralogical shifts (in centrations in A-MOAs of weathering stage II suggest that 2054 R. Mikutta et al. / Geochimica et Cosmochimica Acta 73 (2009) 2034–2060

Fig. 12. (A–C) Plot of the syringyl-to-vanillyl ratio versus the cinnamyl-to-vanillyl ratio of source vegetation leaves, organic layers, and mineral–organic associations (A), MOC-normalized concentration of lignin phenols (K8) versus the ratio of benzenecarboxylic acids (BA) to lignin phenols (B), and the time-dependent change of the lignin-to-carbohydrate ratio in organic layers and mineral–organic associations (C). Roman numbers indicate the different weathering stages (see text). Abbreviations: A-MOA and B-MOA, mineral–organic associations from A and B horizon, respectively.

non-cellulosic carbohydrates significantly contribute to sta- mature mineral assemblage (Table 1). The large K8/carbo- bilized OM in MOAs, despite being generally considered hydrate ratio, however, indicates that at the oldest site lig- an easily available C source. nin-derived phenols were again preferentially associated with minerals (Fig. 12C). This can be ascribed to the pres- 3.2.7. Concentration and source of non-cellulosic ence of minerals which still offer more variable surface carbohydrates in mineral-associated organic matter charge than those at the youngest site (Chorover et al., Fig. 11C and D show that the MOC-normalized concen- 1999), thus favoring the sorption of lignin-derived phenols trations of carbohydrates decline from the 0.3-kyr to the over that of carbohydrates. 4100-kyr site, with the 1400-kyr site being an exception. In order to assess the origin of sugar components, we The large contribution of carbohydrates in the A-MOA calculated the ratios of microbial-derived hexoses/deoxy- of the youngest site, with lignin phenols contributing least, hexoses to plant-derived pentoses in MOAs (Murayama, suggests that sugar components are more relevant in MOAs 1984; Oades, 1984) according to GM/AX = (galact- formed with less reactive primary minerals. Similarly, the ose + mannose)/(arabinose + xylose) and RF/ K8/carbohydrate ratio in A-MOAs tended to increase with AX = (rhamnose + fucose)/(arabinose + xylose) (Fig. 10E time (Fig. 12C), indicating that, relative to lignin phenols, and F). GM/AX values for plant sources are typically carbohydrates are more strongly retained at the youngest <0.5 and those of microorganisms >2 (Oades, 1984). site whereas in mineralogical settings where reactive sec- Increasing RF/AX values from 0 to 0.38 during the biodeg- ondary minerals prevail, the amount of lignin phenols is rel- radation of straw has been ascribed to the microbial synthe- atively larger. Hence, at least for weathering stage II, we sis of deoxysugars (Murayama, 1984). At each site, the infer that carbohydrates are discriminated against other GM/AX ratio increased steadily from organic layers (litter: OM components, including lignin phenols, either by 0.53 ± 0.05; O layer: 0.91 ± 0.09) to MOAs (1.69 ± 0.10; adsorption and/or precipitation processes. At the 1400- n = 12) (Fig. 10E). The high GM/AX values suggest a pre- kyr site, MOC-normalized carbohydrate and lignin-derived dominant, though not exclusive, microbial origin of hydro- phenol concentrations were largest, suggesting that the min- lyzable non-cellulosic carbohydrates in MOAs. This finding eral assemblage at Kolekole is particularly suitable in stabi- clearly highlights the importance of microbial carbohydrate lizing both, carbohydrate and lignin components (Figs. 10A resynthesis in the formation of MOAs. Although absolute and D and 11A and C). Although the K8/carbohydrate ra- carbohydrate concentrations in A-MOAs increased during tio slightly declined with respect to the 400-kyr site, the ra- weathering stage I and II, the proportion of microbial hex- tio was still larger than at the youngest site (Fig. 12C). At oses tended to decrease with time, as reflected by slightly the oldest site, carbohydrate concentrations were smallest declining GM/AX and RF/AX values (Fig. 10E and F). among all sites, both absolute and when normalized to This suggests a somewhat larger availability of OM for sus- OC (Figs. 10D and 11C). The relatively minor occurrence taining microbial activity at the younger sites with mineral– of carbohydrates (and lignin phenols) at the 4100-kyr site organic interactions being less strong, which translates into is consistent with the limited surface reactivity of the most a larger microbial necromass, hence more mineral-bound Evolution of mineral–organic associations upon substrate aging 2055 microbial hexoses/deoxyhexoses. As the mineral–organic range between 332 and 348 K. Except from the oldest sam- associations at later weathering stages become more stable, ple, all B-MOAs revealed one significant and reproducible microbial utilization of OM and production of microbial transition located between 366 and 382 K. Fig. 14 depicts sugars is likely more restricted. the step transition temperature (T*) of MOM at the differ- ent sites. Except from the 20-kyr site, the T* values in the 3.3. Time-dependent physical alteration of mineral-associated A-MOAs were comparable across all weathering stages organic matter and were always significantly smaller than in the respective B-MOAs (Fig. 14). Generally, the T* in the A-MOAs are in The binding of OM to minerals that differ in grain size line with those reported in the literature (Hurrass and and surface reactivity should influence the mobility of or- Schaumann, 2005; Schaumann and LeBoeuf, 2005) while ganic side chains to different extents, thus reducing their de- in the B horizons the T* values exceed the range of data gree of freedom to alter their three-dimensional structure published for A and O horizons. The larger T* values of upon hydration or to interact with solutes. In order to MOM in the B horizons than in A horizons indicate a lar- investigate the potential change in OM rigidity across the ger rigidity of MOM in the B horizon. However, the LSAG, differential scanning calorimetry (DSC) was applied roughly similar T* values across the different weathering re- to the A- and B-MOAs of the six sites (Fig. 13). All A- gimes suggest that the mineral composition as well as the MOAs exhibit a clear step transition in the temperature spatial arrangement of the mineral–organic fabric have no

Fig. 13. Representative DSC thermograms of mineral–organic associations isolated from A (left) and B horizons (right). The asterisk (*) indicates samples where not all thermograms revealed a significant transition. The sample marked with (**) showed a weak reversing transition, which has been eliminated by the baseline correction.

Fig. 14. Time-dependent change of the DSC step transition temperature (T*) of organic matter in mineral–organic associations from A (A- MOA) and B horizons (B-MOA). Marked samples did not show transitions in every replicate (*) or showed a weak, but reversing transition only (**). 2056 R. Mikutta et al. / Geochimica et Cosmochimica Acta 73 (2009) 2034–2060 major effect on the rigidity of MOM. The observed differ- than at the youngest and older sites (Torn et al., 1997). This ence in rigidity between A- and B-horizon MOM is proba- likely is a combined result of the following factors: (i) the bly due to a more intimate association of OM with minerals strong chemical binding of OM to the PC minerals via li- in the B horizon, i.e., a larger extent of direct attachment to gand exchange reactions at low pH, (ii) the sterically favor- the minerals, which in line with the smaller VOM/Vmineral able association of OM with mineral micropores, (iii) the ratios in B-MOAs (Table 2). A larger extent of mineral–or- effective ‘embedding’ of OM in the mineral matrix due to ganic interaction is generally favored at smaller OM frac- the saturation of OM functional groups with small mineral tional coverages (as indicated by the larger CBET colloids or monomeric/polymeric metals, thus also physi- constants of B-MOAs; Table 2), because OM can bind in cally protecting OM, and (iv) the large intrinsic stability an ‘uncoiled formation’ with a maximum number of ligands of bound aromatic compounds (Kalbitz et al., 2005; Scheel involved (Kaiser and Guggenberger, 2003). Thus, complex- et al., 2007). Factors i–iii all hamper the desorption of OM ation with minerals and metals can reduce the mobility of and thus its microbial utilization. the structural parts of OM. Although not present in all (2) Accessibility and reactivity of mineral surfaces to- samples and in all measurement repetitions, a second step wards organic matter. Despite about one-third to one-half transition was measured at T = 353–373 K in many A- of the mass of A-MOAs in stage-II soils being attributable MOAs samples showing the presence of the more rigid to OM, the mineral surfaces appear only moderately cov- mineral–organic associations in the A horizons as well. ered by OM. This suggests that even under productive The detection of this transition, however, is difficult due wet forest ecosystems that constantly deliver organic detri- to the overlap with the strong lower-temperature transi- tus and dissolved OM to the soil (Neff et al., 2000), mineral tion. We assume that this high-temperature transition surfaces are not completely covered by OM. In the youn- probably reflects the property of OM in close proximity gest and oldest soil, mineral surfaces are less favorable for to the mineral interface. Consequently, the lower-tempera- OM sorption because of their small reactivity, i.e., less ture transition prevailing in A-MOAs might be more SSA, microporosity and variable surface charge (Chorover indicative of the bulk OM structure where the mineral–or- et al., 2004). In contrast, in the intermediate-aged soils, sig- ganic binding has less impact on matrix rigidity. This idea nificant surface area as probed by N2 may be located in is supported by the fact that only in A-MOAs, T* corre- small agglomerations/aggregates that cannot be penetrated lated positively with structural properties of MOM that is by OM. It is also possible that OM units are simply too the XPS-based aromatic C:aliphatic C ratio (r2 = 0.76; large to effectively coat agglomerations of nanometer-sized, p < 0.05). The stronger rigidity of OM in B-MOAs might, high-SSA minerals. due to a larger extent of mineral–organic binding and the (3) Mineral control on the chemical composition of min- formation of multiple bonds, limit OM desorption and/or eral-associated organic matter. The attempt to reveal defi- microbial utilization and can therefore help explaining the nite mineralogical controls on the composition of MOM longer mean residence time of OM in subsoil compared to using XPS yielded ambiguous results. The MOM in A hori- topsoil horizons. zons showed little compositional variation despite of the pronounced mineralogical changes, suggesting that the 4. SUMMARY AND IMPLICATIONS chemical signature of topsoil MOM predominantly reflects the uniform source vegetation. In contrast, in subsoil (1) Composition of mineral–organic associations and MOAs containing PC aluminosilicates and Al-organic com- implications for organic matter stabilization. The application plexes, a larger abundance of acidic and aromatic OM com- of density fractionation in combination with multiple ana- ponents is consistent with the high affinity of acidic lytical techniques provided detailed insight into the time- aromatic compounds for variable-charged minerals. Of dependent evolution of MOAs across a long-term mineral- note, in both A- and B-MOAs, we found that PC Fe min- ogical soil gradient. Mineral transformations along this erals significantly explained the concentrations of aliphatic 0.3–4.100 kyr soil chronosequence caused a differential and amide C, but we have no unequivocal explanation for storage of OM in MOAs that is primarily driven by the this result. This finding suggests more complex pathways weathering-induced formation of metastable PC Fe, Al, Si in the formation of MOAs, possibly involving microbial phases and Al-organic complexes (stage II). Both, the sur- activity in connection with precipitation reactions during face C and N enrichment data as well as OM–mineral vol- redox-oscillations (Thompson et al., 2006). umetric ratios suggest that OM at the youngest and oldest Lignin and carbohydrate biomarker analyses proved site is bound to larger mineral grains on average, likely to more suitable in revealing distinct mineral–organic patterns edge faces of primary/secondary silicates, leaving much across the mineralogical soil gradient. Lignin-derived phe- mineral surface area uncovered by OM. In contrast, the nols accumulated preferentially in MOAs rich in high- MOAs at the 20–400 kyr old sites are composed of a ‘well SSA PC minerals and metal–organic complexes during mixed’ homogeneous array of OM and nanometer-sized weathering stage II (20–400 kyr), but were, in absolute mineral phases, with the OM volume partly exceeding the numbers, less important at the youngest and oldest site. mineral volume. Such volumetric ratios indicate that in Independent from the mineralogical regime, acidic lignin OM-rich surface soil horizons, precipitation in addition to phenols were selectively associated with minerals. However, adsorption reactions is involved in the formation of MOAs. given that only a minor fraction of aromatic components In these stage-II MOAs, particularly in the subsurface ones, (as shown by XPS) is attributable to lignin-derived phenols, OM is more effectively protected against biodegradation additional aromatic sources like tannins/polyphenols Evolution of mineral–organic associations upon substrate aging 2057 should contribute and/or the lignin in MOAs might already APPENDIX A. SUPPLEMENTARY DATA be highly altered and thus be invisible for the CuO tech- nique. The analysis of the carbohydrate fraction showed Supplementary data associated with this article can be that sugars are an integral part of MOAs and can become found, in the online version, at doi:10.1016/ effectively stabilized particularly in presence of PC mineral j.gca.2008.12.028. phases and hydrolyzable metal species during weathering stage II. In contrast to lignin-derived phenols, non-cellu- REFERENCES losic carbohydrates were relatively more important in the formation of MOAs at the youngest site where primary Aiken G. R. and Leenheer J. A. (1993) Isolation and chemical minerals prevail. Consequently, the fractionation of OM characterization of dissolved and colloidal organic matter. during adsorption or precipitation reactions is an important Chem. Ecol. 8, 135–151. pathway that imposes differences in the chemical composi- Amelung W., Kaiser K., Kammerer G. and Sauer G. (2002) tion of MOM. 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