contribution of dissolved organic matter to carbon storage in forest mineral soils

2008 WILEY-VCH Verlag GmbH & Co. KGaA, Weinheim www.plant-soil.com

52 DOI: 10.1002/jpln.200700043 J. Plant Nutr. Soil Sci. 2008, 171, 52–60

Review Article

Contribution of dissolved organic matter to carbon storage in forest mineralsoils§

Karsten Kalbitz1* and Klaus Kaiser2

1 Department of Soil Ecology, University of Bayreuth, 95440 Bayreuth, Germany2 Soil Biology and Soil Ecology, Martin Luther University Halle-Wittenberg, Weidenplan 14,

06108 Halle (Saale), Germany

AbstractDissolved organic matter (DOM) is often considered the most labile portion of organic matter insoil and to be negligible with respect to the accumulation of soil C. In this short review, we pre-sent recent evidence that this view is invalid. The stability of DOM from forest floor horizons,peats, and topsoils against microbial degradation increases with advanced decomposition of theparent organic matter (OM). Aromatic compounds, deriving from lignin, likely are the most stablecomponents of DOM while plant-derived carbohydrates seem easily degradable. Carbohydratesand N-rich compounds of microbial origin produced during the degradation of DOM can be rela-tively stable. Such components contribute much to DOM in the mineral subsoil. Sorption of DOMto soil minerals and (co-)precipitation with Al (and probably also with Fe), especially of the inher-ently stable aromatic moieties, result in distinct stabilization. In laboratory incubation experi-ments, the mean residence time of DOM from the Oa horizon of a Haplic Podzol increased from<30 y in solution to >90 y after sorption to a subsoil. We combined DOM fluxes and mineraliza-tion rate constants for DOM sorbed to minerals and a subsoil horizon, and (co-)precipitated withAl to estimate the potential contribution of DOM to total C in the mineral soil of a Haplic Podzol inGermany. The contribution of roots to DOM was not considered because of lack of data. TheDOM-derived soil C ranges from 20 to 55 Mg ha–1 in the mineral soil, which represents19%–50% of the total soil C. The variation of the estimate reflects the variation in mineralizationrate constants obtained for sorbed and (co-)precipitated DOM. Nevertheless, the estimates indi-cate that DOM contributes significantly to the accumulation of stable OM in soil. A more preciseestimation of DOM-derived C in soils requires mineralization rate constants for DOM sorbed toall relevant minerals or (co-)precipitated with Fe. Additionally, we need information on the contri-bution of sorption to distinct minerals as well as of (co-)precipitation with Al and Fe to DOMretention.

Key words: biodegradation / chemical composition / (co-)precipitation / dissolved organic matter / DOM /forest soils / sorption

Accepted May 31, 2007

1 Introduction

Growing concern about climate change evoked interest in therole of dissolved organic matter (DOM) in the global C bal-ance (Hedges, 2002). Fluxes of dissolved organic C (DOC)into the marine environment via runoff from terrestrial settingsare estimated to 0.2 (Harrison et al., 2005) to 0.4 × 1015 g(IPCC, 2001). These fluxes are small compared to the Cfluxes due to primary productivity and heterotrophic respira-tion in terrestrial ecosystems. Total DOC in the oceans isapprox. 685 × 1015 g (Hansell and Carlson, 1998) and contri-butes only to 1.8% of marine C storage. Therefore, DOC

fluxes are generally not considered to be of relevance for theglobal C cycle (Neff and Asner, 2001). However, DOC fluxesin terrestrial ecosystems are several folds larger than thosewith rivers. In certain environments—particularly in peatlandsand at northern latitudes—soluble C can substantially contri-bute to the ecosystem C cycle (Neff and Asner, 2001). Largeinternal DOC fluxes in soils relative to small outputs intoaquatic ecosystems may result from mineralization or stabili-zation, thus, accumulation in the soil, as suggested by Kaiserand Guggenberger (2000).

* Correspondence: Dr. K. Kalbitz;e-mail: [email protected]§ Topical Issue Soils as a source and sink for CO2 – Mechanismsand regulation of organic matter stabilisation in soils (editors: I. Kögel-Knabner and E. Matzner). Synthesis of the DFG Priority ProgramSPP 1090 (German Research Foundation—“Deutsche Forschungs-gemeinschaft”).

* Erscheinungsjahr und Bandnummer waren falsch

Litter layers and the upper, organic-rich mineral horizons arethe main sources of DOM in soils whereas the deeper mineralhorizons are the major sinks. In forest ecosystems, fluxes ofDOC from the forest floor into the mineral soil have been esti-mated to 115–500 kg C ha–1 y–1, representing up to 35% ofthe annual litterfall C (Guggenberger and Zech, 1993; Currieet al., 1996; Michalzik and Matzner, 1999; Solinger et al.,2001; Kaiser et al., 2001; Michalzik et al., 2001). The fluxesfrom forest floor and topsoil horizons down into deeper soilhorizons decrease because of degradation in soil solution,sorption to Fe and Al oxides/hydroxides and clay minerals,and (co-)precipitation by polyvalent cations (summarized byKalbitz et al., 2000). Therefore, concentrations of DOC indeeper soil horizons as well as the export from mineral sub-soil are usually small (Guggenberger and Zech, 1993; Michal-zik and Matzner, 1999; Solinger et al., 2001). Retention inmineral subsoils ranges from 40 to 370 kg DOC ha–1 y–1 (Cur-rie et al., 1996; Guggenberger and Kaiser, 2003). Given thatmost of the reduction in DOC is due to abiotic reactions andnot mineralization, DOC could be a major contributor to thetotal C accumulation in soil, including O layers, which mayreach up to 1440 kg C ha–1 y–1 (Schulze et al., 2000). Basedon estimates of Michalzik et al. (2001) and Guggenbergerand Kaiser (2003), the retention of DOC in the mineral soil of17 temperate forest sites studied in Europe and N Americaaverages to 191 kg DOC ha–1 y–1. The proportion to whichthe decrease of DOC fluxes with soil depth is either due tomicrobial consumption or abiotic storage via sorption or preci-pitation, however, is mostly unknown. Besides knowledgeabout DOM fluxes, we need robust estimates of DOM degra-dation in soil solution and after sorption and (co-)precipitationto rate the contribution of DOM to the formation of stable soilorganic matter (SOM).

In this short review, we summarize recent findings on the sta-bilization of DOM in soils, defined as reduced respiratory Crelease. We evaluate the stability of DOM against microbialdegradation in soil solution, after sorption to mineral subsoiland minerals and after precipitation by polyvalent cations.Furthermore, we estimate the portion of DOM-derived OM inthe mineral soil of a Podzol using DOM fluxes measured inthe field and degradation rates of DOM from that site.

2 Degradation of DOM in soil solution

Laboratory incubations showed 5%–93% of the DOM in soilsolutions to be potentially microbially degradable (Jandl andSletten, 1999; Yano et al., 2000; Kalbitz et al., 2000; Sachseet al., 2001; Kalbitz et al., 2003a; Don and Kalbitz, 2005;Kiikkilä et al., 2006). Marschner and Kalbitz (2003) consid-ered intrinsic properties of DOM to control its biodegradation.Dissolved organic matter with a large portion of C in theXAD-8-adsorbable fraction (i.e., hydrophobic by definition),rich in aromatic structures, and poor in carbohydrates is littlebiodegradable (Qualls and Haines, 1992; Jandl and Sollins,1997; Jandl and Sletten, 1999; Volk et al., 1997; Amon et al.,2001; Kalbitz et al., 2003a; Kiikkilä et al., 2006). Kalbitz et al.(2003a) concluded that microbial degradation of DOM de-rived from forest floors, peats, and A horizons decreased withincreasing degree of decomposition of the parent material.That seems reasonable because of the preferential utilization

of labile material at earlier states of decomposition(Haider, 1992; Dai et al., 2001). In consequence, DOMpercolating into the mineral soil should be largely stablebecause it derives predominately from the more decomposedparts of the forest floor (Fröberg et al., 2006). That alsosuggests that microbial degradation cannot explain thedecrease in DOC concentrations in the mineral soil withincreasing depth.

Aromatic compounds, possibly deriving from lignin, whichsurvive biodegradation preferentially, seem to be the moststable constituents of DOM from forest floors and topsoils(Kalbitz et al., 2003b). Thus, the partial biodegradation leavesbehind compounds with favorable properties for sorption tominerals (McKnight et al., 1992; Kaiser et al., 1997) and con-sequently for further stabilization (Kalbitz et al., 2005; seebelow).

In acid forest soils, Al and Fe can form stable complexes withDOM. Schwesig et al. (2003a) found a decreased mineraliza-tion after addition of Al to four different DOM solutions. How-ever, DOM complexation by metals will not necessarilyreduce its biodegradability but instead may enhance micro-bial activity by reducing the free metal-ion concentrations andthus their toxicity (Marschner and Kalbitz, 2003).

The properties of DOM change with soil depth due to prefer-ential sorption and (co-)precipitation of aromatic, lignin-de-rived portions, resulting in increasing portions of carbohy-drate- and N-rich compounds in solution (McKnight et al.,1992; Kaiser et al., 1997; Kalbitz, 2001; Schwesig et al.,2003b; Kaiser et al., 2004). Despite its large content of carbo-hydrates, subsoil DOM is less degradable than the largelystable DOM rich in aromatic compounds from Oa horizons(Boyer and Groffman, 1996; Schwesig et al., 2003b). Where-as plant-derived carbohydrates seem to be the preferred sub-strate for microorganisms (Kalbitz et al., 2003b), microbialcarbohydrates and N-rich compounds released during OMdegradation can be hardly degradable (Ogawa et al., 2001;Kalbitz et al., 2003b). Consequently, subsoil DOM, enrichedin carbohydrates, amino sugars, and other compounds ofmainly microbial origin (Guggenberger and Zech, 1994;Kaiser et al., 2004), is not labile per se as assumed previous-ly (e.g., Qualls and Haines, 1992).

3 Degradation of DOM after sorption tominerals

Sorption to mineral subsoils, clay minerals, and Fe oxide/hydroxides results in strong stabilization of DOM, as indi-cated by smaller mineralization rate constants for sorbed OMthan for DOM (Tab. 1; Kalbitz et al., 2005; Mikutta et al.,2007). The mean residence time of OM derived from the Oahorizon of a Podzol increased from <30 y in solution to 91 yafter sorption (Kalbitz et al., 2005). These results are inagreement with stabilization, sometimes even complete sup-pression of mineralization, of simple organic compounds likeglucose, citrate, oxalate, and malate after sorption to miner-als and mineral soils (Jones and Edwards, 1998; Ström et al.,2001; van Hees et al., 2003). In turn, Keil et al. (1994) and


J. Plant Nutr. Soil Sci. 2008, 171, 52–60 Stabilization and dissolved organic matter 53

Nelson et al. (1994) found mineralization rates for OM toincrease up to five orders of magnitude after desorption frommarine sediments and soils, respectively. Fluorescence spec-tra, UV absorbance, and 13C signature suggest that aromaticcompounds were preferentially stabilized by sorption (Kalbitzet al., 2005).

The two main reasons responsible for the stabilization ofDOM by sorption are: (1) increased stability of organic com-pounds due to interactions with the mineral phase (Kögel-Knabner et al., 2008a, this issue, pp. 61–82) and (2) selectivesorption of inherently more stable components. Formation ofstrong chemical bonds between the mineral phase andorganic molecules likely reduces the mineralization of C aftersorption (Saggar et al., 1994; Kaiser and Guggenberger,2000, 2007). Furthermore, physical separation of OM frommicroorganisms and hydrolytic enzymes by sorption intosmall pores (Mayer, 1994) and sterical changes due to multi-ple ligand attachment to the sorbing surface (Kaiser andGuggenberger, 2003) can contribute to the reduced C miner-alization.

Sorption is strongly selective for the inherently recalcitrantaromatic and complex compounds (Kaiser and Guggenber-ger, 2000; Kalbitz et al., 2005). Sorptive stabilization of stablecompounds is consequently much more efficient than of theless strongly sorbing labile ones (Ransom et al., 1998; Kaiserand Guggenberger, 2000; Kalbitz et al., 2005). Therefore, thestronger the sorption of a given DOM, the stronger is thesorptive accumulation and stabilization of OM (Kaiser andGuggenberger, 2000; Kalbitz et al., 2005).

4 Degradation of DOM after (co-)precipitation

Precipitation with Al and Fe can be efficient in immobilizingDOM in acidic soils (Zysset and Berggren, 2001; Nierop etal., 2002; Jansen et al., 2005). At molar Al : C ratios typical

for acidic forest soils (0.1), up to 80% of DOC can (co-)preci-pitate (Nierop et al., 2002; Scheel et al., 2007). Co-precipita-tion (also termed co-flocculation) means precipitation of themineral phase (e.g., Al hydroxide) with simultaneous adsorp-tion of organic molecules (e.g., Jekel, 1986). It is difficult todistinguish between true precipitation, i.e., formation of inso-luble organic–metal complexes, and co-precipitation. Bothprocesses may occur simultaneously. Therefore, we use theterm (co-)precipitation. As for the adsorption of DOM tominerals, spectroscopic analyses indicated preferential (co-)precipitation of aromatic compounds, suggesting that mainlyinherently stable compounds tend to precipitate (Römkensand Dolfing, 1998; Blaser et al., 1999; Dolfing et al., 1999;Sharpless and McGown, 1999; Scheel et al., 2007). Schwe-sig et al. (2003a) noticed formation of flocculated OM andsimultaneously decreasing mineralization of DOM after addi-tion of Al to soil solutions.

Mineralization of DOM (co-)precipitated with Al is small ascompared to DOM in solution (Tab. 1; Boudot et al., 1989;Scheel et al., 2007). It depends on the Al species involved in(co-)precipitate formation (Scheel et al., 2007). The propor-tion of hydrous Al species seems to be crucial, and it in-creases with increasing pH and Al : C ratios. Stabilizationdecreased with increasing proportions of hydrous Al speciesinvolved in (co-)precipitation as indicated by larger minerali-zation of precipitates formed at pH 4.5 in comparison to 3.8and with increasing Al : C ratios in the precipitates (Scheelet al., 2007). Probably, OM interacts stronger with Aln+ thanwith Al hydroxides.

The observed stability of (co-)precipitates could contributesignificantly to stabilization of DOM in soils. The Al concentra-tions in the solution of acidic soils increase with depth(Matzner et al., 2004). Besides adsorption on Al and Fe oxi-des, (co-)precipitation with Al (and other polyvalent cations)could explain the decreasing DOC concentrations with soil


Table 1: Results of the three incubation experiments used in this study. Parameters were calculated on a basis of 50 d incubation at 20°C. Forall incubation experiments, the same DOM (from the Oa horizon of a Podzol; Kalbitz et al., 2005) was used. A and k were calculated by a singleexponential model: Cmineralized (%) = A (1 – exp[–kt]); k: mineralization rate constant (d–1); t: time (d); A: mineralizable DOC (%).

Study / treatment Mineralized C(%)

A(%)

k(d–1)

Conversion factor a

Results of Kalbitz et al. (2005) b

OM in solution 18.6 17.4 0.1033OM sorbed to mineral soil 1.4 1.5 0.0482 1.000

Results of Mikutta et al. (2007)OM in solution 8.7 8.5 0.1675OM sorbed to goethite c 2.1 2.4 0.0397 1.214OM sorbed to pyrophyllite c 3.5 4.0 0.0571 0.844

Results of Scheel et al. (2007)OM in solution 4.8 4.1 0.2466OM (co-)precipitated with Al d 0.5 0.7 0.0212 2.274

a ratio between the mineralization rate constant of OM sorbed to the mineral soil and the constant of OM sorbed to minerals or (co-)precipitatedwith Alb results of the long-term incubation of Kalbitz et al. (2005); data after 50 d incubationc sorption with NaCl as background electrolyte (Mikutta et al., 2007)d precipitation at pH 3.8; molar Al : C ratio: 0.1 (Scheel et al., 2007)

54 Kalbitz, Kaiser J. Plant Nutr. Soil Sci. 2008, 171, 52–60

depth as well as accumulation and stabilization of OM in sub-soils. Co-precipitates rich in OM can also form with Fe(Schwertmann et al., 2005), especially if weathering of Fe(II)-bearing minerals is rapid or if interchanging redox conditionsprevail. However, so far no studies addressed the stability ofOM precipitated with cations such as Fe3+ and Ca2+.

5 Estimation of DOM contribution to C in thesubsoil of a forested site

The contribution of DOM to the accumulation of OM in themineral soil depends on the amount of C retained and itsmineralization. The retained C is the difference between theinput of DOC into and its export from the mineral soil. Thedecreases in DOM in the mineral soil should be caused byabiotic processes such as sorption and (co-)precipitationbecause microbial degradation of OM in solution likely is ofminor importance (see above).

We estimated the contribution of DOM to the organic C (OC)in the mineral horizons of a Podzol under Norway spruce inthe “Fichtelgebirge”, Germany (Kögel-Knabner et al., 2008b,this issue, pp. 5–13; Tab. 2). To the best of our knowledge, itis the only site where DOM retention and DOM mineralizationin soil solution, after sorption to the mineral soil and differentminerals as well as after precipitation by Al, has been studied.The observed retention of 260 kg C ha–1 y–1 between 0 and90 cm depth (Kalbitz et al., 2004a) is well within the range ofthe 17 sites (40–370 kg C ha–1 y–1) mentioned in the introduc-tion. According to Fröberg et al. (2006), much of the retentionof DOC occurs in the B horizons. All mineralization rate con-stants used (Tab. 1, 3) were obtained by incubation of DOMfrom the Oa horizon of the study site. Also, the mineral soilhorizon (Bw) used to test the effect of DOM sorption (Kalbitzet al., 2005) is from that site. Therefore, we have a unique setof data that allows for estimating the contribution of DOM toOM in the mineral soil.

5.1 Methods

Generally, the pool size of OM (A) with constant input flux(F; e.g., DOM) and constant exponential mineralization rateconstant (k) is determined as

A(t) = F (1 – exp[–kt ])/k (Olson, 1963).

The maximum value of OM accumulation from the input flux(Aend) can be calculated by assuming time to approachinfinity:

Aend = F / k.

As described earlier, mineralization of DOM decreases afteradsorption to minerals or after (co-)precipitation with poly-valent cations. Therefore, mineralization of the retained DOMcan be described by different mineralization rate constants.The pool size of OM derived from the constant DOM input (A)can be calculated as follows:

A(t) = F1 / k1 (1 – exp[–k1t ] + F2 / k2 (1 – exp[–k2t ]) + ......+ Fn / kn (1 – exp[–knt ]).

F1 to Fn represent the amounts of DOM retained by sorptionto different soil minerals and (co-)precipitation; k1 to kn arethe respective mineralization rate constants. Considering thedifferent mineralization rate constants, the limit value of OMaccumulation from DOM input (Aend) can be calculated:

Aend = F1 / k1 + F2 / k2 + ...... + Fn / kn.

To estimate the long-term contribution of DOM to the OMpool in the mineral soil, we used the mineralization rate con-stants of the stable C pool (kslow) for OM sorbed to themineral soil (Bw horizon), as obtained by long-term (375 d)incubation (Tab. 3; Kalbitz et al., 2005). The mineralizationrate constants for OM after sorption to soil minerals (Mikuttaet al., 2007) and after (co-)precipitation by Al (Scheel et al.,2007) could not be used directly because of shorter incuba-tion periods (precipitated OM: 50 d; sorbed OM: 90 d).Furthermore, mineralization rate constants computed with adouble-exponential model were partly not significant (p >0.1), and the standard errors were often larger than the con-stants themselves. Also Sleutel et al. (2005) observed thatdouble-exponential models did not fit data obtained in degra-dation experiments with incubation periods <100 d. There-fore, we normalized all mineralization rate constants to thoseof obtained for the long-term incubation of OM sorbed tomineral soil. To do so, we calculated a mineralization rateconstant for the first 50 d of the long-term incubation of OM


Table 2: Soil properties, C storage (Kalbitz et al., 2004b), 14C contents, and radiocarbon age (Rumpel et al., 2002) at the Waldstein site (“Fich-telgebirge”, Germany; Kögel-Knabner et al., 2008b, this issue, pp. xx–xx).

Horizon Thickness(cm)

Sand Silt(%)

Clay pH(CaCl2)

C stock(t C ha–1)

14C(pmC)

Radiocarbon age(y BP)

Oi 0.5 3.60 3.6Oe 5 2.90 25.4Oa 3 2.60 31.0EA 10 51.6 38.0 10.4 2.90 27.4 93.6 525 ± 30Bh 2 34.0 49.6 16.4 3.30 8.5 98.5 120 ± 25Bs 18 44.7 44.8 10.4 3.90 51.1 91.1 745 ± 40Bw 25 45.8 43.4 10.8 4.30 20.7 82.2 1570 ± 25C1 15 56.4 34.0 9.6 4.20 1.5 62.0 3840 ± 70C2 15+ 50.8 38.0 11.2 4.10 0.9R 170.1


sorbed to the mineral soil (Tab. 1). That mineralization rateconstant was divided by those of sorbed and precipitated OMobtained in short-term incubations (50 d). The calculatedratios (conversion factors, Tab. 1) were then divided by themineralization rate constant (kslow) determined for OM sorbedto the mineral soil by long-term incubation (Tab. 3). Weassume the difference in rate constants between the soilminerals and precipitates in the short-term incubation to holdthrough also for long-term incubation.

From the study of Mikutta et al. (2007), we selected two soilminerals, goethite and pyrophyllite (an illite-like material butwith smaller permanent charge), with NaCl as the back-ground electrolyte for sorption. From Scheel et al. (2007), weselected the precipitates obtained at pH of 3.8 and a molarAl : C ratio of 0.1; because these values reflect average con-ditions in the different mineral soil horizons (Tab. 2).

Our data set did not allow for distinguishing between differentmineral soil horizons. Instead, we assumed a uniform reten-tion throughout the mineral soil and neglected preferentialflow (Kalbitz et al., 2005). Furthermore, all mineralization rateconstants were adjusted to the mean annual temperature ofour study site (5°C) using the approach of Tjoelker et al.(2001) and Kalbitz et al. (2005):

kx = kx’ / Q10(Tx’ – Tx) / 10°C,

Tx = 5°C (specific temperature: aim of conversion),

Tx’ = 20°C (reference temperature: incubation),

kx = mineralization rate constant at 5°C,

kx’ = mineralization rate constant at 20°C.

As we do not know to which extent sorption by different soilminerals and (co-)precipitation with Al contributed to theDOM retention, we tested each variant independently by allo-cating the total retention of 260 kg C ha–1 y–1 (Kalbitz et al.,2004a) to the various processes. This procedure gives therange of the contribution of DOM to OC in the mineral soil.The four scenarios tested were:

(1) DOM retention only by sorption to a mineral soil horizon(Bw; Kalbitz et al., 2005);

(2) DOM retention only by sorption to goethite (Mikutta et al.,2007);

(3) DOM retention only by sorption to pyrophyllite (Mikuttaet al., 2007);

(4) DOM retention only by (co-)precipitation with Al at pH of3.8 and a molar Al : C ratio of 0.1 (Scheel et al., 2007).

The study soil is characterized by large stocks of hydrous Feand Al phases and large concentrations of dissolved Al, thussorption as well as (co-)precipitation of DOM are both likely.

5.2 Results

Using the mineralization rate constants for DOM sorbed tothe Bw horizon, the soil OC deriving from DOM was esti-mated to 24 Mg C ha–1, which represents 22% of the total soilOC (Fig. 1; Kalbitz et al., 2005). The maximum contribution ofDOM-derived C was estimated to 50% (55 Mg C ha–1),assuming that all retained DOM is (co-)precipitated with Al.This is because of the 2.3-fold larger mean residence time ofOM (co-)precipitated with Al than sorbed to the Bw horizon(Tab. 1). Sorption of all retained DOM to goethite would resultin an accumulation of 28 Mg C ha–1 in the mineral soil, whichcorresponds to 26% of the stored OC. The larger mineraliza-tion rate constant for OM sorbed to pyrophyllite results in asmaller contribution to soil C (20 Mg ha–1, corresponding to19% of soil OC). Variations of the amounts of retained DOMand of the mineralization rate constants will result in propor-tional changes in the final contribution of DOM to SOMbecause of the linearity of the used equation when approach-ing the point of infinity.


Table 3: Input parameters (kslow) to estimate the contribution of DOMto OM in the mineral soil at the Waldstein site (“Fichtelgebirge”,Germany) after long-term incubation (375 d; Kalbitz et al., 2005). Thedecomposition rate constants kslow were converted to 5°C accordingto Tjoelker et al. (2001) and Kalbitz et al. (2005); kslow values derivedfrom a double-exponential model: mineralized C (% of total C) = a (1– exp[–kfast t]) + (100 – a) (1 – exp[–kslow t]); a: labile C (%), 100 – a:stable C (%), kfast: mineralization rate constant of labile C (d–1), kslow:mineralization rate constant of stable C (d–1), t: time (d). The amountof retained DOC is 260 kg ha–1 y–1 (Kalbitz et al., 2004a).

kslow of OM sorbed to mineral soil 0.000030kslow of OM matter sorbed to goethite 0.000025a

kslow of OM sorbed to pyrophyllite 0.000035a

kslow of OM (co-)precipitated withAl at pH 3.8 and Al : C = 0.1 0.000013a

a calculated by dividing the kslow of OM sorbed to mineral soil(determined by long-term incubation) by the conversion factor(Tab. 1)

0 200 400 600 800 1000

time (years)

0

10000

20000

30000

40000

50000

60000

110000

120000

C s

tock

in th

e m

iner

al s

oil (

kg C

ha-1

)

sorption to Goethite

sorption to the mineral soil

(co)precipitation by Al

sorption to pyrophilite

measured C stock in the mineral soil

Figure 1: Contribution of DOM to the build-up of OM in the mineralsoil at the site Waldstein (“Fichtelgebirge”, Germany) assuming aconstant DOM retention of 260 kg C ha–1 y–1 and constantmineralization rate constants. Dissolved organic matter retentionwas allocated to four different scenarios. Mineralization rateconstants for sorption and (co-)precipitation are given in Tab. 3.


5.3 Discussion

Our results highlight the contribution of DOM to the accumu-lation of OM in the mineral soil. The similar chemical compo-sition of DOM in forest floor percolates and of OM in acidicsubsoil is another indication that OM in the mineral soil canderive from retained DOM (e.g., Kaiser and Guggenberger,2000). In many acidic subsoils, most of the OM dissolvesupon dissolution of the mineral soil matrix with hydrofluoricacid (Eusterhues et al., 2003). That means the OM is poten-tially soluble, likely because representing sorbed or precipi-tated DOM. Decomposition of roots is likely also linked toproduction of DOM, which can be subsequently sorbed orprecipitated, thus contributing to OM in the subsoil.

Although our data do not allow to allocate the estimated stabi-lization of DOM to the soil horizons of the study site, we canassume that most of DOM-derived OM is located in B and Chorizons. The estimated DOM-derived C stock (20–55 Mg Cha–1) comprises 25%–66% of the OC in B and C horizons ofthe study soil. Organic matter in the EA horizon is mainly par-ticular, of high age and with large contribution of insoluble,long-chain alkyl compounds (Rumpel et al., 2004). Mineral-associated OM derived from sorption or (co-)precipitation isonly of minor importance in this horizon. Distribution, compo-sition, and age of C in particle-size fractions and the largeproportion of mineral-associated OC indicate a predominantinfluence of DOM on SOM in B and C horizons of the studysoil and also of a Dystric Cambisol under European beech(Rumpel et al., 2004). The large abundance of carboxyl C inthe Bs horizon is a further proof that the OM derives fromwater-soluble OM from the forest floor (Rumpel et al., 2004).The apparent mean radiocarbon age in the Bh horizon (120 y;Tab. 2) is between the mean residence time of sorbed and(co-)precipitated OM as determined by incubation studies(Fig. 1; Kalbitz et al., 2005). This is another indication for OMin the Bh horizon to derive largely from DOM.

The different scenarios tested suggest that DOM-derived C inthe mineral soil can vary over a wide range (20–55 Mg Cha–1), depending on the mineralization rate constants. Theultimate contribution of DOM to OC in the mineral soildepends on the portions of OM sorbed to different mineralsand (co-)precipitated. These portions, however, we do notknow. Adsorption to hydrous oxides (and acidic subsoils) aswell as (co-)precipitation with Al and Fe can remove DOMfrom Oa horizons almost entirely from solution (Kaiser et al.,1996; Kaiser and Zech, 1997; Nierop et al., 2002; Scheelet al., 2007). In the review of Kalbitz et al. (2000), sorption inthe mineral soil is considered to be the main process of DOMretention. On the basis of measured DOC fluxes, Guggenber-ger and Kaiser (2003) estimated the sorption capacity to beexhausted within a few years, given no neoformation of reac-tive mineral phases. However, field data did not indicate adecrease in DOC retention. In contrast, DOC concentrationsdecreased at 20 and 90 cm depth at the study site during thelast 10 y despite constant input from the forest floor (Kalbitzet al., 2004a). On the other hand, it needs approx. 300 y toreach the equilibrium between sorption and mineralization ofDOM (Fig. 1). Mineral weathering and neoformation of Feand Al oxides/hydroxides may provide sufficiently large sorp-

tion capacity over such a period of time. Thus, the actualsorption capacity cannot be used to predict the long-termpotential of soils to stabilize DOM via sorption. Also, weather-ing releases metals such as Al and Fe which can induce (co-)precipitation of DOM, thus, maintaining steady retention aswell as stabilization of DOM (Zysset and Berggren, 2001;Kalbitz et al., 2005; Scheel et al., 2007).

The apparent mean 14C age of DOM in the B horizon of twopodzolic soils in Sweden was 50 and 300 y, respectively, andexceeded the age of DOM in the forest floor by far (Fröberget al., 2006). Also NaCl-extractable OC was much older inthe B horizon of the study soil (14C content: 99.9 pmC) thanin the Oe (114.4 pmC) and in the Oa horizon (102.6 pmC;Michalzik et al., 2003). The age of DOM might reflect the timeto reach equilibrium between retention and mineralization ofDOM which is approx. 300–400 y, depending on the minerali-zation rate constant of the DOM retained. A younger age ofDOM than the equilibrium time may indicate a portion ofDOM having the 14C bomb signal to be directly leached fromthe forest floor into deeper mineral horizons.

The mineralization rate constants of (co-)precipitated OMwere lower than those of sorbed OM (Scheel et al., 2007;Mikutta et al., 2007; Tab. 1). Therefore, a stronger stabiliza-tion of OM by (co-)precipitation than by sorption seems prob-able. However, also sorption of DOM might result in much lar-ger stabilization of OM than measured by Kalbitz et al. (2005)and Mikutta et al. (2007). So, DOM sorption to a multi-domaingoethite and to amorphous Al hydroxide resulted in an almostcomplete inhibition of C mineralization (Kaiser and Schneider,unpublished data). The same observation was made byJones and Edwards (1998) after adsorption of simple organiccompounds to certain minerals such as ferrihydrite. Also, wehave no data on the mineralization of DOM co-precipitatedwith Fe. Considering the strong stabilization of organic com-pounds by ferrihydrite, the dominant mineral phase in suchco-precipitates (Schwertmann et al., 2005), the protection ofthe contained OM is likely rather efficient. These examplesimply that our estimate of the potential contribution of DOM tosoil C is largely conservative.

5.4 Limitations and uncertainties of the estimation

One limitation of our study is due to the fact that we do not haveestimates of the contribution of each of the individual com-pounds (different minerals, [co-]precipitates) to total DOM reten-tion in the mineral soil. OM sorbed by different minerals or con-tained in (co-)precipitates is mineralized at different rates (seeabove). Our estimate of DOM-derived C stocks in the mineralsoil is only 1/3 to 1/2 the estimate by Michalzik et al. (2003) whomodeled the DOM dynamics in forested ecosystems. Theyfound that 73%–89% of the OM in the mineral soil of two sitesderived from DOM, the rest from roots. In contrast, DOC con-tributed to 25% of total soil C according to a modeling studyby Neff and Asner (2001). This range of estimates points toother DOM sources like roots, not included in our estimation,and to the conservative character of our calculations.

Roots may be an important source of DOM in soil, either byreleasing exudates or due to water-soluble compounds pro-



duced during their decomposition (Kalbitz et al., 2000). Theportion of root-derived DOM, however, is largely unknown. Itis not included or at least underestimated in soil solutionsampled by lysimeters in the field because of their small sizerelative to the large spatial variability of root exudation anddecomposition. It is even impossible to estimate the contribu-tion of roots to SOM although root litter is generally lessdecomposable than leaf litter (reviewed by Rasse et al.,2005). Rumpel et al. (2004) found large amounts of root-derived compounds in SOM of the mineral soil. These find-ings suggest a larger contribution of roots to SOM than pre-viously assumed. Thus, root-derived compounds may alsocontribute significantly to DOM in the subsoil.

Mineralization of DOM results in microbial production ofwater-soluble compounds (Kalbitz et al., 2003b). This pro-cess cannot be quantified but might be important particularlyin the deeper mineral soil where other C sources are sparse.The 14C content in subsoil horizons decreases with depth(Rumpel et al., 2002); this is paralleled by a large contributionof smaller molecules of microbial origin to SOM in the deepermineral soil (Rumpel et al., 2004). These findings support theassumption of microbial DOM production directly in the sub-soil. Sorbed or (co-)precipitated OM is slowly mineralized(k values: Tab. 1, 3) which likely is accompanied by therelease of water-soluble microbial carbohydrates and pep-tides (Kalbitz et al., 2003b). These compounds will be lea-ched to deeper soil horizons.

Besides of quantitatively unknown other DOM sources in themineral soil, our estimation of the contribution of DOM to OCstorage might be conservative because we used mineraliza-tion rate constants obtained in the laboratory under optimumconditions. We assume that mineralization in the laboratory isfaster than in the field although the rate constants wereadapted to the mean annual temperature of the study site.Furthermore, mineralization rate constants depend on thetime of incubation. The rate constants for the mineralizationof C in DOM solutions from the Oi and Oa horizon of thestudy site obtained in short-term incubations (50 d) were 2.6-to 4.6-fold larger than those derived from long-term incuba-tion (375 d) (data from Kalbitz et al., 2005). From a long-termperspective, we have to consider that the mineralization rateconstants used in our calculations are probably too large.The mean residence time of SOM in the Bh horizon accordingto 14C analysis (120 y) is somewhat longer than that ofadsorbed OM obtained by long-term incubation (91 y; Kalbitzet al., 2005). The “true” mean 14C age might be even highertaking possible inputs of bomb 14C into account. If weassume the true mineralization rates in the field to be half ofthose obtained in the laboratory, the DOM-derived soil Cwould double. Then, the proportion of DOM-derived C stockin the mineral soil would be within the range of proportionsestimated by modeling (Michalzik et al., 2003).

6 Conclusions

Dissolved OM is an important contributor to the OC in mineralsoils of forest sites. Adsorption onto minerals and (co-)preci-pitation with Al results in stabilization against microbial decay.At least 20 to 55 Mg C ha–1 in the mineral soil of the test site

derives from DOM, which represents 19%–50% of total soilOC. These numbers are probably underestimates because ofthe unknown contribution of roots to the DOM pool and addi-tional DOM production in the mineral soil. (Co-)precipitationwith Al seems to reduce the mineralization stronger thansorption to phyllosilicates and goethite. However, reactiveminerals with large surface areas (multi-domain goethite, fer-rihydrite, amorphous Al hydroxide) can inhibit C mineraliza-tion of sorbed OM almost completely. Therefore, the type ofminerals and the formation of (co-)precipitates determine theDOM-derived reservoir of OC in the mineral soil. An improvedevaluation of DOM-derived soil C therefore requires robustestimates on contribution of individual minerals and (co-)pre-cipitation with Al and Fe to the total DOM retention. Specialemphasis has to be given to obtain mineralization rate con-stants under field conditions.

Acknowledgments

We gratefully acknowledge the financial support by theDeutsche Forschungsgemeinschaft Priority Program 1090“Soils as sinks and source of CO2 – mechanisms and regula-tion of organic matter stabilisation in soils”. C. Dörfler,L. Haumaier, R. Mikutta, T. Scheel, J. Schmerwitz, M. Schnei-der, and D. Schwesig contributed much to the scientific back-ground of the manuscript. For discussions and support wethank G. Guggenberger, B. Marschner, and E. Matzner.

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contribution of dissolved organic matter to carbon storage in forest mineral soils

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