dissolved organic matter in soil: challenging the paradigm of sorptive preservation

Dissolved organic matter in soil: challenging the

paradigm of sorptive preservation

Georg Guggenberger*, Klaus Kaiser

Lehrstuhl fur Bodenkunde und Bodengeographie, Universitat Bayreuth, 95440 Bayreuth, Germany

Received 7 January 2002; accepted 9 December 2002

Abstract

During the last decade, research in sedimentary systems led to the paradigm of sorptive

stabilization of organic matter (OM). Studies on soils also show that sorptive interactions between

dissolved organic matter (DOM) and mineral phases contribute to the preservation of soil OM. In the

first part of the paper, we summarize evidence for sorptive stabilization of OM in forest soils

including (a) pronounced retention of DOM in most subsoils, (b) strong chemisorptive binding

exhibiting strong hysteresis, and (c) similarity in the composition of DOM and OM in illuvial soil

horizons and clay-sized separates. However, the capacity of soils for sorption of DOM is not infinite.

In the second part of the paper, we present a case study where we relate the yearly retention of

dissolved organic carbon (DOC) in the mineral soil to the available sorption capacity of seven forest

soils. We estimate that the saturation of the sorption complex would occur within 4–30 years.

Assuming these soils are in steady-state equilibrium with respect to carbon cycling, this suggests a

mean residence time of the sorbed organic carbon (OC) of about the same time, therefore providing

little evidence for a long-term stabilization of sorbed OM.

One explanation for this discrepancy may be because in forest soils most surfaces are not

characterized by juvenile minerals but are covered with OM and colonized by microorganisms. This

is the case mainly in topsoil horizons but occurs also along preferential flow paths and on aggregate

surfaces. Biofilms develop particularly at sites receiving high input of nutrients and organic

substrates, i.e., DOM, such as preferential flow paths. The OM input enhances the heterotrophic

activity in the biofilm, converting the DOM into either organic compounds by microbial resynthesis

or inorganic mineralization products. Recent studies suggest that Fe hydrous oxides embedded

within the biofilms may serve as a sorbent and shuttle for dissolved organic compounds from the

surrounding aqueous media. We assume that sorption of DOM to the biofilm does not lead to a

stabilization of OM but is a prerequisite for its rapid turnover. Only when DOM is transported by

mass flow or diffusion to fresh, juvenile mineral surfaces, may sorption effectively stabilize OM.

0016-7061/02/$ - see front matter D 2002 Elsevier Science B.V. All rights reserved.

doi:10.1016/S0016-7061(02)00366-X

* Corresponding author. Present address: Institut fur Bodenkunde und Pflanzenernahrung, Martin-Luther-

Universitat Halle-Wittenberg, 06099 Halle, Germany. Tel.: +49-345-5522535; fax: +49-345-5527116.

E-mail address: [email protected] (G. Guggenberger).

www.elsevier.com/locate/geoderma

Geoderma 113 (2003) 293–310

This stabilization would involve complexation of functional groups, changed conformation, and

incalation in small pores.

D 2002 Elsevier Science B.V. All rights reserved.

Keywords: Organic matter; Organic carbon; Stabilization; Biofilms; Dissolved organic matter; Sorption

1. Introduction

Three types of pathways are commonly considered in the formation of stable organic

matter (OM) in soils (Christensen, 1996; Sollins et al., 1996). Selective enrichment of

organic compounds refers to the inherent recalcitrance of specific organic molecules

against degradation by microorganisms and enzymes. Chemical stabilization involves all

intermolecular interactions between organic substances and inorganic substances leading

to a decrease in availability of the organic substrate due to surface condensation and

changes in conformation, i.e., sorption to soil minerals and precipitation. Physical

stabilization is related to the decrease in the accessibility of the organic substrates to

microorganisms caused by occlusion within aggregates.

Recently, increasing evidence from studies in soils and sedimentary systems indicates

that sorptive protection of OM may be of particular importance, although chemisorption

of OM to clay-sized particles and physical protection of OM within organo-clay

aggregates often cannot be clearly distinguished. A first indication for the importance

of sorptive protection in soils is the frequently reported positive relationship between the

organic carbon (OC) content and the clay content (e.g., Burke et al., 1989; Hassink,

1997). Additional evidence comes from close relations between OC and BET surface

areas in coastal sediments (Mayer, 1994a; Keil et al., 1994) and subsoil horizons

(Mayer, 1994b), giving calculated surface loadings of 0.6–1.5 mg OC m� 2. These

loadings were considered to represent the ‘‘monolayer equivalent’’ (ME) range for OM

associated with mineral particles (Mayer, 1994a). Hedges and Keil (1995) assumed that

this finding is indicative of dissolved organic matter (DOM) sorption to mineral grains.

Keil et al. (1994) showed that simple desorption of OM from marine sediments with

water increased mineralization rates by up to five orders of magnitude. Similar results

were observed for soils by Nelson et al. (1994).

A reappraisal of the BET surface data using calculations of the reaction enthalpies

(Mayer, 1999; Mayer and Xing, 2001) and TEM studies (Ransom et al., 1997; Salmon

et al., 2000) showed that only a fraction of the mineral surface was covered with OM,

which occurred as patches and formed microaggregates with clay-sized particles. These

OM patches seem to be related to mineralogy. Ransom et al. (1998) reported that OM

appears to be preferentially sequestered in sediments, being rich in smectite but also in

metal oxyhydroxides, suggesting an influence of mineralogy on sorptive protection.

Likewise, Kaiser and Guggenberger (2000) calculated close correlations between

measures for Al and Fe hydrous oxides and OC concentrations in soils. In a

chronosequence study, Torn et al. (1997) identified the concentrations of short-range

ordered and noncrystalline minerals as the primary control for soil OC concentrations

and turnover.

G. Guggenberger, K. Kaiser / Geoderma 113 (2003) 293–310294

Together, these results suggest that stabilization of OM by interactions with distinct

mineral matrices may be the single largest factor controlling OM preservation on the

Earth’s surface today (Keil et al., 1994). Because sorption is defined as the transfer of a

solute (the sorbate) from solution to an existing solid phase (the sorbent) (Sposito,

1984), a prerequisite for the sorptive stabilization of OM is that it must occur in a

dissolved state prior to sorption. Likewise, if precipitation (the accumulation of a solute

as a new solid phase) is considered as a process of OM stabilization (Boudot et al.,

1989), the precursor substance must be dissolved. Hence, the paradigm of sorptive

stabilization includes a significant fraction (if not all of the OM) must have passed

through the dissolved phase before undergoing sorption or precipitation (Hedges and

Keil, 1995). This gives rise to the investigation of DOM sorption to minerals not only in

rivers and sedimentary systems but also in soils.

In this discussion paper, we will first summarize results from DOM research in forest

soils to show that DOM sorption to soil minerals contributes to the formation of stable

soil OM. Then we will challenge on the paradigm of sorptive stabilization. Using a

combination of field data on dissolved organic carbon (DOC) fluxes in forest soils and

laboratory data on the soils’ sorption capacity for DOC, we will provide evidence that

the turnover of sorbed OC may be quite rapid. We will present a line of evidence

showing that, depending on the type of surface, DOM sorption can be a stabilizing

process but it may be also a prerequisite for OM mineralization.

2. Processes of dissolved organic matter retention in forest soils

In the last two decades, many studies have dealt with the dynamics of DOM in forest

ecosystems. It has been reported that about 10–40 g DOC m� 2 year� 1 is translocated

from the organic surface layer into the mineral soil horizons (summarized in Michalzik et

al., 2001; Fig. 1). This means that about 10–25% of total C input to the forest floor with

litter fall is leached from the organic surface layers (McDowell and Likens, 1988;

Guggenberger, 1992). In deeper mineral soil horizons, the DOC fluxes decline to about

1–10 g m� 2 year� 1 (Michalzik et al., 2001), suggesting pronounced DOM retention in

the subsoil horizons. Most authors consider that the DOC fluxes at a soil depth of about

90–100 cm represent the DOC export by leaching.

Sorption studies using disturbed soil samples (Kaiser and Zech, 1998) and intact soil

columns (Guggenberger and Zech, 1992) suggest that the DOC retention occurs rapidly.

Kaiser and Zech (1998) showed that 60–90% of the added DOC was retained by subsoil

horizons within 15 min of addition to the soil. On the other hand, Qualls and Haines

(1992) observed that only 14–33% of the DOC collected from litter layers and mineral

soil horizons decomposed during a 134-day incubation. Kalbitz et al. (2003), using a

double exponential model, found that 94% of DOM derived from a spruce Oa horizon

was stable and had a mean half-life of 8.6 years. Rapid sorption combined with slow

microbial decomposition of the larger part of DOM suggests that sorption processes are

primarily responsible for the pronounced DOC retention in the mineral subsoil. This is

corroborated by the fractionation of DOM occurring during its passage through the soil.

The strongly colored hydrophobic DOM fraction is preferentially retained by the soil

G. Guggenberger, K. Kaiser / Geoderma 113 (2003) 293–310 295

matrix (Kaiser and Zech, 1997), whereas the less colored hydrophilic DOM fraction is

relatively enriched in the mineral soil output (Vance and David, 1991; Kaiser et al.,

2002a,b). Because mineralization primarily affects the hydrophilic DOM fraction (Qualls

and Haines, 1992; Kalbitz et al., 2003), any considerable contribution of mineralization

to the DOM retention would induce the relative enrichment of the hydrophobic fraction

in the mineral subsoil.

One possible group of sorbent for DOM is phyllosilicates. In the presence of metal ions

such as Ca2 +, Al3 +, Fe2+/3 +, the anionic organic materials may be bound to the negatively

charged clay surfaces by formation of cation bridges (Theng, 1976; Baham and Sposito,

1994). This process may be particularly important in soils with neutral to slightly alkaline

pH or very acidic pH, where the solute concentration of Ca2 + or Al3 + and Fe2+/3 + are

high. However, recently it appeared that the primary sorbents for DOM are minerals of

variable charge. Using multiple regression analysis, Moore et al. (1992) and Kaiser et al.

(1996) related DOC sorption to the quantities of Fe and Al hydrous oxides of soils. This

statistical relationship was corroborated by sorption experiments carried out on soil

material of a spodic B horizon with different quantities of amorphous Al(OH)3 and

goethite (a-FeOOH) coatings (Kaiser and Guggenberger, 2000). The increasing DOC

sorption at higher concentrations of oxalate-extractable Al (Alox) and of dithionite–

citrate–bicarbonate extractable Fe (FeDCB) may be partly explained by the strong increase

of the surface area at higher loadings of the hydrous oxides (Kaiser and Guggenberger,

2000). However, the charge of the surface is also altered by the oxide coatings. Aluminum

and Fe hydrous oxides possess a net positive charge at slightly acidic soil pH (Theng and

Orchard, 1995). Such positively charged surfaces are a prerequisite for a strong sorption of

Fig. 1. Schematic illustration of the fate of dissolved organic carbon (DOC) in forest soils.


the negatively charged organic polyelectrolytes. Indeed, the strong dependence of DOM

sorption on pH, the competition of DOM with specifically binding inorganic anions (e.g.,

phosphate), and the release of OH� during the sorption suggest that surface complexation

of functional groups via ligand exchange is the most important process in the sorption of

DOM on minerals and soils (Tipping, 1981; Gu et al., 1994; Weigand and Totsche, 1998).

Using DRIFT spectroscopy, Kaiser et al. (1997) provided direct evidence for a

chemisorptive bonding of DOM to hydrous oxides. Sorption of DOM to goethite and

amorphous Al(OH)3 resulted in a sharp decrease of the band at 1715 cm� 1 due to

protonated carboxylic groups. This was accompanied by a strong increase of the

carboxylate band together with a shift from 1625 to 1600–1605 cm� 1 and increasing

absorption at 1400 cm� 1. Such changes are due to complexation of carboxyl groups with

metals on the mineral surface resulting from ligand exchange reactions (Parfitt et al., 1977;

Gu et al., 1994). The same authors reported that the formation of bidentate complexes

between two organic ligands in ortho position (i.e., one carboxylic group and one phenolic

group) of an aromatic ring and a metal at the surfaces of oxides and hydroxides causes a

strong (innersphere) chemisorptive bonding. Indeed, increased absorption occurred also at

1270 cm� 1 indicating that phenolic groups are involved in the sorption of DOM on

hydrous oxide surfaces as well. Because the major part of the carboxylic groups participate

in the complexation reactions at the mineral surfaces, Kaiser et al. (1997) concluded that

each organic macromolecule is sorbed by many bonds. Sorption of DOM with multi-

dentate bondings per molecule, called the ‘‘octopus effect’’ (Podoll et al., 1987), might

alter strongly the conformation and electron distribution of the organic molecules (Stotzky

et al., 1996). Such changes may result in an inhibition of enzymatic decomposition of the

OM, caused by the inability of the enzymes either to detect or to react with the substrate

(Khanna et al., 1998). At high OM loadings, fewer carboxylic groups are involved in the

bonding than at low OM coverage, suggesting a weaker bonding and possibly less

conformational changes of the OM molecules.

The strong chemisorptive bonding coincides with a pronounced hysteresis of the

sorption. Gu et al. (1994) reported that about 60–90% of DOC sorbed to Fe oxides was

irreversibly bound. Likewise, Kaiser and Zech (1999) reported that 24 h after sorption,

less than 3% of sorbed OC could be released from goethite and amorphous Al(OH)3under solution conditions similar to those during the sorption step. Even at desorption

with H2PO4�, an anion that forms strong bondings on Al and Fe hydrous oxide surfaces

via complexation–ligand exchange, less than 50% of the sorbed hydrophobic DOC

could be desorbed (Kaiser and Zech, 1999). This proves the high affinity of the

hydrophobic DOM to Al and Fe hydrous oxide surfaces. In contrast, hydrophilic DOM

could be almost completely removed from the surface, suggesting weaker, and possibly

nonspecific bondings.

The strong and hardly reversible sorption of DOM to Al and Fe hydrous oxide

surfaces may help to explain the strong relationship between the concentration of

mineral-associated soil OM and the concentration of these mineral phases (Kaiser and

Guggenberger, 2000). Dissolved organic matter sorption is also considered to be

responsible for the strong resemblance in the structural composition of soil OM in an

alluvial B horizon as well as in the clay fraction of a topsoil with hydrophobic DOM

from the overlying forest floors (Kaiser and Guggenberger, 2000).


The solute composition strongly accentuate the sorption processes. Inorganic cations,

e.g., H+, Ca2 +, Al3 +, Ca2 +, highly influence the net electrical charge of DOM. Protonation

or binding of other cations to the acidic functional groups of OM reduces the negative

charge of the molecules and thus the water solubility (Tipping and Hurley, 1988; Tipping

and Woof, 1990). At high concentrations, cations may give rise to precipitation of DOM.

According to Dahlgren and Marrett (1991), precipitation of DOM is controlled by the

DOC/(Fe +Al) ratio and the pH in the soil solution. Critical ratios are 10 to 20 mol C

mol� 1 Fe plus Al (Buurman, 1985; Dahlgren and Marrett, 1991). Buurman (1985) argues

that in podzols more narrow ratios can occur due to buffer reactions of H+ with the

concurrent release of polyvalent metals and the mobilization of metals from the labile

metal pool in the course of the podzolization process. However, in subsoil horizons, such

narrow ratios may also be a product of DOC sorption leaving a solution behind that is

relatively enriched in polyvalent metals. Furthermore, at the given DOC concentrations in

the mineral soil input of about 20 to 50 mg l� 1, precipitation further requires a

concentration of polyvalent metals in the range of several micromoles per liter, which is

unlikely for most soils (Guggenberger, 1992).

In summary, there is ample evidence that sorptive preservation of OM is an important

process in the OM accumulation and stabilization in soil. One has to bear in mind,

however, that sorption experiments usually were carried out on juvenile minerals or

subsoil samples with little OM and showing little, if any, microbial activity. Furthermore,

the timescale of sorption experiments is usually in the range of several minutes to days.

This is too short to assess mineralization processes of the sorbed OM. Thus, the role of the

decomposer community in soil is neglected in these studies. For this reason, it is difficult

to scale up the laboratory results about sorptive stabilization of OM on free mineral

surfaces to the situation in the field.

3. Capacity of soils to sorb organic matter

In the previous section it was shown that relatively large quantities of DOM enter the

mineral horizons. Biodegradation of organic matter in the dissolved phase is too slow to

remove a large portion of the DOM percolating through the soil (Qualls and Haines, 1992;

Kalbitz et al., 2003), whereas sorption to mineral phases is an efficient sink for DOM in

subsoil horizons (Kaiser and Zech, 1998; Kaiser and Guggenberger, 2000). However,

sorption of DOM to mineral phases and mineral soil is not infinite but approaches sorption

maxima at large DOM additions (Kaiser and Zech, 1997). This enables an estimation of

the sorption capacity of soils for DOM.

To do this, we carried out sorption experiments as outlined by Kaiser et al. (1996) for

all mineral soil horizons of seven forest soils, including cambisols, podzols, an arenosol,

and a leptosol. Briefly, 75 ml of an artificial soil solution containing 0 to 101 mg OC l� 1

was added to 3 g of soil material. The solutions were zero-tension extracted from the

organic layers under spruce, pine, or beech, depending on the type of forests growing on

the soils. For the organic layer under spruce, the major cations in the solutions were NH4+

(0.26 mmol l� 1) and K+ (0.18 mmol l� 1); the major inorganic anions were SO42� (0.21

mmol l� 1) and H2PO4� (0.06 mmol l� 1). The ionic strength of the solutions obtained from


the organic layers under pine was 0.002 M with the major inorganic constituents NH4+, K+,

and SO42�. In the solutions extracted from beech litter, the ionic strength was 0.008 M; the

major inorganic solutes were Ca2 + and SO42�. Hence, in all cases it was assured that no

precipitation reactions took place during the sorption experiment. Suspensions were

shaken gently for 24 h, then passed through 0.45-Am membrane filters, and the filtrates

were analyzed for DOC. The sorbed amount of OC was calculated by the difference

between the DOC concentration in the initial solution and in the filtrate. Maximum

sorption was estimated from the initial mass plots (Nodvin et al., 1986) as that point where

further DOC additions did not increase the amount of DOC sorbed any more (Fig. 2).

Because the sorption maximum characterizes the amount of OC that can be sorbed in

addition to the OC already stored in a particular horizon, this signifies the available

sorption capacity. The available sorption capacity can be either related to mass per unit

fine earth ( < 2 mm) or calculated as stock in an individual horizon based on its thickness,

bulk density, and proportion of coarse fragments. This assumes that coarse fragments do

not contribute to the sorption of DOM. The sorption capacity of the whole soil was

estimated by adding up the sorption capacity of all the horizons.

It is clear from Table 1 that the available sorption capacity of the soils is limited.

Corroborating earlier results (Kaiser et al., 1996), the available sorption capacity was

largest for the B horizons (Fig. 2, Table 1). This can be explained by the enrichment of

the strongly sorbing Fe and Al hydrous oxides caused by in situ formation in cambic

horizons and by illuvial accumulation in addition to in situ formation in spodic horizons.

In contrast, A and C horizons showed a low available sorption capacity. In A horizons

this is due to the high concentrations of indigenous OM already present, whereas in C

horizons this is caused by low concentrations of Fe and Al hydrous oxides (Kaiser et al.,

1996). When scaled up over the whole profile, the available sorption capacity was about

300 to 400 g OC m� 2 in cambisols and podzols and about 40 g OC m� 2 in the leptosol

Fig. 2. Examples for the estimation of the sorption maximum of DOC to two horizons of a haplic podzol and a

rendzic leptosol, using the initial mass relationship (Nodvin et al., 1986).


and the arenosol. The latter soils are weakly developed with little accumulation of Fe

and Al hydrous oxides (Kaiser et al., 2002a,b).

When we now relate the available sorption capacity of the different horizons to the

amount of total C already stored in the horizons, it is obvious that in the topsoil horizons

Table 1

Available sorption capacity of seven soils for OC related to the storage of total OC and OC found in density

separates >1.6 g cm� 3 (i.e., the mineral-associated OC fraction)

Site/soil (FAO) Horizon Depth Available OC storage Available sorption capacity

(cm) sorption

capacity

(g m� 2)

Total OC

(g m� 2)

OC >1.6 g

cm� 3

(g m� 2)

% of

total OC

% of

OC >1.6

g cm� 3

Oberwarmensteinach/

cambic podzol

(Guggenberger and

Zech, 1993)

E

Bhs

Bw

C

0–10

10–35

35–75

75–90

6

110

211

42

2208

3840

1092

101

1632

3360

1092

101

0.3

2.9

19.3

41.0

0.4

3.3

19.3

41.0

Total 0–90 369 7241 6185 5.1 6.0

Wulfersreuth/

dystric cambisol

(Guggenberger and

Zech, 1993)

AE

Bsw

Bw1

Bw2

2BC

0–7

7–22

22–35

35–70

70–90

2

58

73

182

64

1940

2178

546

588

120

1490

1980

437

588

120

0.1

2.7

13.4

31.0

53.0

0.1

3.0

16.8

31.0

53.0

Total 0–90 380 5372 4615 7.1 8.2

Hohe Matzen/haplic

podzol (Guggenberger

and Zech, 1993)

Ah

E

Bh

Bs

Bws

C

0–10

10–20

20–26

26–36

36–68

68–90

1

4

9

36

145

139

1260

900

1320

2550

4066

924

300

360

1080

2150

3872

924

0.1

0.5

0.7

1.4

3.6

15.0

0.4

1.2

0.9

1.7

3.8

15.0

Total 0–90 334 11,020 8386 3.0 4.0

Betzenstein/rendzic

leptosol

(Kaiser et al., 2000)

Ah1

Ah2

C

0–10

10–25

25–90

2

8

35

9240

7452

4290

8003

7128

3900

0.0

0.1

0.8

0.0

0.1

0.9

Total 0–90 45 20,982 19,030 0.2 0.2

Sey bothenreuth/typic

arenosol

(Kaiser et al., 2000)

Ah

C1

C2

0–5

5–15

15–90

1

3

36

1050

160

1200

490

160

1200

0.1

2.0

3.0

0.2

2.0

3.0

Total 0–90 40 2410 1850 1.7 2.2

Waldstein/cambic

podzol (Kalbitz and

Matzner, unpublished;

Kaiser et al., 2002a)

AE

Bh

Bs

Bw

BwC

C

0–10

10–12

12–30

30–55

55–70

70–85

1

2

45

167

42

41

2660

911

4586

1960

144

144

878

765

4173

1823

137

137

0.0

0.2

1.0

8.5

29.0

28.5

0.1

0.3

1.0

9.2

30.7

29.9

Total 0–85 298 10,406 7913 2.9 3.7

Steigerwald/dystric

cambisol (Solinger

et al., 2001;

Kaiser et al.,

2002a)

Ah

Bw1

Bw2

BC

Total

0–5

5–24

24–50

50–80

0–80

2

150

127

133

411

3630

3663

832

510

7635

1271

2198

707

485

4660

0.0

5.6

15.3

26.0

5.4

0.2

6.8

18.0

27.4

8.8


the percentage of the total sorption capacity that is available is < 1% (Table 1). This

suggests that the capacity of these horizons to store chemically stabilized OC (Sollins et

al., 1996), i.e., the OC of the >1.6 g cm� 3 fraction, is almost exhausted and agrees well

with field data, showing that A horizons are a source rather than a sink for DOC (Kaiser et

al., 2002a,b). With increasing soil depth, the percentage of the available sorption capacity

increases by up to 50% in subsoil horizons of well-developed soils. From this it may be

concluded that the subsoil horizons have a high potential for sorptive retention of DOM.

The weakly developed rendzic leptosol and typic arenosol are the only exceptions. In these

soils, the contents of Fe and Al hydrous oxides are low throughout the whole profile

(Kaiser et al., 2002a,b) and this limited amount of sorbents appears to be almost

completely occupied by OC. Consequently, the DOC concentration in the mineral soil

output is quite high with 19 g l� 1 (rendzic leptosol) and 25 g l� 1 (typic arenosol) (Kaiser

et al., 2002a,b). In general, there seems to be a negative relationship between the

percentage of the total sorption capacity for OC that is still available and the DOC losses

of the soils (Fig. 3). These results suggest that the available sorption capacity of the subsoil

horizons may be an important variable that governs the DOC output from soil.

Considering that the OM which is sorbed to the minerals must be previously dissolved,

the OC storage and the available sorption capacity can be related to the DOM dynamics in

the field (Fig. 4). The annual DOC retention in the seven soils, as calculated by subtraction

of the DOC fluxes at 80 to 90 cm depth from those in the mineral soil input, ranged from 4

to 31 g m� 2 (Table 2). This variation can be explained by factors controlling DOC

mobilization in the organic layers such as temperature and precipitation regime and the

litter quality, and the efficiency of DOC immobilization in the mineral soil horizons

primarily due to sorption to soil minerals (for a more detailed discussion we refer the

reader to the original papers cited in Table 1 and to the review of Kalbitz et al., 2000).

Neglecting other C input to the mineral-associated OM such as decomposition products of

Fig. 3. Available sorption capacity of the whole profile and the deepest mineral soil horizon related to log DOC

output of the seven soils shown in Table 1.


roots or microbial metabolites, a simple relation of the annual DOC retention (g OC m� 2)

in the mineral soil horizons to the mineral-associated OC (i.e., OC in the fraction >1.6 g

cm� 2) stored in the soil (g OC m� 2) provides an estimate for the residence time of the

indigenous mineral-associated OC in the range of 186 to 1730 years (Table 2). From this,

it can be concluded that the soils store and stabilize a large amount of the incoming OC

very efficiently.

A different picture emerges when the DOC retention is related to the available sorption

capacity. The estimated number of years until exhaustion of the sorption capacity for OC

of the seven soils ranged from 4 to 30. When we now consider that these forest soils are

close to a steady-state equilibrium with respect to carbon sorption and mineralization, the

amount of new OC sorbed within this period of time is in the order of the amount of

sorbed OC being mineralized. This means a surprisingly low mean residence time of the

sorbed OC of < 50 years, especially in the biologically active rendzic leptosol. Even if we

consider that not all of the DOC retained is sorbed but that it is partly mineralized from the

solution, this will add only a few years to the estimated mean turnover time of the sorbed

OC. Of course, a complete saturation of the soils with OC is not possible. Depending on

the distribution coefficient, at a given solution concentration of OC there is a unique

concentration of sorbed OC. Thus, the available sorption capacity will never be filled up

completely. Consequently, the number of years for the estimated mean turnover time of the

sorbed OC must be reduced. Although considering a quite large error in the estimates of

the mean turnover time of the currently sorbed OC, it is obvious that it is in the range of

decades or less.

Summarizing the results, strong discrepancies on the sorptive stabilization of OM

in forest soils are observed. Sorption experiments on minerals and subsoil horizons

Fig. 4. Annual DOC sorption in the mineral soil horizons of a haplic podzol related to the available sorption

capacity.


suggest strong chemisorptive retention of OM. This is corroborated by the high

residence time of indigenous OC already sorbed to the minerals as calculated from

the relation of the yearly DOC input and the stocks of mineral-associated OC. On

the other hand, the available sorption capacity of forest soils is limited, in particular

in the topsoil and upper subsoil horizons. Our OC budget estimates suggest a

relatively short mean residence time of the OC currently entering the mineral soil as

DOM.

As mentioned above, sorption experiments that have been carried out may not be well

suited to assess mineralization processes. This is further complicated by the spatial

heterogeneity of soil with respect to microbial biomass and activity. In all soils, there

are horizons or microsites having high microbial biomass and activities that may affect the

sorptive retention of DOM differently than locations with low microbial activity and large

amounts of juvenile mineral surfaces.

Table 2

Available sorption capacity of seven soils for OC as measured in the laboratory related to the published annual

DOC retention in the mineral soil horizons as determined in the field

Site/soil (FAO)/

reference for annual

DOC retention

Horizon Depth

(cm)

Available

sorption

capacity

Annual

DOC

retentiona

Estimated

residence

timeb

Time until

exhaustion

of sorption

g kg� 1 g m� 2 (g m� 2) (years) capacityc (years)

Oberwarmensteinach/ Total 0–90 n.d. 369 16 386 23

cambic podzol

(Guggenberger

and Zech (1993)

(E to C)

Wulfersreuth/dystric Total 0–90 n.d. 380 14 330 27

cambisol (Guggenberger

and Zech, 1993)

(AE to 2BC)

Hohe Matzen/haplic podzol Total 0–90 n.d. 334 31 270 11

Guggenberger

and Zech, 1993)

(Ah to C)

Betzenstein/rendzic leptosol Total 0–90 n.d. 45 11 1730 4

(Kaiser et al., 2000) (Ah1 to C)

Sey bothenreuth/typic arenosol Total 0–90 n.d. 40 4 463 10

(Kaiser et al., 2000) (Ah to C2)

Waldstein/cambic podzol Total 0–85 n.d. 298 10 791 30

(Michalzik and

Matzner, 1999)

(AE to C)

Steigerwald/dystric cambisol Total 0–80 n.d. 411 25 186 16

(Solinger et al., 2001) (Ah to BC)

n.d. = not determined.a Annual DOC retention by the soil is based on the assumption that DOC fluxes in the soil are primarily

controlled by sorption. It is calculated by subtracting the DOC output from the soil measured at a 80 to 90 cm

depths from the DOC fluxes entering the mineral soil. Data are taken from the references shown in the first

column.b Estimated residence time is calculated as OC (>1.6 g cm� 3) storage divided by annual DOC retention.c Time until exhaustion of sorption capacity is calculated as available sorption capacity divided by annual

DOC retention.


4. Sorption of organic matter to surfaces differing in the microbial activity

In podzols, sorption of DOM has been considered an important process in the formation

of OM in subsoil horizons. The Bh horizon is generally thought to result from DOM

sorption. Often, Bh horizons have large OC concentrations, which is a prerequisite for the

high density of microorganisms found in these horizons (McKeague et al., 1986; Fig. 5).

Concurrently, the 14C age of OM shows a minimum in the Bh horizon and is only about

100 years (Rumpel et al., 2002; Fig. 5). This is about the same order of magnitude as the

mean residence time of OM estimated in Table 2, and suggests a rapid mineralization of

the OM in the Bh horizon. The substrate utilized by the microbial community can be

efficiently replenished by the large DOM input. At large OM loadings, which are typical

for Bh horizons, not all functional groups are involved in the chemisorptive bonding, and

steric changes of the substrate are probably not as pronounced than at smaller OM loadings

(Kaiser et al., 1997). Hence, the stabilizing effect of sorption on microbial degradability

may not be as efficient than at lower surface loadings. In the deeper soil horizons, the

percentage of available sorption sites is larger (Table 1). Accordingly, almost all functional

groups are involved in the bonding (Kaiser et al., 1997). Recent results further suggest that

at small loadings most OM is associated with micropores and/or other high-surface sites

(Kaiser and Guggenberger, 2003). This strong sorption in the horizons beneath the Bh

horizon leads to a more pronounced stabilization of OM, which is illustrated by the high14C age of the OM (Fig. 5).

Such compartmentalization is not only important between different horizons but also

with respect to the pathway of water, ions, and DOM fluxes in soil. Preferential flow paths

have been identified as biological hot spots (Bundt et al., 2001a,b). Compared to the soil

Fig. 5. Microbial density (left; from McKeague et al., 1986) and 14C age (right; from Rumpel et al., 2002) in a

podzol. The measured 14C activity was corrected for isotope fractionation and given in percent modern carbon

(pMC) according to the international standard. The radiocarbon age in years BP was calculated according to

Stuiver and Polach (1977). Please note that both analyses were carried out at two different soil profiles.


matrix, soil material collected from preferential flow paths is enriched in C and in different

parameters assigning the microbial biomass (Fig. 6). This reflects to more favorable living

conditions for the microorganisms in terms of nutrient and substrate supply. The higher

organic matter content in the preferential flow paths could be partly due to a high input of

rhizodeposition (Bundt et al., 2000). Kaiser et al. (2002) showed that during rainstorm

events, fluxes of DOC and dissolved organic nutrients occur primarily along preferential

pathways. It is thus likely that preferential input of DOM from the forest floor along flow

paths and its sorption at the respective soil surfaces is a prerequisite for the formation of

these biological hot spots. After formation of the biofilms, DOM sorption to the micro-

bially active surfaces provides an important substrate supply for the microbial community.

Also at smaller scale, heterogeneity of sites with different microbial activity needs to be

considered. Foster (1988) reported that bacteria living at the outer surface of micro-

aggregates and occupying pores >6 Am are large while bacteria living in microaggregates

often are small and many exist in a nongrowing or dormant state. Killham et al. (1993)

suggested that bacterial cells at the surface of microaggregates are able to intercept most

soluble organic substrates that diffuse within the soil matrix. Thus, they efficiently prevent

readily decomposable substrates from reaching microorganisms in the interior of micro-

aggregates. From the viewpoint of sorptive stabilization, this means also that micro-

organisms located on aggregate surfaces prevent DOM from reaching mineral surfaces in

the aggregate interior where it can be chemically bound.

Regardless of whether the soil microorganisms are existing in small colonies on

microaggregate surfaces or as complex communities living in biofilms along preferential

pathways they must compete in their interception of DOM with the sorption of DOM to

the minerals. Biofilms consist of bacterial cells growing within an organic matrix of

extracellular polymeric substances (EPS) synthesized by attached bacteria. Their structure

is similar to a porous gel containing 90–95% water and is characterized by interstitial

voids, channels, and cell clusters (Lewandowski et al., 1994). Because the EPS contains

Fig. 6. Depth profiles of concentrations of total carbon, microbial carbon (Cmic), DNA, and cell counts

(DAPI = 4V,6-diamidino-2V-phenylindole) in preferential flow paths and matrix soil (n= 4 and error bars repre-

senting the standard error) (from Bundt et al., 2001a,b).


large concentrations of sugar acids, biofilms usually possess a negative charge (Morgan et

al., 1990). The net negative surface charge of biofilms is likely to decrease sorption of

negatively charged OM to the biofilm and its diffusion within the biofilm (Carlson and

Silverstein, 1998). Recently, Lunsdorf et al. (2000) identified microgranular clusters of

Fe–O material and phyllosilicates directly associated with the bacterial cell envelope in

electron spectroscopic images. Within the biofilms, the authors also found granules of

storage OM and concluded that the minerals within the biofilm may remove OM from the

surrounding aqueous media and transport the organic substrate to the decomposers.

Sorption of DOM to positively charged Fe hydrous oxides associated with biofilms

may thus be an efficient pathway for OM uptake into biofilms, where the organic substrate

can be utilized by the microbial consortium. Interestingly, the active mineral compound

involved in the sorption into biofilms appears to be the same as for whole soils, i.e., Fe

hydrous oxides of low crystallinity.

5. Conclusion

Based on these findings, we propose a model shown in Fig. 7 where the location of the

OM is decisive for its fate. In forest soils, organic matter in the soil solution is mineralized

at rates that are much slower than the mean residence time of DOM in the mineral soil

(e.g., Qualls and Haines, 1992). This suggests that mineralization cannot be responsible

for the DOM retention in the mineral soil. Kalbitz et al. (2003) identified that the stable

DOM component, which dominates OM in the solution percolating into the mineral soil, is

associated with high UV absorbance and high aromaticity. According to Guggenberger et

al. (1994), such structures can be best described as lignocellulose-degradation products.

Fig. 7. Conceptual model of the fate of OM in the soil solution and sorbed to soil surfaces differing in their

biological activity.


For a complete degradation of such compounds, a consortium of microorganisms is

necessary (Shevchenko and Bailey, 1996). However, the density of bacteria in the soil

solution is low and individuals are physically separated. Hence, there is little possibility

for degradation of complex organic molecules in the soil solution.

However, these DOM components have a large affinity for sorption to positively

charged mineral surfaces, i.e., Fe and Al hydrous oxides (Kaiser et al., 1997). The sorption

is predominantly a chemisorptive process that goes along with changes in conformation. In

case of juvenile surfaces, part of the sorbed OM seems to be located in small pores that are

inaccessible for microorganisms (Kaiser and Guggenberger, 2003). Hence, OM sorbed to

juvenile surfaces is efficiently stabilized against microbial mineralization and sorptive

preservation is certainly taking place. According to Table 1, such conditions rarely occur

in topsoil horizons but primarily in subsoil horizons.

With time, more OM is sorbed to the mineral surfaces. The sorption capacity is almost

saturated and the bondings are not as strong as at smaller OM loadings (Kaiser et al., 1997).

This attracts heterotrophic microorganisms to colonize the OM-loaded mineral surfaces.

Cell colonies and biofilms develop at such sites of high nutrient and organic substrate

availability. These are topsoil horizons in general, but may also include preferential flow

paths and surfaces adjacent to interaggregate pores. Biofilms compete with inorganic

surfaces for OM sorption and are probably an effective sink for DOM in the mineral soils.

Iron hydrous oxides within the biofilms may serve as sorbents and OM shuttles to provide

substrate to the microbial community (Lunsdorf et al., 2000). Efficient sorption of OM into

the biofilms leads to a concentration of OM at this microsite. Likewise, the diverse

communities living in biofilms enable a more rapid mineralization of complex organic

molecules than of the same organic compounds in the soil solution. Sorption of OM into

biofilms may thus be a prerequisite for a quite rapid OM mineralization, and we conclude

that this is largely responsible for the short mean residence time estimated for the DOM

retained in the seven mineral soils shown in Table 2.

Therefore, to predict the fate of DOM in soil, the organic compounds not only need to

be classified into pools of different physicochemical and biochemical properties. It seems

even more important to address the probability of whether DOM is sorbed to juvenile

mineral surfaces or into biofilms. This requires detailed research on the flow paths of

DOM, the location and activity of biofilms, the sorption of DOM into and mineralization

by biofilms, and the location of active Fe and Al hydrous oxides in the natural soil

environment.

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dissolved organic matter in soil: challenging the paradigm of sorptive preservation

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