The dynamics of Si cycling during weathering in two small catchments in the Black Forest (Germany) traced by Si isotopes

Accepted Manuscript The dynamics of Si cycling during weathering in two small catchments in the Black Forest (Germany) traced by Si isotopes

Grit Steinhoefel, Jörn Breuer, Friedhelm von Blanckenburg, Ingo Horn, Michael Sommer PII: DOI: Reference:

S0009-2541(17)30384-4 doi: 10.1016/j.chemgeo.2017.06.026 CHEMGE 18383

To appear in:

Chemical Geology

Received date: Revised date: Accepted date:

26 January 2017 13 June 2017 19 June 2017

Please cite this article as: Grit Steinhoefel, Jörn Breuer, Friedhelm von Blanckenburg, Ingo Horn, Michael Sommer , The dynamics of Si cycling during weathering in two small catchments in the Black Forest (Germany) traced by Si isotopes, Chemical Geology (2017), doi: 10.1016/j.chemgeo.2017.06.026

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ACCEPTED MANUSCRIPT The dynamics of Si cycling during weathering in two small catchments in the Black Forest (Germany) traced by Si isotopes Grit Steinhoefel a,1, Jörn Breuer b,2, Friedhelm von Blanckenburga, Ingo Hornc, Michael Sommerd,e a

Deutsches GeoForschungsZentrum GFZ, Telegrafenberg, D-14473 Potsdam, Germany

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b

Landesanstalt für Landwirtschaftliche Chemie (710), Universität Hohenheim, Emil-Wolff-Str. 14, D-

Institut für Mineralogie, Leibniz Universität Hannover, Callinstr. 3, D-30167 Hannover, Germany

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c

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70593 Stuttgart, Germany

Institute of Soil Landscape Research, Leibniz-Centre for Agricultural Landscape Research (ZALF) e.V.,

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Eberswalder Strasse 84, D-15374 Müncheberg, Germany

University of Potsdam, Institute of Earth and Environmental Sciences, Karl-Liebknecht-Str. 24–25,

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14476 Potsdam, Germany Present Address: 1

G. Steinhoefel

Am Handelshafen 12

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Alfred-Wegener Institute

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Marine Biogeosciences

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D-27570 Bremerhaven, Germany 2

J. Breuer

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Landwirtschaftliches Technologiezentrum Augustenberg Neßlerstr. 25

D-76227 Karlsruhe, Germany

Corresponding author: Grit Steinhoefel email [email protected] phone + 49 (0)471 4831 2093 fax +49(471)4831-2020

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ACCEPTED MANUSCRIPT Abstract Silicon stable isotopes have emerged as a powerful proxy to investigate weathering because Si uptake from solution by secondary minerals or by the vegetation causes significant shifts in the isotope composition. In this study, we determined the Si isotope compositions of the principle Si pools within two small catchments located on sandstone and paragneiss, respectively, in the

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temperate Black Forest (Germany). At both settings, clay formation is dominated by mineral

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transformation preserving largely the signature of parental minerals with 30Si values of around -

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0.7‰. Bulk soils rich in primary minerals are similar to bulk parental material with 30Si values close to -0.4‰. Topsoils are partly different because organic matter degradation has promoted intense

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weathering leading to 30Si values as low as -1.0‰. Water samples expose highly dynamic weathering processes in the soil zone: 1) after spring snowmelt, increased release of DOC and high

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water fluxes trigger clay mineral dissolution which leads to 30Si values down to -0.7‰ and 2) in course of the summer, Si uptake by the vegetation and secondary mineral formation drives dissolved

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Si to typical positive 30Si values up to 1.1‰. Groundwater with 30Si values of around 0.4‰ records

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steady processes in bedrock reflecting plagioclase weathering together with kaolinite precipitation. An isotope mass balance approach reveals incongruent weathering conditions where denudation of

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Si is largely driven by physical erosion. Erosion of phytoliths contributes 3 to 21% to the total Si export flux, which is in the same order as the dissolved Si flux. These results elucidate the Si dynamics

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during weathering on catchments underlain of sedimentary origin, prevailing on the Earth surface and provide therefore valuable information to interpret the isotope signature of large river systems. Keywords Weathering, sedimentary rocks, biogeochemical Si cycle, Silicon isotopes, UV femtosecond laser ablation

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ACCEPTED MANUSCRIPT 1. Introduction Dissolved Si (DSi) is a major solute in rivers and is released by the complex interaction of abiotic and biotic weathering processes in the soil zone and underlying regolith and bedrock. Ratios of its stable isotopes provide the opportunity to characterise and quantify the associated weathering and transport processes (Frings et al. 2016). Clay formation incorporates preferentially light Si isotopes

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compared to the original silicate material, with the degree of fractionation depending on the mass

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balance and fractionation factors of the reaction as inferred from field observations, experimental

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and theoretical approaches (Bern et al. 2010; Cornelis et al. 2010; Georg et al. 2007, 2009a, b; Meheut et al. 2007; Meheut and Schauble 2014; Oelze et al. 2014, 2015; Opfergelt et al. 2010, 2012;

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Steinhoefel et al. 2011; Ziegler et al. 2005a, b). Adsorption of Si on Fe-oxides and amorphous Alhydroxides is likewise associated with the preference of the light Si isotopes in the adsorbed Si phase

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(Delstanche et al. 2009; Oelze et al. 2014; Opfergelt et al. 2009). Silicon is also a beneficial element for plants, which preferentially utilize light Si isotopes from soil solutions (Ding et al. 2008a, b;

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Opfergelt et al. 2006; Ziegler et al. 2005a). It is stored in plants in percent concentrations as bio-opal

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called phytoliths. Each of these processes, clay formation, Si adsorption and Si uptake by the vegetation fractionates light Si isotopes into solid phases driving DSi in fluids to higher 30Si/28Si ratios

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(denoted in permil as δ30Si values). Conversely, dissolution of clay minerals and phytoliths as well as desorption of Si release light Si isotopes into solution. Multiple cycles of dissolution/reprecipitation

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of secondary minerals and adsorption/desorption of Si are common during weathering and thus govern the Si isotope composition of DSi in soil solution and groundwater. Especially the amorphous silica pool (abiotic and biotic) appears to be a readily accessible pool with high turnover rates (e.g. Sommer et al. 2006, 2013). But, also diatom blossoms can have a large impact on surface waters as lakes and large river systems sequestering preferentially light Si isotopes into the diatom pool (e.g. Alleman et al. 2005). The complex interaction of abiotic and biotic processes causes a wide range for continental water reservoirs with soil solution and forest floor leachates ranging from about -2.1 to 2‰ (Cornelis et al. 2010; Frings et al. 2016; Pogge von Strandmann et al. 2012; Ziegler et al. 2005a), 3

ACCEPTED MANUSCRIPT groundwater between -1.4 and 0.4‰ (Georg et al. 2009a,b; Opfergelt et al. 2013; Pokrovsky et al. 2013) and rivers and lakes from -0.1 to 4.7‰ (Alleman et al. 2005; Cardinal et al. 2010; Cockerton et al. 2013; De la Rocha et al. 2000; Delvaux et al. 2013; Ding et al. 2004, 2011; Engstrom et al. 2010; Fontorbe et al. 2013; Georg et al. 2006a, 2007; Hughes et al. 2011b, 2012, 2013; Opfergelt et al. 2013; Pokrovsky et al.2013; Sun et al. 2013; Ziegler et al. 2005a). To examine sources of DSi together

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with driving weathering processes leading to this range in riverine DSi, small well-investigated

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catchments provide ideal field laboratories.

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Most studies using Si isotopes as tracer for weathering have focused on catchments underlain by igneous rocks. However, about two thirds of the continents are covered by sedimentary rocks, where

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sandstone and shale make up about 50% of all exposed rocks (e.g. Amiotte-Suchet et al. 2003; Dellinger et al. 2014). Thus, we have studied weathering processes by characterizing the Si isotope

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composition of the major Si pool in rocks, soils and water in a small catchment underlain by a mixture of sandstone and siltstone in the temperate climate of the Black Forest (Germany). Additionally, we

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investigate the Si isotope dynamics on a hillslope developed on paragneiss in a nearby field site.

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Steinhoefel et al. (2011) presented femtosecond laser ablation isotope data from in situ high resolution Si isotope analyses in weathering environments, together with first Si isotope data from

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soils from both catchments. This study extends the previous Si isotope data on soils and includes data from the analysis of water samples. We focus on the Si cycling during dissolution/reprecipitation

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processes of primary and secondary minerals in the weathering zone and quantify the chemical and erosional weathering fluxes using an isotope mass balance approach. 2. Site descriptions and sample materials The field sites of this study, two small well-investigated catchments on sandstone and paragneiss, respectively, are located within the cold, perhumid Black Forest, Germany (e.g. Fiedler et al. 2002; Sommer et al. 2001; Steinhoefel et al. 2011). The climate is similar at both locations with a mean annual temperature of 6°C and evenly distributed precipitation throughout the whole year. Soils 4

ACCEPTED MANUSCRIPT have developed from periglacial debris since at least 10 ka (Raab et al. 2007). Denudation rates from cosmogenic nuclides for the former non- or little glaciated parts of the Black Forest vary between 11 and 55 mm ky-1 for sandstone areas and 24 and 82 mm ky-1 for crystalline areas, respectively and decrease with decreasing slope (Morel et al. 2003; Meyer et al. 2010). The vegetation has been influenced by anthropogenic impact at least since the 12th century and is dominated today by

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cultivated Norway spruce forest (Picea abies (L.) H. Karst). Steinhoefel et al. (2011) presented first Si isotope data for two Cambisol profiles together with mineralogical investigations from both

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catchments. The data include Si isotope ratios of bulk soils (< 2 mm) from different horizons,

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separated clay fractions, parental mineral from rock fragments and phytolith samples. Investigated

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soil properties for these profiles include soil organic content (SOC), pH, crystallinity of Fe and Si, grain size distribution and the mineralogy of the clay fractions.

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The sandstone catchment Seebach covers an area of about 4.3 km2 in the Northern Black Forest (Fig. 1a) and have been investigated in detail by various studies (e.g. Georgiadis 2011; Hinderer 1995;

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Hinderer et al. 1998; Schlöser 1991; Sommer et al. 2001; Steinhoefel et al. 2011). The mean annual

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amount of precipitation is 1935 mm y-1 with an estimated total runoff of 1300 mm y-1. Most of the catchment is underlain by Triassic sandstone (Buntsandstein) characterized by sequences of reddish

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conglomerate and sandstone with intercalated siltstone layers and lenses. The lowermost part of the catchment consists of Forbach granite. In general, north-facing slopes are dominated by Stagnosols

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whereas south facing slopes are covered predominately by Podzols featuring evidence for vertical as well as lateral translocation of micro-sized particles and solutes along the slopes (Sommer et al. 2000; 2001). Cambisols occurs at lower, flat landscape positions on granite and feldspar-rich sandstones. Soil formation processes and soil characteristics including mineralogy have been investigated in detail (Georgiadis 2011; Georgiadis et al. 2013, 2014; Sommer et al. 2001). Representative soil profiles for Si isotope analysis (Fig. 1a) were selected from three soil profiles investigated at each location after carefully evaluation of physical and chemical properties and soil augering. Soil profile Sc1 (E-Podzol) developed on the stratigraphic unit “Middle Buntsandstein” 5

ACCEPTED MANUSCRIPT (smb-sst) consisting of 77% quartz and 22% alkali feldspar (orthoclase). The soil profiles downslope Sc2 (Bs-Podzol) and Sc3 (Cambisol) are underlain by “Lower Buntsandstein” (su-sst) which contains in average 62% quartz and 36% alkali feldspar (orthoclase). Interlayered siltstone consists almost exclusively of illite. Bulk soils are dominated by the sand fraction, which contains > 90 wt% quartz whereas the remainder is comprised of mainly alkali feldspar. Excavated sandstone blocks reveal a

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similar composition but contain slightly more feldspar. Clays make up about 15% of the soils. Their content decreases with depth. In soils, we assume that the majority of clays is sedimentary illite

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incorporated from the bedrock’s siltstone layers or derived from it. The soil clay assemblage consists

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of variable amounts of illite (20 - 80%), mixed layered illite/vermiculite (10-40%), vermiculite (< 20%),

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kaolinite (10 - 30%) and minor chlorite. Al-hydroxy interlayered vermiculite (HIV) makes up to onethird of the clay minerals in soil profile Sc2 downslope. Detailed clay mineralogy for the soils

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investigated in the profiles Sc1 and Sc2 is presented in Sommer et al. (2001) and for the profile Sc3 in Georgiadias et al. (2014). Clay formation in soils is characterized by following mineral

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transformations: 1) feldspar → kaolinite, 2) illite → vermiculite, 3) vermiculite → HIV and 4)

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vermiculite and HIV → kaolinite + Fe-oxyhydroxide. Based on the water chemistry of springs and streams, Hinderer (1995) distinguishes between surface water flowing through the soil zone and upper periglacial layers, subsurface water originated from underlying pre-weathered periglacial

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debris layers and base flow passing fractured sandstone and granite aquifers (Table 1). Depending on

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the seasonality and actual hydrological condition, the stream Seebach is fed by variable proportions of these water sources. Extreme conditions are given after snowmelt or intense rain event with up to 95% surface flow and during dry periods with almost only base flow giving an annual average of 46% surface/ subsurface water and 54% base flow downstream at monitoring point Sb13. The Herrenwieser See represents a small lacustrine component within this system. The study area on paragneiss, Wildmooswald, is a small catchment of about 0.065 km2 on a northfacing hillslope within the Central Black Forest (Fig. 1b) (Fiedler et al. 2002; Georgiadis2011; Jungkunst et al. 2008; Lamers et al. 2007; Steinhoefel et al. 2011). The mean annual precipitation is 6

ACCEPTED MANUSCRIPT 1600 mm y-1 with a discharge rate of 900 mm y-1. The bedrock is Paleozoic paragneiss and consists on average of 50-60% feldspar, which is mainly plagioclase besides minor orthoclase, 10-20% quartz, 1020% biotite and rare hornblende (Wimmenauer and Schreiner, 1999). A catenary sequence of different soils has developed from pre-weathered periglacial layers: Cambisol, Stagnosol, Histosols and patches of reddish-brown Cambisols showing lateral Fe transport by Fe reduction and

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subsequent accumulation as Fe(oxy)-hydroxides downslope. Weathering processes and typical soil characteristics have been described by Fiedler et al. (2002). Bulk soils are dominated by the sand

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fraction making up over 50% whereas the content of clay-sized particles is about 10 - 20%, which

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decreases with depth. Higher clay contents are observed in organic rich horizons. Soil clays consist

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typically of 10-35% vermiculite, 20-40% Al-hydroxy interlayered vermiculite (HIV), up to 20% illite and about 20% kaolinite. Clays have formed mostly by biotite weathering and to less extent by feldspar

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weathering. Most water is either surface runoff or subsurface water flowing through pre-weathered periglacial layers as the boundary to bedrock prevents largely deep water percolation. The hillslope is

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drained by two small springs, Wm-Q and Wm-QC integrating over the whole area. The latter one is

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characterized by prominent Fe-oxide precipitation. An artificial trench (Wm-G) and a small raised bog (Wm-M) provide access to water from the soil zone.

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To complement data from the previously studied Cambisols (Steinhoefel et al. 2011), we generated Si isotope data from catenary soil sequences. We used the least weathered rock fragments

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incorporated within the soil profiles to provide the best estimates for the parental minerals to these soils. Water samples were taken in streams, spring outlets and a bog twice in the year 2009, during spring in April/May after snowmelt and during late summer in September at the end of the growing season (Table 1). An exception is the spring Wolfsbrunnen (Sb-WB), which is located outside of the sandstone catchment but represents the fissured aquifer of the Forbach granite draining into the stream Seebach. Water samples for Si isotope analyses were filtered (0.45 µm) and acidified with HCl. 3. Methods 7

ACCEPTED MANUSCRIPT 3.1 Soil analyses Soil types and soil horizons were named according to the World Reference Base for Soil Resources (IUSS Working Group WRB 2014) except for Podzols (Sommer et al. 2001; Sommer et al. 2000). The analyses to characterize soil properties were described in detail in Steinhoefel et al. (2011) and are briefly summarized here. Bulk soil samples were air dried, gently crushed and sieved at 2 mm. Bulk

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soil samples were digested after incineration to remove organic matter by alkaline fusion using

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lithium metaborate and dissolved in 0.5 M HNO3. The element concentrations were determined

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using inductively coupled plasma optical emission spectrometry (ICP-OES) (Varian Vista Pro). Extractions using dithionite and dark acid-oxalate reveal the low crystalline Fe and Si fractions (Feo,

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Sio) and higher crystalline Fe and Si fractions (Fed, Sid), respectively. The particle size distribution of the bulk fine soil samples (< 2 mm) was determined using wet sieving and the pipette method after

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treatment with H2O2 (10 vol.%) to remove organic matter and dispersion in 0.01 M Na4P2O7. To investigate the Si isotope composition of clays, the coarse to medium clay (0.2 to 2 µm) and the

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fine clay (< 0.2 µm) fractions were separated as follows. Bulk soil samples (< 2 mm) were treated with H2O2 (10 vol. %) to remove organic matter and were afterwards shaken in deionized water to disintegrate aggregates. The suspension was diluted in 1 L deionized water and pH was adjusted to 8

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by NH3 to facilitate dispersion of Fe-(oxy)hydroxide-rich samples. After an appropriate time for sedimentation of particles > 2 µm the supernatant suspension was carefully removed and stored in

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plastic bottles. This procedure was repeated until the supernatant suspension remained clear. The fine clay fraction (< 0.2 µm) was separated from the coarse to medium clay fraction (0.2 to 2 µm) by a similar procedure but by using a centrifuge to obtain higher gravity. The fine clay was flocculated from the supernatant by adding MgCl2. The separated clay fractions were suspended in deionized water, dialyzed against deionized water to remove salts, dried and ground in a mortar made of ZrO2. For some soil samples, the fine clay fraction could not be separated because of strong aggregation with Fe-(oxy)hydoxides, which were not removed by this procedure. Clay minerals are considered as

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ACCEPTED MANUSCRIPT the dominant Si pool in these clay fractions, however clay-sized Fe-oxides, primary minerals and amorphous silica phases are also abundant (Georgiadis, 2014).

3.2. Water chemistry The characteristics of the water samples including temperature, pH, conductivity and major anion

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and cation contents have been measured directly in the field or contemporary in the laboratory of the University of Hohenheim on filtered water samples. The concentration of Al, Ca, Fe, K, Mg, Na

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and Si were determined by ICP-OES (Varian Vista Pro) whereas elements with low concentrations as

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Mn and Sr were measured by ICP-MS (Perkin Elmer Sciex Elan 6000). The anions Cl-, N03-, PO43- and SO42- were determined by ion chromatography (Dionex ICS 2000). The dissolved organic carbon

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(DOC) content was investigated by liquid chromatography-organic carbon detection (LC-OCD) at the

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GFZ Potsdam (Huber et al. 2011).

3.3. Si isotope measurements

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3.3.1. Sample Preparation of water samples

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Si isotope ratios of water samples were measured after employing a fusion technique (Georg et al. 2006b) to destroy organic components followed by Si purification using cation exchange

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chromatography. High concentrations of DOC in natural samples for instance in soil solutions might induce difficulties during Si isotope analysis by multicollector inductively coupled mass spectrometry

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(MC-ICP-MS) such as matrix effects, interferences, clogging within the introduction system and deposits in the mass spectrometer (Hughes et al. 2011a; van den Boorn et al. 2009). Here we developed an effective method to remove high amounts of DOC from water samples by adapting the fusion technique for silicate digestion of Georg et al. (2006b). Water samples containing an equivalent of about 40 to 100 μg of Si were evaporated to about 1 to 10 ml in 80 ml Teflon vials on hot plates at about 110 °C. The residual solution was then transferred step-wise into 5 ml silver crucibles and evaporated completely. Residual strongly adhesive sample material was transferred 9

ACCEPTED MANUSCRIPT from the teflon vials with pure aceton into the crucibles followed by evaporation. To remove the organic carbon of the sample, the crucibles were covered with lids and placed into a muffle furnace which was heated from room temperature to 750°C within two hours. After cooling 1 ml of 1 N NaOH was added to each crucible, mixed with MilliQ water to reach the initial fill level and evaporated on hotplates. This procedure ensures that the residual of the water samples covering the interior of the

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crucible up to initial fill level is completely covered by NaOH salt to provide complete digestion by alkaline fusion. The following procedure is similar as described by Georg et al. (2006b). Samples were

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fused at 750 °C for 15 min. The fusion product was then dissolved in about 11 ml MilliQ water for 24

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h followed by about 11 ml 0.03 N HCl for 3h in darkness. Pure concentrated HCl was added to the

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final solution to adjust to a pH of 1.5 avoiding Fe-oxide precipitation. This procedure results in a solution of about 22 ml containing the cations of the original water samples including 3 to 5 ppm Si in

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0.03 N HCl, whereas DOC concentrations are usually close to the blank level. ICP-OES analysis demonstrated quantitative digestion of Si by alkaline fusion.

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Typical soil solutions under forest stands in temperate humid climate range between 50 and 150

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ppm for DOC with extreme values up to 500 ppm while Si concentrations vary between 2 and 60 ppm (e.g. Campbell and Beckett, 1988). To explore the effect of these DOC amounts to our method we

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added to test samples using 80 ml MilliQ water dissolved humic acid (Aldrich Chemistry) amounting to 25, 50, 100, 250 and 500 ppm DOC, respectively, which equals 2, 4, 8, 20 and 40 mg of organic

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carbon. These test samples were evaporated in silver crucibles. Measurements with LC-OCD demonstrated that the organic matter was reduced to blank levels for all test samples after 2 h in which the muffle furnace was heated up to 750 °C. Because there is no certified Si isotope standard of water available, we decomposed well characterized materials including IRMM-17, diatomite, Big Batch and BHVO-2 (e.g. Reynolds et al. 2006) using the fusion technique as described above. The digested materials dissolved in 0.03 N HCl were then treated in the same way as water samples and digested again. The separation of Si by ion exchange chromatography also followed the method of Georg et al. (2006b). 15 ml of the prepared sample solution containing ca. 3 to 5 mg L-1 Si were 10

ACCEPTED MANUSCRIPT loaded onto pre-cleaned columns filled with 1.5 ml of BioRad DOWEX 50W-X8 200-400 resin and Si was eluted with 5 ml Milli-Q water. Purification of the sample and a Si column yield of > 95% was checked by ICP-OES (Varian 720-ES) after column separation. Within analytical errors, all re-digested standard materials reveal identical isotope composition as the original solutions, which are comparable with published isotope values.

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3.3.2. Sample preparation of solid samples

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For Si isotope analysis of solids using fs laser ablation (LA) coupled to MC-ICP-MS, methods were

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investigated and described in detail by Steinhoefel et al. (2011). Briefly, sample powder of bulk soil samples (< 2 mm) were fused to glass beads using the iridium strip heater at the Max Planck Institute

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of Chemistry in Mainz (Germany), which is a flux-free method for bulk sample analysis of geological

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material by LA-ICP-MS (Stoll et al. 2008). Separated clay fractions and bulk soil samples (< 2mm) from the organic horizons were analyzed as pressed powder pellets obtained by a hydraulic press. Parental

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mineral phases are analyzed in situ using thin section of collected less weathered rock fragments.

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3.3.3. Si isotope analysis by MC-ICP-MS

The Si isotope ratios on water samples were determined on a Thermo Fisher Scientific Neptune with

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a Jet Interface upgrade at GFZ Potsdam. The measurements were performed as described in more detail by Oelze et al. (2016). All analyses were conducted in high-resolution mode (m/ΔmM > 5000,

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5%) allowing interference-free measurements of all Si isotopes. Measurements were performed using an H-skimmer cone and dry plasma conditions. Purified solutions were introduced into the mass spectrometer via a desolvation system (Apex, ESI) using a nebulizer with an uptake rate of 120 µl/min. Sample and standard solutions were doped with Mg and measured applying the standardsample-bracketing technique to correct for instrumental mass bias and drift. Standard and sample solutions contained equal amount of Si and added Mg of about 1 mg L-1, respectively, in 0.1 M HCl giving typical signal intensities of between 10 and 15 V for 28Si on a Faraday cup equipped with a 1011 Ω resistor. For standards and samples, Si and Mg isotopes were measured in dynamic mode for 30 11

ACCEPTED MANUSCRIPT cycles with an integration time of 4 s using an idle time of 3 s after magnet switching. Each sample was analyzed at least three times in at least two independent runs. The reproducibility of Si isotope analysis, based on multiple digestions and isotope analysis of standard rock materials and natural water samples, is ± 0.14‰ for δ30Si (2 SD) (Oelze et al. 2016). During the course of this study several reference materials including IRMM-17, Diatomite, BHVO-2 and Big Batch were dissolved and

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reprocessed together with the water samples. Their repeated analysis revealed 30Si values of -1.38 ±

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0.11‰ (2 SD, n =11), 1.27 ± 0.12‰ (2 SD, n =15), -0.29 ± 0.10‰ (2 SD, n =11), and -10.65 ± 0.11‰ (2 SD, n =9), respectively, which compares well with published values (see overview in Oelze et al. 2016

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and Schuessler and von Blanckenburg, 2014).

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Si isotope analysis on solid samples were performed on fused glass beads or pressed powder pellets using an UV femtosecond laser ablation system coupled to a ThermoFinnigan Neptune MC-ICP-MS at

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Institute for Mineralogy at the Leibniz University of Hannover (Germany). The analytical set up is described in detail by Horn and von Blanckenburg (2007). The analysis followed the analytical

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protocol of Steinhoefel et al. (2011) and was checked with a variety of reference materials (Chmeleff

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et al. 2008; Steinhoefel et al. 2011). The ICP-MS was equipped with a Ni X skimmer cone revealing typical signal intensities for Si28 of 10 to 15 V using a 1011 Ω resistor. The mass resolution was similar

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as for solution ICP-MS allowing interference-free measurements. Sample and standard materials were investigated in the raster-mode using a spot diameter of 30 µm and covering an area of about

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0.01 mm2. The external reproducibility for bulk analysis on fused glass beads and pressed powder pellets and for in situ analysis on thin sections at high spatial resolution is ± 0.3‰ for δ30Si (2 SD) as revealed by repeated measurements of standard and sample materials in several analytical sessions (Steinhoefel et al., 2011). The both types of analyses were conducted by standard-sample-bracketing using the quartz reference material NBS28 as bracketing standard. Several replicates were obtained for each sample material. All obtained Si isotope ratios are reported relative to the reference material NBS28 in delta notation expressed in per mil (‰): 12

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(1)

(2)

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3.4. Estimating mass loss by chemical weathering The relative loss of cations by chemical weathering can be quantified by the Chemical Depletion

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Factor CDF, which describes the time-integrated chemical mass loss relative to the total time-

(3)

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integrated denudation (Riebe et al. 2001a, b):

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Where, CZr,p and CZr,w are the Zr concentration of the parental and the weathered material, respectively. In both catchments, soils have been developed from preweathered periglacial debris

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layers (Fiedler et al. 2002; Sommer et al. 2001). Therefore, we assign the C horizons to represent the

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parent material onto which Holocene weathering was superimposed to form the soils. The calculated CDF values are minimum estimates, because the periglacial debris samples might have been experienced various degrees of Holocene weathering. A challenge is also the heterogeneity of clastic

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sedimentary and the hillslope position of most of the soil profiles. Especially in the sandstone

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catchment, the mixed lithology involving variable proportions of siltstone and sandstone of different compositions (Sommer et al. 2001) leads to variable chemical composition of the C horizon. The sandstone sites show also admixture of debris of quartzitic conglomerates from the uppermost part of the catchment. Because of this complexity we only consider the CDF values of the Cambisol profiles (Sc1 and Pc1) to gain an idea of the chemical weathering intensity for both catchments. These soils have been developed on relatively flat landscape positions and represents therefore onedimensional weathering processes.

3.5. Estimating Si distribution in soil profiles 13

ACCEPTED MANUSCRIPT The relative loss or gain of elements in soil profiles can be quantified by normalizing its concentration with an immobile element such as Ti, Zr or Nb using the mass transfer coefficient τi, j (Brimhall and Dietrich 1987; Chadwick et al. 1990):

(4)

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Where, the concentration C of the immobile element i and mobile element j refers to weathered

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material w and unweathered parental material p, respectively. The Si concentration is enriched for τ > 0 and depleted for τ < 0 relative to the parental material, respectively, whereas for τ = 0 there is

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neither enrichment nor depletion. For the various reasons discussed above, we consider only the τ

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values of Cambisol profiles (Sc1 and Pc1) by using Zr as immobile element and taking the average of C horizons as parental material to gain an idea of the Si distribution in soils.

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4. Results

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4.1. Water chemistry

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Water chemical data are presented in Table A.1 (Appendix). Water samples from the sandstone catchment reveal differences between surface water and subsurface/baseflow water and are

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consistent with long-term observations of Hinderer (1995). Surface and subsurface waters Sb1, Sb2, Sb3, Sb7 show typically low pH values around 4 (4.2 to 4.4 in spring, 4.8 to 5.8 in late summer), which

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increased to about 6 for typical base flow waters Sb8 and Sb-WB (pH = 5.6 to 6.0 in spring, pH = 5.8 to 6.2 in later summer). The concentration of DOC is highest for surface water sources representing the soil zone (e.g. Sb2 and Sb7) with up to 24 mg L-1 and decrease with depth of the water source. Surface water, Sb2 and Sb7 show higher concentration of Al, Fe and Mn compared to deeper water sources. An exception is spring Sb16, which drains a clay-rich sequence within the sandstone and show likewise high Al and Mn concentrations. Concentrations of other major elements such as K, Na, Mg, Ca and Si and anions including NO3-, SO42- and Cl- reveal more spatial variations.

Low

concentrations observed in surface waters during spring are due to a dilution effect after snow melt. 14

ACCEPTED MANUSCRIPT Springs Sb-WB represent the granitic aquifer underlying the sandstone and are characterized by higher Ca and Mg but lower Al concentrations compared to the sandstone draining water sources. At the outlet, the stream Seebach shows a very similar water chemistry including pH and DOC as Sb8, showing that base flow dominates during the time of sampling in spring and late summer. On the paragneiss hillslope, the water sources are dominated by surface and subsurface water. The

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pH values range between 4.2 and 6.3 with higher values in late summer. Most water samples except

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for Wm-OQ reveal high DOC values between 9.6 and 32.5 mg L-1. Despite a large range, most samples

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show a lower Si content in spring after snowmelt (12 to 143 µmol L-1) compared to late summer (67 to 365 µmol L-1). A distinct water source in this area is the ochre-rich spring (Wm-OQ) with Fe-oxide

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and -hydroxide precipitation. At the spring outlet the water is rich in Si (143 to 365 µmol L-1) but poor in dissolved Fe (< 0.35 to 34 µmol L-1) whereas downstream the Si content decreases (61 to 96 µmol

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L-1) but the Fe content can be much higher especially in late summer (141 µmol L-1).

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4.2. Soil properties and element loss by weathering

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The major elements of the different soil horizons and soil properties including grain size distribution, soil pH, organic carbon content and extractable Fe (Feo und Fed) and Si (Sio und Sid) fractions are

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reported in Tables A.2 and A.3 (Appendix). The element loss by chemical weathering is quantified by calculating CDF values for both Cambisol

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for horizons with organic content of below 10%. CDF values vary in the same range between 0.01 and 0.27 giving averages of 0.19 in the sandstone catchment and 0.10 in the paragneiss catchment, respectively (Table A.2 (Appendix)). These values are minimum estimates as the C horizons were used as best approach for parental material. Despite of different parental rocks, the degree of chemical weathering is similar at both sites and accounts for 10 to 20% of the total denudation in these landscapes. Calculated τZr,

Si

values below the organic horizons for these soil profiles show

similar ranges between -0.06 and -0.30 indicating an average Si loss of 22% in the sandstone

15

ACCEPTED MANUSCRIPT catchment and 15% in the paragneiss catchment, respectively, compared to parental material (Table A.2 (Appendix)).

4.3. Si isotope data of natural water samples The Si isotope data of the water samples are presented in Table 2 and depicted in Fig. 1 and 2. In the

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sandstone catchment, water flowing either through sandstone or periglacial debris shows 30Si values between 0.29 and 0.52‰. The sampled granitic aquifer reveals similar values between 0.34

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and 0.49‰. Water draining predominately the soil zone yields lower 30Si values during spring of

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between -0.66‰ and 0.07‰ and tends to higher values in late summer as up to 0.83‰. The run-off of the lake Herrenwieser See shows 30Si values of 0.10‰ to 0.20‰. The main stream Seebach

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reveals 30Si values of 0.42‰ during spring and 0.72‰ during late summer, respectively at the

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sampling site Sb13.

A similar pattern results from the paragneiss hillslope. Investigated water samples from a small bog

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and a trench, which are dominated by soil water reveal low 30Si values in spring (-0.01‰ and

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0.51‰, respectively) and high 30Si values during late summer (0.89% and 0.80%, respectively). Springs draining the whole area containing surface and subsurface water reveal positive values with

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little seasonal influence ranging between 0.51‰ and 1.08‰ in 30Si.

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4.4. Si isotope data of soils and parental minerals Isotope composition of the investigated soil profiles is presented in Table 3 and Fig. 3. The Si isotope data on the Cambisol profiles, Sc3 and Pc1, and on separated phytoliths from both locations have been published in Steinhoefel et al. (2011). For the sandstone catchment, investigation of various rock fragments reveal following average δ30Si values for parental minerals: quartz with -0.35‰, feldspar with -0.49‰ and illite with -0.68‰. Using the average estimated modal abundance of 70% quartz and 30% feldspar from the investigated samples and assuming a clay content of 10%, the average bulk parental material has an isotope composition of -0.42‰, which is close to those of the 16

ACCEPTED MANUSCRIPT lower most soil samples from the C horizon. Bulk soils (< 2 mm) are within a similar range between 0.50‰ and -0.24‰. The clay fraction 0.2 to 2 µm show δ30Si values between -0.74‰ and -0.51‰, the grain-size fraction < 0.2 µm reveal isotope composition between -1.05‰ and -0.84‰. In the paragneiss catchment, parental minerals investigated in rock fragments reveal average δ30Si values of -0.22‰ for quartz, -0.48‰ for feldspar and -0.69‰ for biotite, respectively. The calculated

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isotope composition of the parental material is -0.37‰ assuming average modes of 50% quartz, 40‰

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feldspar and 10% biotite. Bulk soils (< 2 mm) below the organic horizon range between -0.48‰ and

fractions yield δ30Si values between -0.34‰ and -0.81‰.

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5. Discussion

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0.05‰. Samples from the organic horizons yield δ30Si values as low as -0.97‰. Separated clay

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5.1. Preservation of initial Si isotope signatures during soil formation The formation of clay minerals favors the incorporation of isotopically light Si (see overview in

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Opfergelt and Delmelle 2012). Their isotope composition integrates over the time scale of

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pedogenesis. Results from clay-size fractions from both catchments reveal 30Si values lower than bulk soils ranging between -1.04 and -0.33‰. However, this signature is largely inherited from the

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minerals they are derived from. In the sandstone catchment, the majority of soil clays is a mixture of primary illite from parental siltstone and neoformed clays derived from it including illite/vermiculite

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integrates, vermiculite and HIV (Georgiadis et al. 2014; Sommer et al. 2001). On the paragneiss hillslope, the dominate soil clay phases are vermiculite, HIV and illite derived from biotite alteration (Fiedler et al. 2002). In both cases, Holocene clay formation is dominated by transformation processes of primary illite and biotite, respectively in which the original silicate structure remains unmodified (Acker and Bricker 1992; Wilson 2004). Therefore, bulk clays in soils inherit largely the isotope signature of parental illite and biotite, respectively, which carry similar isotope signatures of around -0.7‰ (Table 3). Lower δ30Si value down to -1.04‰ are attributed to the presence of kaolinite which contributes up to 20% in the clay fractions. Kaolinite precipitation followed both 17

ACCEPTED MANUSCRIPT feldspar dissolution and as the final weathering product of clay transformation can be associated with significant fractionation with the average of which is -1.5‰ lower in its isotope composition compared to parental minerals (Georg et al. 2007; Méheut et al. 2007; Ziegler et al., 2005b). At both settings, bulk soil below the organic horizons feature uniform Si isotope compositions of 0.31 ± 0.08‰ (2 SD, n = 9) in the sandstone catchment and -0.28 ± 0.16‰ (2 SD, n = 6) in the

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paragneiss catchment. These isotope signatures show no significant fractionation compared to the

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estimates for the parental material, i.e. underlying periglacial debris layers of -0.32 ± 0.01‰ (2 SD, n

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= 3) and -0.44 ± 0.12‰ (2 SD, n = 2), respectively and bedrock compositions of -0.42‰ for sandstone and -0.37‰ for paragneiss, respectively. The CDF values suggest that only about 20% of mass has

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been lost by chemical weathering from both substrates above the C horizon. Both observations are consistent with these soils being dominated by the sand fraction (70-90%, Table A.3 (Appendix)) that

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is comprised of mainly quartz and some feldspar (Fiedler et al. 2002; Sommer et al. 2001). Therefore, the isotope composition of these minerals of about -0.3 and -0.5‰, respectively, controls the

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signature of bulk soils in both settings (Table 3, Fig. 3).

5.2. Intense weathering in organic-rich soil horizons

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In the paragneiss catchment, bulk organic horizons reveal significantly lower 30Si values between 0.97 and -0.48‰ and higher clay contents compared to deeper soil horizons (Fig. 3, Table A.3

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(Appendix)) (Georgiadis 2011; Steinhoefel et al. 2011) which are likely strongly depleted in heavy Si isotopes. Under conditions with low pH values, a high moisture content, and high concentrations of organic acid, biotite, plagioclase and soil clays can weather rapidly to form isotopically light kaolinite (Acker and Bricker 1992; Giesler et al. 2000; Wilson 2004; Ziegler et al. 2005b). Phytoliths with an isotope composition of -0.4‰ represents a very minor Si pool in these soils. Therefore, their influence on bulk soil signatures is negligible (Steinhoefel et al. 2011). In contrast, in the sandstone catchment, clay abundance and Si isotope composition are homogeneous throughout the profiles including the organic horizons (Georgiadis 2011; Steinhoefel et al. 2011). 18

ACCEPTED MANUSCRIPT 5.3. Seasonal controlled processes in the weathering zone The isotope composition of water samples reflects Si release and precipitation processes that integrate over the residence times of soil pore water. Thus, they are suited to resolve seasonal variations such as observed in previous river studies (e.g. Cardinal et al. 2010; Engstrom et al. 2010; Georg et al. 2006a; Hughes et al. 2011b, 2013; Pokrovsky et al. 2013). Water samples in both

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catchments yield lower δ30Si values in spring compared to those in summer, which is most

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pronounced in the soil zone (see Fig. 1 and 2). In spring after snowmelt, a common phenomenon is

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intense organic matter degradation in the organic-rich upper horizons due to rising water tables, which releases high amounts of DOC. The combination of DOC release and undersaturation of soil

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pore waters facilitates dissolution of secondary and biogenic Si phases (Agren et al. 2007; Boyer et al. 1997; Kalbitz et al. 2000). Dissolution of isotopically light phases such as phytoliths, amorphous silica

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associated with Fe-oxides, and clay minerals is supported by low water δ30Si values that prevail where pH is low and DOC values are high (Fig. 4 a, b). That clays are dissolved is indicated by the

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inverse correlation between δ30Si values and Al/Na ratios in water (Fig 4c). Furthermore, calculated

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saturation indices (SI) using the software Phreeqc, version 3 (Parkhurst and Appelo 2013) (Fig. 5) show that chlorite and vermiculite become unstable under these conditions. Dissolution of phytoliths

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and Fe-oxides are expected to contribute to DSi as they represent readily accessible pools with fast dissolution kinetics (Fraysse et al. 2010; Gerard et al. 2008; Jones, 1998; Sommer et al. 2013). The

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effect of dissolution is less pronounced in the paragneiss catchment where prominent Fe-oxide precipitations in soils and spring outlet drives DSi to higher δ30Si values (Delstanche et al. 2009). The addition of low δ30Si to DSi by mineral dissolution arising from high rates of water infiltration and DOC concentrations has been identified before. Very low DSi δ30Si values between -1.38 to -2.05‰ together with low Ge/Si ratios in forest floor leachates from organic-rich soil horizons in a temperate granitic setting in France have been attributed to dissolution of secondary clay minerals (Cornelis et al. 2010). Low δ30Si have been found in organic-rich rivers (“Black rivers”) in the tropic Amazon basin and the Congo River system (Cardinal et al. 2010; Hughes et al. 2013). In contrast, preferential uptake 19

ACCEPTED MANUSCRIPT of isotopically light Si by both vegetation and secondary mineral precipitation due to oversaturation drive DSi in the soil zone to typical positive δ30Si values in summer (Fig. 2) (Ding et al. 2008b; Opfergelt et al. 2006; Ziegler et al. 2005a). Clay mineral formation is supported by low Al/Na ratios together with high 30Si values and calculated SI values indicating stable condition for most clay minerals and Fe-oxides (Fig. 4c, 5).

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The deeper-sourced waters, those that drain periglacial debris layers and bedrock in the sandstone

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catchment record steady weathering processes with no seasonal variations (Fig. 2). Groundwater

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contributes substantially to stream discharge (e.g. Andermann et al. 2012; Calmels et al. 2011) that is also observable for the stream Seebach for which groundwater makes up in average 50% of the

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stream water (Hinderer 1995). Detailed hydrochemical investigations indicated dissolution of feldspar in the sandstone units and additional of biotite in the underlying granite followed by

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kaolinite formation (Hinderer 1995). This weathering reaction is supported by Si isotopes because precipitation of light kaolinite with a fractionation factor of ca. -1.5‰ (Georg et al. 2007; Méheut et

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al. 2007; Ziegler et al., 2005b) leads to high 30Si values for DSi as revealed for deep-sourced water

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with an average of 0.40 ± 0.12‰ (2 SD, n = 8). Seasonal variations as observed in the Black Forest have been observed in large river systems under

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various climate conditions (Engstrom et al. 2010; Georg et al. 2006a; Hughes et al. 2011b, 2013; Pokrovsky et al. 2013; Mavromatis et al. 2016). In particular, arctic and subarctic watersheds disclose

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similar controlling mechanisms for the release of DSi: 1) low δ30Si compositions in spring is governed by mineral dissolution; and 2) increasing δ30Si values in course of the summer reflect secondary clay and amorphous silica precipitation and Si uptake by the vegetation (Engstrom et al. 2010; Pokrovsky et al. 2013; Mavromatis et al. 2016).

5.4. Quantification of controlling mechanism of DSi release An isotope mass balance model considering the distribution of Si from dissolving rock with 30Sirock into either secondary minerals with 30Sisec or the complementary dissolved compartment with 20

ACCEPTED MANUSCRIPT 30Sidiss can describe the observed Si isotope composition of the water samples. For this model, we assume the absence of fractionation during dissolution and a steady state system for which the regolith production rate equals the denudation rate. Such a system can then be consider as a “closed system” (Johnson et al. 2004) or “batch reactor” (Bouchez et al. 2013). Bouchez et al. (2013) investigated this model approach in detail for the weathering zone, which we simplified for the

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purpose of this study. (5) (6)

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Where fdiss describes the relative flux proportion remaining in solution after partial precipitation into secondary phases. Observed and experimental determined Si isotope fractionation for clay and oxide

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formation and for plant uptake range around -1.5‰ (see review of Opfergelt and Delmelle 2012), which we take as bulk fractionation factor for secondary mineral formation 30Sisec-diss to describe

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abiotic and biotic processes by the same model. Combining equation (5) and (6) shows that 30Sidiss is

(7)

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into secondary minerals.

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controlled by the fraction of Si remaining in solution fdiss or in turn by the relative fraction removed

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This model implies a large range of precipitation fractions that also vary with depth (Fig. 6). Both catchments approach apparently congruent weathering condition in the soil zone during spring after snowmelt converging to zero net formation of secondary minerals. Water samples with fdiss > 1 even indicate preferential dissolution of clays. In course of the summer, the conditions turn into incongruent weathering with a high degree of Si uptake into secondary minerals. This opposing behavior is most pronounced in surface dominated water whereas water from periglacial layer and bedrock shows moderately incongruent weathering conditions. Overall, the derived conditions from water samples compare well with those indicated by clay-sized fractions, which give a time21

ACCEPTED MANUSCRIPT integrated weathering signal consistent with overall moderately incongruent weathering conditions (Fig. 6). In this model, phytoliths with isotope compositions similar to bedrock reveal almost quantitative Si uptake from solution by the vegetation, which is typical during the growing session (Farmer et al 2005). However, this observation is only partly consistent with the conditions inferred from water and likely indicate a different pool, for instance tightly bound soil water, which is a

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commonly accessed by trees (Brooks et al 2010).

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5.5. Physical erosion versus chemical weathering

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Taking this mass balance approach, we can use the isotope signatures of parental material and weathering products to estimate the relative proportions of the DSi flux WSi and erosional Si flux ESi

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out of the system. This relationship has been explored in detail by Bouchez et al. (2013). The isotope

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composition of the dissolved flux 30Sidiss is represented by stream water whereas the isotope

(8)

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composition of the particulate flux 30Sised is approach by the soil isotope composition.

We use the average isotope composition of soils above the C horizon for 30Sised (-0.33‰ for the sandstone catchment and -0.50‰ for the paragneiss catchment, respectively), the average isotope

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composition of the samples from the C horizon as approach for periglacial debris for 30Sirock (-0.32‰

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for the sandstone catchment and -0.44‰ for the paragneiss catchment, respectively) and average stream water compositions as 30Sidiss (0.57‰ for the sandstone catchment and 0.92‰ for the paragneiss catchment, respectively). This equation reveals that the Si export is predominately governed by the erosional flux with 98% for the sandstone catchment and 96% for paragneiss catchment whereas the remaining Si flux leaving as dissolved flux is very small with 2 and 4%, respectively. Taking maximal and minimal obtained soil isotope compositions, respectively (Table 3), we can estimate minimal and maximal values for the dissolved Si flux giving 0 to 17% for the sandstone catchment and 0 to 28% for the paragneiss catchment. 22

ACCEPTED MANUSCRIPT Because of the stochastic nature of particle transport in rivers, it is difficult to access ESi by sediment gauging. But, we can compare these estimates for the erosional flux with those calculated via the dissolved load WSi and the Si flux from bedrock into the weathering system, which equals the Si flux via denudation DSi out of the system by assuming steady state between regolith production and

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denudation. (9)

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The flux WSi can be calculated from the discharge q (1.34*109 L y-1 km-2 for the sandstone area and

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0.90*109 L y-1 km-2 for the paragneiss area, respectively) and the average Si concentration in stream water [Si]diss (74.8 µmol L-1 and 192.4 µmol L-1 for the sandstone and the paragneiss area,

(10)

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respectively).

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parageniss catchment, respectively.

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This reveals a flux WSi of 2.8 t km-2 yr-1 in the sandstone catchment and 4.9 t km-2 yr-1 in the

The flux DSi is a function of the Si content of bedrock [Si]rock (average values are 82 and 60 wt% SiO2 for sandstone and paragneiss, respectively) and the denudation rate D approached by using minimal

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and maximal denudation rates from cosmogenic nuclides, which are

29 and 142 t km-2 yr-1,

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respectively for sandstone and are 64 and 219 t km-2 yr-1, respectively for crystalline areas in the Black forest (these values are recalculated from data published by Meyer et al. (2010) and Morel et al. (2003) using an average rock density of 2.6 g cm-3 for sandstone and 2.7 g cm-3 for gneiss, respectively). (11) This approach reveals for the flux DSi 8.4 to 51.5 t km-2 yr-1 for the sandstone area and 13.4 to 57.0 t km-2 yr-1 for the paragneiss catchment.

23

ACCEPTED MANUSCRIPT Using these estimates on minimal and maximal denudation rates for sandstone and crystalline areas in the Black Forest, respectively, equation (9) gives similar ranges for the erosional Si flux ESi for both catchments, which is 75 to 95% on sandstone and 73 to 92% on paragneiss whereas the remainder is dissolved Si flux WSi (5 to 25% on sandstone and 8 to 27% on paragneiss, respectively). Another method to estimate the dissolved export flux are chemical depletion indices. Calculated CDF

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values reveal that chemical weathering accounts for about 10 to 20% of the total denudation. τZr, Si

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values indicate the same development but explicitly for Si with 15 to 20% average loss of Si from

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soils.

All this approaches disclose a similar range for the dissolved Si flux, which compares well with the

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estimates from the isotope mass balance (equation 8) (see summary in Table 4). Therefore, such a

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system with similar isotope signature of soils and parental material but distinct signatures for DSi, indicate a high erosional flux relative to the dissolved flux. Part of the erosional flux is the export of biogenic opal, i.e. phytoliths. The vegetation represents a dynamic pool acting as sink and source for

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Si but is responsible for a net transfer of Si from the DSi pool into the solid pool (e.g. Cornelis and Delvaux, 2016). In general, plants take up DSi from soil solution forming phytoliths, which are recycled to the soil zone where they are either dissolved or removed by erosion. Taking the Si uptake

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rate determined for Norway spruce on granite under temperate climatic conditions of 4.35 t km-2 y-1 (Cornelis et al. 2010) and assuming that 40% of recycled phyoliths are preserved in soils (Sommer et

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al. 2013) we can estimate the erosional Si flux via phyoliths, which is 1.7 t km-2 y-1. This estimate reveals that phyoliths represents a substantial export flux contributing 3 to 21% of the total Si flux in sandstone catchment and 3 to 13% in the paragneiss catchment, respectively. Thus, phytolith erosion contributes in the same order as the dissolved export flux to the desilification of the system. Several studies on various small catchment underlying by a range of lithologies including shale, sandstone and granite in temperate climatic conditions have disclosed comparable results for various isotope systems including Li, B, Mg and Si: namely soils with very similar isotope signatures as 24

ACCEPTED MANUSCRIPT parental rock but distinct isotope signatures for dissolved phases (Bolou-Bi et al. 2012; Lemarchand et al. 2010; Ma et al. 2015; Noireaux et al. 2014). Based on the isotope mass balance approach above (equation 8), the considered elements have to be predominately removed from the system by the erosional flux whereas the dissolved flux is rather small, which is typically for kinetic-limited systems. In turn high erosion rates triggered by high denudation rates leads to little fractionated soil but highly

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fractionated water under incongruent weathering conditions due to short regolith residence times.

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Conclusion

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This study shows that Si isotopes can decipher the complex release of Si in the weathering zone and underlying bedrock. Independently of bedrock lithology, soils carry a similar isotope composition as

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their parental material whereas DSi in soil water and groundwater is highly fractionated that also

-

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shows large spatial and temporal variations in δ30Si.

At both filed sites, chemical weathering is dominated by mineral transformation processes

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leaving the original silicate lattice largely unmodified. Therefore, neoformed clays have

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largely inherited the isotope signature from parental minerals. i.e. illite in the sandstone catchment and biotite in the parageniss catchment, respectively. Organic matter degradation together with a high water flux in organic horizons enhance

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-

dissolution processes accompanied with increased precipitation of isotopically light clays, i.e.

-

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kaolinite which leads to strong depletion of isotopically heavy Si in these horizons. Si fluxes and isotope ratios of DSi in the soil zone are largely controlled by seasonal effects: 1) Enhanced release of DOC in spring after snowmelts triggers low pH values, which favors dissolution of isotopically light phases such clays and amorphous Si leading to DSi with low 30Si values, 2) In the course of the summer, secondary mineral formation and Si uptake by plants drives DSi to typical positive 30Si values. -

Groundwater records deep weathering processes and contributes substantially to the discharge of the streams. In the sandstone catchment, weathering in periglacial debris layers 25

ACCEPTED MANUSCRIPT and underlying bedrock is governed by plagioclase weathering and the precipitation of kaolinite, leading to an enrichment of isotopically heavy Si in DSi compared to bedrock. -

A mass balance approach based on isotope ratios reveals that Si export results predominately by erosional flux whereas the dissolved flux is small for both investigated settings. Therefore, weathering systems with soils carrying a similar isotope signature as

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their parental material but showing highly fractionated water sources are dominated by physical erosion whereas chemical weathering contributes little to observed denudation

Plant activity causes a net transfer of Si from the DSi pool to the solid Si pool by forming

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-

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rates, which is typical for kinetic-limited systems.

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biogenic opal, i.e. phytoliths. Estimates show that erosion of phytoliths represents a substantially export flux contributing up to 21% of the total Si flux. This amount is

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comparable with the flux leaving the system as DSi.

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Acknowledgments

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We thank K. P. Jochum for organizing the preparation of fused glass beads and S. Fiedler and M. Hinderer for their help in field work. M. Oelze and M. Tatzel provided support for Si isotope analysis.

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The manuscript has benefited from discussions with M. Oelze, S. Hynek, A. Dere, S. Fiedler and M. Hinderer. We appreciate the constructive comments by D. Cardinal and an anonymous reviewer and

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the editorial handling by J. Gaillardet. This study was supported by the Deutsche Forschungsgemeinschaft (DFG) – PAK 179 “Multiscale analysis of Si cycling in terrestrial biogeosystem” (SO 302/3-1). GS was supported by a fellowship provided by the Deutsches GeoForschungsZentrum GFZ.

26

ACCEPTED MANUSCRIPT Figure Captions Fig. 1 Investigated field sites in the Black Forest (Germany): A) sandstone catchment and B) hillslope developed on paragneiss. The Si isotope composition 30Si in ‰ of water samples are shown in blue for spring and in red for summer, respectively. The green dots denote locations of the investigated soil profiles in the sandstone catchment (Sc1, Sc2 and Sc3) and on the hillslope developed on

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paragneiss (Pc1, Pc2 and Pc3), respectively.

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Fig. 2 Si isotope composition for water samples in the sandstone catchment (A) and the paragneiss

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hillslope (B), shown against the source of water. Filled symbols: spring; open symbols: summer.

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Fig. 3 Si isotope data of the principle solid Si pools in the investigated soil profiles. The data of Cambisols, Sc3 and Pc1, are published in Steinhoefel et al. (2011). Organic horizons are coloured in

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dark brown whereas mineral soil horizons are depicted in light brown. The reference of the depth information is the boundary been the organic horizons and the mineral soil horizons. Plotted isotope

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data represent each the average value of 2 to 5 measurements. The depicted uncertainty bar refers

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to the long-term reproducibility, detailed information about the uncertainty for each data point is given in Table 3.

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Fig. 4 Si isotope composition of water versus pH and DOC, Al, Si concentrations in the sandstone catchment (red circles) and the paragneiss catchment (blue diamonds). Filled symbols: spring; open

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symbols: summer.

Fig. 5 Calculated saturation indices (SI) for relevant mineral phases in representative water sources. (A) Sandstone catchment: Sb3 (mixture near-surface/subsurface water), Sb1 (subsurface water) and Sb8 (base flow) and (B) hillslope on paragneiss: Wm-G (surface water) Wm-Q and Wm-OQ (both mixture for near-surface/subsurface water). Filled symbols: spring; open symbols: summer.

27

ACCEPTED MANUSCRIPT Fig. 6 Isotope mass balance model showing the exchange of Si between secondary phases and water in a batch reactor model for the sandstone catchment (A) and the hillslope on paragneiss (B). Filled symbols: spring; open symbols: summer. The DSi fraction fdiss is calculated using equation (7) for the water sample. Likewise, fdiss can be calculated for clay fractions and phytolith samples

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). All displayed 30Si values are measured data.

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Alleman LY, Cardinal D, Cocquyt C, Plisnier PD, Descy JP, Kimirei I, Sinyinza D, Andre L (2005) Silicon

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isotopic fractionation in Lake Tanganyika and its main tributaries. J Great Lakes Res 31: 509-

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519.

Amiotte-Suchet P, Probst J-L, Ludwig W (2003) Worldwide distribution of continental rock lithology:

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Implications for the atmospheric/soil CO2 uptake by continental weathering and alkalinity river transport to the oceans. Glob Biogeochem Cycles 17: 1038, doi:10.1029/2002GB001891

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Andermann CL, Longuevergne L, Bonnet S, Crave A, Davy P, Gloaguen R (2012) Impact of transient groundwater storage on the discharge of Himalayan rivers. Nat Geosci 5: 127-132.

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Bern CR, Brzezinski MA, Beucher C, Ziegler K, Chadwick OA (2010). Weathering, dust, and biocycling

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effects on soil silicon isotope ratios. Geochim Cosmochim Acta, 74: 876-889. Bolou-Bi EB, Vigier N, Poszw A, Boudot J-P, Dambrine E (2012) Effects of biogeochemical processes

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ACCEPTED MANUSCRIPT

Table 1 Location and origin of water samples Sample name

Elevation AMSL [m]

UTM coordinates

spring / late summer

Source of water samples

Sandstone catchment

32U 447112 5390877

919

periglacial debris layers

Sb2-s / dried out

32U 447381 5390672

962

Sb3-s / Sb3-a

32U 447496 5391372

827

soil zone / periglacial debris layers

Sb5-s / Sb5-a

32U 448034 5391417

771

soil zone / periglacial debris layers

Sb7-s / dried out

32U 448234 5391525

Sb8-s / Sb8-a

32U 448831 5391189

Sb13-s / Sb13-a

32U 449067 5390458

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PT

Sb1-s / Sb1-a

MA

NU

SC

soil zone

D

774

soil zone

sandstone

702

main stream

879

sandstone

32U 441414 5381664

674

granite

Wm-M-s / Wm-M-a

32T 433827 5312297

1085

soil zone

Wm-G-s / Wm-G-a

32T 433981 5312318

1092

soil zone

Wm-Q-s / Wm-Q-a

32T 433644 5312298

1068

soil zone / periglacial debris layers

Wm-OQ-s / Wm-OQ-a

32T 433829 5313045

1062

soil zone / periglacial debris layers

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32U 449579 5391106

AC

Sb16-s / Sb16-a

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746

SB-WB-s / SB-WB-a

spring, subsurface water passing lower perglacial layers and influenced by base flow from Bunter Sandstone episodic spring, nearsurface water passing soil zone and upper periglacial layers, reflecting soil water of Dystric Albic Stagnosol mixture of Sb1 and Sb2 run off of Herrenwieser See, small lake fed by near near-surface and subsurface water similar to Sb3 episodic little creek, near-surface water passing soil zone and upper periglacial layers, reflecting soil water of a catchment dominated by Podzols spring, deep base flow at the contact lower Bunter Sandstone/Forbach granite Seebach, main stream fed by 50% base flow and 50% near-surface and subsurface flow water (annual average) spring, base flow from fissured aquifer Mittlere Buntsandstein smb with many clay-rich layers and subsurface flow passing periglacial layers Spring Wolfsbrunnen, fissured aquifer Forbach granite

Paragneiss catchment

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small high raised bog, near-surface water trench, near-surface water passing soil zone spring, near-surface and subsurface water passing soil zone and perglacial layers, drain the catchment excluding Chromic Stagnic Cambisols spring rich in Fe-oxide precipitation, nearsurface and subsurface water passing soil zone

ACCEPTED MANUSCRIPT and periglacial layers

Wm-OQC-s / Wm-OQC-a 32T 433845 5312879

1053

small stream 150 m downstream from WmOQ, inflow of nearsurface water

soil zone / periglacial debris layers

Table 2

29Si

Sample materials

2SD

30Si

0.06

Diatomite

0.63

0.06

BHVO-2

-0.16

0.07

Big Batch

-5.45

0.06

Sb1-a Sb3-a Sb5-a Sb8-a

9

3

0.10

3

0.04

0.07

0.06

3

0.04

0.06

0.10

0.14

3

-0.34

0.02

-0.66

0.07

3

0.25

0.05

0.47

0.12

3

0.24

0.06

0.42

0.10

3

0.14

0.05

0.29

0.02

3

0.17

0.03

0.34

0.11

3

0.18

0.03

0.29

0.09

3

0.42

0.04

0.83

0.10

3

0.13

0.04

0.20

0.12

3

0.26

0.08

0.47

0.13

3

D

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Sb-WB-s Late summer

0.11

0.06

0.03

Sb16-s

11

-10.65

0.30

Sb3-s

Sb13-s

15

0.10

-0.42

-0.22

Sb8-s

0.12

0.04

0.14

Sb2-s

Sb7-s

1.27 -0.29

0.05

Sb1-s

Sb5-s

20

MA

Spring

0.11

NU

Sandstone catchment (Seebach)

n

-1.38

SC

-0.71

2SD

RI

Standard materialsa IRMM-17

PT

Si isotope composition of standard materials and water samples

0.39

0.03

0.72

0.20

3

Sb16-a

0.28

0.07

0.52

0.19

3

Sb-WB-a

0.25

0.05

0.49

0.06

3

CE

Sb13-a

AC

Paragneiss catchment (Wildmooswald) Spring Wm-M-s

0.01

0.10

-0.01

0.07

3

Wm-G-s

0.28

0.05

0.51

0.11

3

Wm-Q-s

0.38

0.03

0.74

0.04

3

Wm-OQ-s

0.46

0.03

0.87

0.02

3

Wm-OQC-s Late summer

0.27

0.04

0.56

0.07

3

Wm-M-a

0.47

0.07

0.89

0.15

3

Wm-G-a

0.40

0.12

0.80

0.15

2

Wm-Q-a

0.55

0.04

1.08

0.07

3

Wm-OQ-a

0.51

0.04

1.00

0.07

10b

Wm-OQC-a

0.23

0.05

0.51

0.04

3

40

ACCEPTED MANUSCRIPT a

Redigested standard materials.b10 aliquotes of sample WM-OQ-a were decomposed and measured each at least three times. Data of reference material refers to replicated measurements on the mass spectrometer. Estimates of the population's SD are calculated as followed:

is the

AC

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D

MA

NU

SC

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where n is the number of samples, x is the observed value and mean value of observations.

41

ACCEPTED MANUSCRIPT

Table 3 Si isotope composition of bulk soil samples and its compartments Sample name

Mineral phase

Method

29Si

2SD

30Si

2SD

n

Sandstone catchment Parent minerals (rock fragments) Quartz

thin section

-0.15

0.14

-0.35

0.27

18

Feldspar

thin section

0.30

19

powder pellet

0.24 0.18

-0.49

Illite*

-0.26 -0.39

-0.68

0.18

5

1202

fused glass

-0.28

0.07

-0.50

0.05

1205

fused glass

-0.14

0.17

-0.29

0.18

1208

fused glass

-0.12

0.12

-0.27

0.11

1210

fused glass

-0.11

0.10

-0.31

0.09

Clay fraction 0.2 - 2 µm powder pellet 1205

-0.28

0.09

-0.66

1206

powder pellet

-0.25

0.20

-0.50

1210

powder pellet

-0.42

0.15

-0.95

1205

powder pellet

-0.46

0.05

1206

powder pellet

-0.45

0.11

1210

powder pellet

-0.61

2201

powder pellet

2202

fused glass

-0.17 -0.12

2203

fused glass

-0.17

2205

fused glass

-0.15

2206

fused glass

2207

fused glass

4

SC

RI

4 4 5

3

0.31

3

0.25

5

-0.92

0.11

3

-0.84

0.14

3

MA

NU

0.35

0.03

-1.04

0.11

3

0.16

0.33

3

0.06

-0.40 -0.24

0.12

3

0.19

-0.32

0.11

2

0.03

-0.32

0.06

4

-0.16

0.08

-0.26

0.18

3

-0.15

0.05

-0.32

0.07

3

powder pellet

-0.62

0.05

-1.04

0.11

3

CE

Clay fraction < 0.2 µm

PT

Sc1 (E-Podzol) Bulk soil (< 2 mm)

-0.38

0.18

-0.57

0.25

4

-0.46

0.19

-0.74

0.33

4

PT E

D

Sc2 (Bs-Podzol) Bulk soil (< 2 mm)

Clay fraction < 2 µm 2205

Clay fraction 0.2 - 2 µm powder pellet 2207 Clay fraction < 0.2 µm

powder pellet

AC

2207

42

ACCEPTED MANUSCRIPT

Table 3 continued Sample name

Mineral phase

Method

29Si

2SD

30Si

2SD

n

Paragneiss catchment Quartz

thin section

-0.12

0.19

-0.22

0.40

18

Feldspar

thin section

-0.24

0.23

-0.48

0.32

12

Biotite

thin section

-0.36

-0.36

-0.69

0.37

16

5201

powder pellet

-0.49

0.31

-0.84

0.37

5

5202

powder pellet

-0.59

0.02

-0.97

0.18

3

5203

powder pellet fused glass

0.23 0.18

-0.74 -0.31

0.38 0.26

4

5205

-0.46 -0.16

5206

fused glass

-0.22

0.16

-0.40

0.28

5207

fused glass

-0.11

0.22

-0.19

5208

fused glass

-0.27

0.09

-0.48

Clay fraction 0.2 - 2 µm powder pellet 5205

-0.53

0.06

-0.81

Clay fraction < 0.2 µm powder pellet 5205

-0.25

0.16

-0.33

5

SC

11

0.34

5

0.21

5

0.14

3

0.33

4

NU

MA

Pc3 (Cambisol) Bulk soil (< 2 mm)

RI

Pc2 (Stagnosol) Bulk soil (< 2 mm)

powder pellet

-0.45

0.01

-0.75

0.13

2

7203

powder pellet

-0.29

0.14

-0.48

0.26

3

7204

fused glass

-0.07

0.09

-0.20

0.08

4

7205

fused glass

4

fused glass

-0.27 -0.32

0.11

7206

-0.09 -0.14

0.20

6

7208

fused glass

-0.25

0.08

-0.40

0.04

5

powder pellet

-0.43

0.15

-0.76

0.25

3

0.13

PT E

Clay fraction < 2 µm

CE

7206

D

7202

0.03

AC

*separated from a clay sample originated from a sandstone outcrop. All data on solid samples were obtained by LA-MC-ICP-MS. Estimates of the population's SD are calculated as followed:

where n is the number of samples, x is the observed value and value of observations.

43

is the mean

PT

Parental minerals (rock fragments)

ACCEPTED MANUSCRIPT

Table 4 Estimates of the relative proportion of Si export fluxes contributing to the total denudation (see text for details)

Physical Erosion ESi (%)

Chemical weathering WSi (%)

Isotope mass balance

Estimate from denudation and dissolved fluxes

CDFZr

Zr, Si

SC

PC

SC

PC

Both catchments

Both catchments

83 - 100

82 - 100

75 - 95

73 - 92

80 - 90

80 - 85

(98a)

(96a)

0 - 17

0 - 18

5 - 25

8 - 27

10 - 20

15 - 20

(2a)

(4a)

PT

Method:

AC

CE

PT E

D

MA

NU

SC

RI

SC = sandstone catchment; PC = paragneiss catchment. aThese values represent the best estimates from the isotope mass balance (see text for details).

44

NU

SC

RI

PT

ACCEPTED MANUSCRIPT

AC

CE

PT E

D

MA

Figure 1

45

MA

NU

SC

RI

PT

ACCEPTED MANUSCRIPT

AC

CE

PT E

D

Figure 2

46

AC

CE

PT E

D

MA

NU

SC

RI

PT

ACCEPTED MANUSCRIPT

Figure 3

47

PT E

D

MA

NU

SC

RI

PT

ACCEPTED MANUSCRIPT

AC

CE

Figure 4

48

PT E

D

MA

NU

SC

RI

PT

ACCEPTED MANUSCRIPT

AC

CE

Figure 5

49

AC

Figure 6

CE

PT E

D

MA

NU

SC

RI

PT

ACCEPTED MANUSCRIPT

50