Geoderma 235–236 (2014) 19–29
Contents lists available at ScienceDirect
Geoderma
journal homepage: www.elsevier.com/locate/geoderma
Silicon isotopes record dissolution and re-precipitation of pedogenic clay
minerals in a podzolic soil chronosequence
Jean-Thomas Cornelis a,b,⁎, Dominique Weis a, Les Lavkulich c, Marie-Liesse Vermeire b,
Bruno Delvaux b, Jane Barling a,1
a
Pacific Centre for Isotopic and Geochemical Research, Department of Earth, Ocean and Atmospheric Sciences, University of British Columbia (UBC), 6339 Stores Road, Vancouver,
BC V6T 1Z4, Canada
b
Soil Science and Environment Geochemistry, Earth and Life Institute, Université catholique de Louvain, Croix du Sud 2/L7.05.10, B-1348 Louvain-la-Neuve, Belgium
c
Soil Science, University of British Columbia (UBC), 127-2357 Main Mall, Vancouver, BC V6T 1Z4, Canada
a r t i c l e
i n f o
Article history:
Received 8 February 2014
Received in revised form 20 June 2014
Accepted 22 June 2014
Available online xxxx
Keywords:
Podzol
Silicon isotopes
Soil formation
Clay minerals
Biogeochemical cycles
a b s t r a c t
By providing the largest part of the reactive surface area of soils, secondary minerals play a major role in terrestrial biogeochemical processes. The understanding of the mechanisms governing neo(trans-)formation of pedogenic clay minerals in soils is therefore of the utmost importance to learn how soils evolve and impact the
chemistry of elements in terrestrial environments. Soil-forming processes governing the evolution of secondary
aluminosilicates in Podzols are however still not fully understood. The evolution of silicon (Si) isotope signature
in the clay fraction of a podzolic soil chronosequence can provide new insight into these processes, enabling to
trace the source of Si in secondary aluminosilicates during podzol-forming processes characterized by the mobilization, transport and precipitation of carbon, metals and Si. The Si isotope compositions in the clay fraction
(comprised of primary and secondary minerals) document an increasing light 28Si enrichment and depletion
with soil age, respectively in illuvial B horizons and eluvial E horizon. The mass balance approach demonstrates
that secondary minerals in the topsoil eluvial E horizons are isotopically heavier with δ30Si values increasing
from − 0.39 to + 0.64‰ in c.a. 200 years, while secondary minerals in the illuvial Bhs horizon are
isotopically lighter (δ30Si = −2.31‰), compared to the original “unweathered” secondary minerals in BC horizon (δ30Si = −1.40‰). The evolution of Si isotope signatures is explained by the dissolution of pedogenic clay
minerals in the topsoil, which is a source of light 28Si for the re-precipitation of new clay minerals in the subsoil.
This provides consistent evidence that in strong weathering environment such as encountered in Podzols, Si released from secondary minerals is partially used to form “tertiary clay minerals” over very short time scales
(ca. 300 years). Our dataset demonstrates the usefulness to measure Si isotope signatures in the clay fraction
to discern clay mineral changes (e.g., neoformation versus solid state transformation) during soil evolution.
This offers new opportunity to better understand clay mineral genesis under environmental changes, and the
short-term impact of the dissolution and re-precipitation of pedogenic clay minerals on soil fertility, soil carbon
budget and elemental cycles in soil–plant systems.
© 2014 Elsevier B.V. All rights reserved.
1. Introduction
Soil is a precious but threatened resource (Banwart, 2011). In order
to protect it for the future we need a better understanding of the soilforming processes controlling the evolution of newly-formed minerals
(secondary minerals). Soil formation progressively modifies parent
rock material and controls the pathways of primary mineral weathering
and secondary mineral synthesis in the clay fraction (Chadwick and
⁎ Corresponding author at: Earth and Life Institute (ELI-e), Université catholique de
Louvain (UCL), Croix du Sud 2, L7.05.10, 1348 Louvain-la-Neuve, Belgium.
E-mail address: [email protected] (J.-T. Cornelis).
1
Now at Department of Earth Sciences, University of Oxford, South Parks Road, Oxford
OX1 3AN, United Kingdom.
http://dx.doi.org/10.1016/j.geoderma.2014.06.023
0016-7061/© 2014 Elsevier B.V. All rights reserved.
Chorover, 2001). The secondary minerals consist of layer-type
aluminosilicates (called pedogenic clay minerals) and Fe-, and Aloxyhydroxides, both of which play a major role not only in soil fertility,
but also in the transfer of elements and pollutants from land to ocean
given their high surface reactivity (Sposito, 2008). Moreover, the
capacity of charged mineral surfaces to form adsorption complexes
can stabilize organic carbon (OC) in soils through the formation of
organo-mineral associations, partly controlling global C budget (Parfitt
et al., 1997; Torn et al., 1997).
The formation of secondary minerals and their evolution during pedogenesis have been studied for over a half century (Wilson, 1999). The
proportion and the chemistry of minerals in the clay fraction change
with soil evolution (Egli et al., 2002; Righi et al., 1999; Turpault et al.,
2008). Some environmental changes (vegetation type, agricultural
20
J.-T. Cornelis et al. / Geoderma 235–236 (2014) 19–29
practices, land-use, climate and drainage) can amplify the modification
of clay mineralogy on very short time-scales (10–1000 years) (Caner
et al., 2010; Collignon et al., 2012; Cornu et al., 2012; Mareschal et al.,
2013). These rapid clay modifications occur in chemically reactive soil
micro-environments, i.e. the part of the soil influenced by roots and
earthworms (Calvaruso et al., 2009; Jouquet et al., 2007), and can play
a key role in geochemical balance of several minor and major elements
in soils and sediments (Michalopoulos and Aller, 1995; Velde and
Meunier, 2008). However, the origin of elements involved in clay
neo(trans-)formation is still not well understood.
Podzol, the focus of this study, is a type of soil that covers more than
3% of the Earth's land surface. The low stock of weatherable minerals,
the acidic conditions and complexing capacity of organic acids in the environment where Podzols developed are responsible for mobilization,
transport and precipitation of carbon (C), metals (Fe, Al) and silicon
(Si) in the soil profile (Lundström et al., 2000). A fully developed Podzol
consists of a leached gray subsurface eluvial E horizon contrasting with
the accumulation of elements in the dark illuvial B horizons. The topsoil
is characterized by the production of organic acids that form soluble
organo-metallic complexes enhancing weathering in the eluvial E horizon. This E horizon overlies the dark C-enriched Bh horizon and reddish
Fe-, Si-, and Al-enriched Bhs/Bs horizons (Lundström et al., 2000). Given
the very acidic conditions in Podzols, besides the weathering of primary
minerals, secondary clay minerals can be dissolved in the podzolic
weathering front (Ugolini and Dahlgren, 1987; Zabowski and Ugolini,
1992), which describes the soil depth where minerals dissolve faster
than they form. A podzolic soil chronosequence, i.e. in which all soilforming factors remain constant except time; represents an ideal natural system for the study of the effect of time on pedogenic clay minerals
behavior in soils.
Stable Si isotopes fractionate during silicate weathering and the biogeochemical Si cycling (Opfergelt et al., 2010; Ziegler et al., 2005), and
as such provide a means of tracing the bio-physico-chemical processes
in terrestrial environments (Cornelis et al., 2011). In addition to its incorporation in the mineral structure during the formation of crystalline
layer-type aluminosilicates, poorly-crystalline aluminosilicates and
pedogenic opal, monosilicic acid (H4SiO4) released into soil solution
can also be transferred into the biosphere to produce biogenic opal
(phytoliths) or be adsorbed onto secondary Fe oxy-hydroxides. The incorporation of Si in mineral structures through neoformation of secondary pedogenic and biogenic precipitates and its adsorption onto the
surfaces of Fe oxides are two processes favoring the retention of light
28
Si in soils and contributing to the enrichment of rivers in heavy 30Si
(Delstanche et al., 2009; Georg et al., 2007; Opfergelt et al., 2006;
Ziegler et al., 2005). Clay minerals can also be unstable in organic and inorganic acidic environments where they dissolve (Sokolova, 2013;
Zabowski and Ugolini, 1992), and enrich soil solutions (Cornelis et al.,
2010) and rivers (Cardinal et al., 2010) in light 28Si. The naturally occurring mass-dependent Si isotopic fractionation is induced by dissolution,
precipitation and adsorption but not by complexation as chemical binding of Si to organic matter is negligible (Pokrovski and Schott, 1998). It
has also been demonstrated that the Si isotopic compositions of secondary clay minerals relates to climatic gradient and its control on clay mineralogy (Opfergelt et al., 2012). However Si isotopes have never been
used to better understand clay mineral modifications induced by soilforming processes under identical geo-climatic conditions. The rapid
modification of clay mineralogy in Podzol is well documented (Caner
et al., 2010; Egli et al., 2002; Righi et al., 1999), but the fate of Si released
in soil solution after clay modification has not yet been studied, even
though it is of crucial importance for identifying the sources controlling
the formation of pedogenic clay minerals in soils.
In this study, we aim to use Si isotope signatures of the clay fraction
in a podzolic soil chronosequence for gaining better insights into the
origin of Si in pedogenic clay minerals.
To achieve this goal, we analyzed Si isotopes, elemental (Ge/Si, Al/Si,
Fe/Si) ratios and determined clay fraction mineralogy for an age
sequence of four soil profiles undergoing podzolization (Cox Bay on
Vancouver Island, Canada) (Fig. 1) and for a single Podzol pedon
(Gaume, Belgium). The Cox Bay chronosequence offers an opportunity
to study the variation of Si isotopic composition and elemental ratios
of the clay fraction in the vertical pedogenic scale: E, Bh, Bhs, Bs, Bw
and BC horizons, and in the horizontal time-dependent scale: duration
of pedogenesis from 0 to 335 years. We used the Belgian Podzol as a
“natural duplicate” in temperate climate to corroborate the processes
documented in the soil samples from the Cox Bay podzolic soil
chronosequence.
2. Materials and methods
2.1. Sample collection and location
We sampled a soil chronosequence undergoing podzolization in Cox
Bay (CB), on the west coast of Vancouver Island (British Columbia,
Canada). At the Cox Bay study site, three main vegetative associations
are identified in the chronosequence. These correspond to Sitka spruce
(Picea sitchensis) in the younger site (CB-120 years), and Sitka spruce
(P. sitchensis) and salal (Gaultheria shallon) in the sites of 175 and
270 years (CB-175 and -270 years). The oldest site (CB-335 years) is
characterized by Sitka spruce (P. sitchensis), Douglas fir (Pseudotsuga
menziesii), salal (G. shallon) and western sword fern (Polystichum
munitum). Heavy mean annual precipitation (3200 mm) coupled with
frequent fogs and sea sprays ensure an abundance of moisture and nutrients year round in this maritime temperate climate (Cfb: without dry
season and with warm summer; Peel et al., 2007). The Tofino Area
Greywacke Unit is the source of the beach sand parent material, from
which soils have developed in the age sequence (Singleton and
Lavkulich, 1987). Sampling sites were located along a transect (0–94
m) perpendicular to the present shoreline (Fig. 1). Dendrochronology
and geomorphology established surface duration of pedogenesis ranging from 0 to 335 years for the four selected pedons. Tree ages were determined counting the tree rings in the increment bores. Assuming that
the beach built towards the ocean in a configuration parallel to the
existing shoreline and that a linear deposition rate occurred with time
between successive oldest trees, the rate of advance of the beach front
was estimated to be 0.26 m per year. At this rate, the 13-m strip of
sand containing tree seedlings would have accumulated in approximately 50 years (Singleton and Lavkulich, 1987). With soil development, there was progressive deepening and differentiation of genetic
horizons during podzolization, resulting in soil classification (World
Reference Base for Soil Resources — WRB) that ranged from Dystric
Cambisol at the youngest sites (CB-120 years; CB-175 years) to a Placic
Podzol at the oldest site (CB-335 years) (Fig. 1). The 335-year-old Podzol is characterized by the following soil horizon development: eluvial
albic E horizon (strongly weathered horizon) → illuvial spodic Bh horizon (enriched in organic matter) → Bhs horizon (enriched in Fe
oxyhydroxides and organic matter) → Bs horizon (enriched in poorlycrystalline aluminosilicates and Fe oxyhydroxides) → Bw horizon
(development of color and structure without illuvial accumulation of
materials) → BC horizon (weakly colored and structured; little affected
by pedogenic processes).
The sampling area of the Podzol in Gaume (Belgium), ranging in altitude from 300 to 350 m above sea level, has an annual rainfall of
1100 mm and a mean annual temperature of 7.7 °C (Herbauts, 1982),
and is also characterized by a maritime temperate climate (Cfb; Peel
et al., 2007). The Podzol is located on the Lower Lias outcrop in Southeast Belgium (Gaume). The bedrock (calcareous sandstone of Lower
Lias age) is covered by a two-layered sheet: an autochthonous sandy
layer, formed by the dissolution of the calcareous bedrock, is overlaid
by a mixture of this sandy material with loessic silt-sized particles.
The Belgian Podzol developed under heather (Calluna vulgaris) is characterized by a similar morphological profile as the Podzol in Cox Bay
J.-T. Cornelis et al. / Geoderma 235–236 (2014) 19–29
21
Vertical scale (m)
50
40
30
20
0,26 m/year
10
Beach
0
0
10
20
CB-0 yr CB-120 yrs
30
40
50
60
CB-175 yrs
CB-270 yrs
E
E
Bh
Bw
C
BC
Bw
BC
Parent
material
Dystric Cambisol
70
80
90
100
110
120
130
140
150
Horizontal scale (m)
CB-335 yrs
E
Bh
Bhs
Bs
0
Bw
BC
BC
Haplic Podzol
Placic Podzol
-0.75 m
Fig. 1. Cross section of the Cox Bay study area showing site locations and soil horizons, depending on their respective age of soil formation: CB-0 year, CB-120 years, CB-175 years,
CB-270 years and CB-335 years.
sequence (CB-335 years) with the following horizons: E–Bh–Bhs–Bs–
Bw–BC.
2.2. Physico-chemical characterizations
The soil samples were air-dried, then sieved and homogenized. The
content of free iron oxides was assessed after selective dissolution of
Fe oxides using Na-dithionite–citrate–bicarbonate and ammonium oxalate–oxalic acid (Fedcb = crystalline Fe oxides, Feox = poorly-crystalline
Fe oxides). The content of Si bound to poorly crystalline aluminosilicates
and weakly-ordered Fe oxyhydroxides was estimated on fine earth by
extraction with ammonium oxalate–oxalic acid (Siox). Al complexed
with organic ligands was assessed using the complexing agent Napyrophosphate at pH 10 (Alp). The total organic carbon (OCtot) content
was measured on ground samples using CNS analyzer.
The clay fraction (b2 μm) was separated using a ‘clean procedure’
without any oxidative treatment. Air-dried soil was dispersed in deionized water and sonicated. The suspension was then separated on a
50 μm sieve, re-suspended in deionized water and sonicated and sieved
until the supernatant was clear after sonication. The fraction retained in
the sieve was collected as the N50 μm sand fraction. Clay (0–2 μm) and
silt (2–50 μm) fractions were then collected by gravimetric sedimentation after dispersion using an ultrasonic probe and Na+ as a dispersion
agent.
2.3. X-ray diffraction patterns
XRD analyses were carried out on the clay-sized fraction (b2 μm) of
soil horizons sampled in the Cox Bay chronosequence (120, 175, 270
and 335 years), using CuKα radiation in a Bruker Advance diffractometer. After removal of the organic matter by treating the sample with 6%
H2O2 at 50 °C, and removal of Fe-oxyhydroxides using dithionite–
citrate–bicarbonate, eight standard treatments were applied to determine mineralogy of the clay fraction: K-saturation (KCl 1 N) followed
by drying and heating at 20, 105, 300 and 550 °C, and Mg-saturation
(MgCl2 1 N) followed by drying at 20 °C and saturation with ethyleneglycol (eg). XRD analysis was also performed on powder samples of
the clay-sized fraction after removal of organic matter and Fe
oxyhydroxides but without any further treatment for quantifying mineralogy of the clay fraction using the Siroquant software V4.0 (Sietronics
Pty Ltd), and for the following horizons: BC horizon (CB-120 years), E
horizons (CB-175, 270 and 335 years), Bh horizon (CB-335 years) and
Bhs horizon (CB-335 years).
2.4. Isotopic and geochemical analyses
Silicon isotope compositions and elemental (Ge, Al, Fe and Si) concentrations were measured on clay-sized fraction (b2 μm) extracted
from all the horizons of the four soil profiles in Cox Bay (clay-CB
120 years; clay-CB 175 years; clay-CB 270 years; clay-CB 335 years)
and the undated podzolic soil profile in Gaume (clay-G), and also on
parent material of soils in Cox Bay (sand fraction of the beach sand;
Beach-CB 0 year). An alkaline digestion with 99.99% pure NaOH is
used to transform solid samples into an aqueous HF-free solution
(Georg et al., 2006). All dissolutions and chemical separations were carried out in Class 100 laminar flow hoods in Class 1000 clean labs, mass
spectrometric analyses were performed in Class 10,000 laboratories at
the Pacific Centre for Isotopic and Geochemical Research (PCIGR) at
the University of British Columbia (UBC). Al, Fe and Si contents of the
dissolved NaOH fusions were analyzed by ICP-OES (Varian 725-ES)
with Europium as the internal standard. For Ge measurements, the dissolved NaOH fusions were dried and re-dissolved in 1% v/v HNO3 with
10 ppb indium (In) for analysis by HR-ICP-MS (Element 2) in medium
resolution.
22
J.-T. Cornelis et al. / Geoderma 235–236 (2014) 19–29
The remaining dissolved NaOH fusion solution was purified for isotopic analyses through cation exchange chromatography (Georg et al.,
2006). The Si isotope compositions were measured on a Nu Plasma
(Nu 021; Nu Instruments Ltd, UK) MC-ICP-MS in dry plasma mode
using type B cones and a Cetac Aridus II desolvating nebulizer system.
Instrumental mass bias was corrected by simple sample-standard
bracketing of measured Si isotope ratios, i.e. one sample measurement
normalized to the average of two bracketing NBS-28 standard measurements. Silicon isotopic compositions are expressed as deviations in 30Si/
28
Si relative to the NBS-28 reference standard using the delta (δ) per mil
(‰) notation: δ30Si = [(30Si / 28Sisample) / (30Si / 28SiNBS28) − 1] × 1000.
Each sample was measured at least twice during different analytical sessions. Silicon isotopic (δ30Si) values are reported as the mean of
replicate isotopic analyses (n N 2) ± 2 standard deviations (SD). The
NBS-28 (quartz standard) which processed through the full analytical
procedure, and analyzed over a period of 7 months during 5 data acquisition sessions gave a value of δ30Si = 0.01 ± 0.18‰ (2SD, n = 66). Accuracy and reproducibility were also checked on reference materials
(diatomite and BHVO-2) at the beginning and at the end of each sample
series. These gave values identical within error to previously published
values: 1.24 ± 0.13‰ (2SD, n = 15) for diatomite and −0.29 ± 0.19‰
(2SD, n = 6) for BHVO-2 (Reynolds et al., 2007; Savage et al., 2012).
3. Results
3.1. Soil mineralogy
The parent material of the soil chronosequence (0–335 years) is Cox
Bay beach sand (Singleton and Lavkulich, 1987), which is comprised of
very well-sorted glacial sands with little wearing off and smoothing
sharp edges and corners. The primary minerals present in the beach
sand C material identified by X-ray diffraction and microscopy are
quartz, amphibole, pyroxene, olivine and feldspars, as well kaolinite
precipitating in the dissolution pits of feldspars. The parent material
does not contain inherited clay minerals, except kaolinite present in
the weathered feldspar. We observe an increase of oxalate-extractable
Siox in Bhs, Bs and Bw horizons of the Podzol (CB-335 years)
(Table 1). We also document a strong mobilization of Fe in Podzol
after 335 years, characterized by an accumulation of crystalline and
amorphous Fe oxides in the Bhs horizon, which is related to an increase
of OC content. This co-accumulation of Fe oxides and OC is also observed
in the Belgian Podzol in Bh and Bhs horizon. The content of clay-sized
minerals is quite constant in the Belgian Podzol while we observe an increase of clay content towards the topsoil in the Canadian podzolic soil
chronosequence (Table 1). The content of clay-sized minerals in the entire soil profiles increases over time in the chronosequence.
The mineralogy of the clay fraction in the Cox Bay podzolic
chronosequence is dominated by quartz, amphiboles, chlorites, vermiculite, mixed-layers minerals (MLM), smectite, illite, and kaolinite and
evolves depending on soil age and the development of soil horizons
(Fig. 2). In the youngest soil profile (CB-120 years), the clay mineralogy
is characterized by the presence of quartz, Na-feldspar and amphiboles
as primary minerals and kaolinite, chlorite and illite as pedogenic clay
minerals (data not shown). XRD patterns display similar mineral compositions in the E horizons of CB-175 years and CB-270 year profiles.
In those soil horizons, peaks at 1.40, 1.00, 0.83 and 0.70 nm, correspond
respectively to chlorite, illite, amphibole and kaolinite (disappearance
of the 0.7 nm peak after K 550 °C treatment). A band at 1.40 nm
(Mg 20 °C treatment) that shifts to 1.60–1.70 nm after Mg–eg treatment
due to swelling indicates the presence of discrete smectite. In addition,
the combination of the peaks at 1.40 nm after Mg-20 °C and Mg–eg, and
the collapse of the peak from 1.10 to 1.00 nm due to the dehydration
after a K-saturation followed by heating correspond to vermiculite. Finally, the presence of a wide peak at 1.20 nm after Mg-20 °C treatment
that shifts after Mg–eg treatment indicates irregularly mixed-layer minerals (MLM).
In the CB-335 years profile, mineralogical differences were observed. In the E horizon, relative to the E horizons of CB-175 years and
CB-270 year profiles, XRD patterns show a strong decrease of the abundance of kaolinite (the 0.70 nm peak has almost disappeared), absence
Table 1
Summary of the major soil physical and chemical characteristics (for the fine earth b2 mm and the clay fraction b2 µm) of the investigated soils.
Horizon
Depth
pH
cm
Sioxa
Soil fractions
Sand
Silt
Sidcb
Feox
Fedcb
Alox
Alp
OCtotb
Clay
Si
−1
%
Clay fraction
g.kg
Al
Fe
Ge
μg·g−1
%
Cox Bay 120 years (Dystric Cambisol)
BC
0–75
5.9
99.2
0.6
0.3
0.1
0.4
1.7
1.8
0.7
0.5
9.5
15.8
8.6
12.0
2.9
Cox Bay 175 years (Dystric Cambisol)
E
0–3
5.4
90.2
Bw
3–44
5.8
99.0
BC
44–75
5.9
99.6
7.1
0.6
0.2
2.7
0.4
0.1
0.1
0.1
0.1
1.1
0.3
0.2
1.4
1.7
1.0
2.8
2.0
1.3
0.4
0.9
0.6
0.3
0.7
0.4
35.2
4.3
2.7
22.0
13.5
16.1
8.3
7.4
8.9
8.1
10.5
9.4
5.7
2.0
3.1
Cox Bay 270 years (Haplic Podzol)
E
0–7
4.6
Bh
7–23
5.1
Bw
23–57
5.3
BC
57–75
5.4
90.8
97.2
97.4
98.2
6.1
1.7
1.8
1.1
3.1
1.0
0.8
0.7
0.1
0.2
0.2
0.2
0.8
0.4
0.4
0.3
1.2
2.3
2.1
1.7
2.5
2.6
2.3
2.2
0.6
1.3
1.1
1.4
0.4
0.9
0.8
1.0
13.3
16.1
10.4
7.6
23.3
15.5
16.1
13.1
9.1
8.7
8.8
10.0
5.6
12.1
12.0
11.3
8.9
3.0
2.8
2.4
Cox Bay 335 years (Placic Podzol)
E
0–16
4.8
Bh
16–23
5.6
Bhs
23–24
Nd
Bs
24–28
5.1
Bw
28–60
5.1
82.3
88.0
90.0
94.9
96.1
14.4
8.7
6.8
4.2
2.4
3.0
2.8
2.9
0.9
1.4
0.1
0.5
1.1
2.9
2.5
0.4
1.0
1.6
1.0
0.9
0.2
3.3
21.5
3.7
2.1
0.5
4.4
44.0
4.0
2.9
0.6
8.8
5.2
7.8
6.4
0.5
5.0
4.4
1.3
1.1
10.7
36.8
17.8
5.2
3.8
26.5
16.9
6.2
14.2
13.2
10.5
14.3
9.4
19.6
20.3
1.9
6.2
30.2
9.3
7.5
12.9
5.8
4.6
3.0
3.8
Gaume (Haplic Podzol)
E
19–35
Bh
35–40
Bhs
40–47
Bs
47–58
BC
70–100
94.0
89.0
90.0
91.6
92.9
3.1
7.0
6.0
3.4
2.5
2.9
4.0
4.0
5.0
4.6
0.0
0.1
0.3
0.5
0.2
0.0
0.3
0.3
0.4
0.1
0.1
4.6
5.5
0.7
0.1
2.1
16.7
16.8
6.8
2.8
0.04
1.3
2.0
2.2
0.6
Nd
Nd
Nd
Nd
Nd
1.3
14.4
6.1
3.6
0.7
12.6
11.0
7.7
9.9
9.8
7.6
7.8
9.6
12.7
13.9
10.9
14.5
19.7
12.9
13.8
2.8
2.9
1.5
1.9
1.4
5.0
4.7
4.8
5.1
4.6
Nd = not determined.
a
Dithionite- (dcb), oxalate- (ox) and pyrophosphate- (p) extractable contents of Fe, Al and Si.
b
Total organic carbon.
J.-T. Cornelis et al. / Geoderma 235–236 (2014) 19–29
A
B
CB-175 yrs : E
23
CB-270 yrs : E
1.39
0.83
0.99
1.39
1.62
Mg eg
1.20
0.83
1.19
0.99
Mg eg
Mg 20°C
Mg 20°C
K 550°C
K 550°C
K 300°C
K 300°C
K 150°C
K 150°C
K 20°C
4
10
K 20°C
4
10
2 - theta [°]
C
2 - theta [°]
D
CB-335 yrs : E
CB-335 yrs : Bh
0.70
1.64
1.39
1.39
0.99
0.83
0.70
Mg eg
1.19
0.83
0.99
Mg eg
Mg 20°C
1.20
K 550°C
Mg 20°C
K 550°C
K 300°C
K 300°C
4
K 150°C
K 150°C
K 20°C
K 20°C
10
2 - theta [°]
4
10
2 - theta [°]
Fig. 2. XRD patterns of the clay-sized fraction (b2 μm) of soils of the Cox Bay soil chronosequence after six treatments: K-saturation followed by drying at 20, 105, 300 and 550 °C, and Mgsaturation followed by drying at 20 °C and saturation with ethylene-glycol. (A) CB-175 years E horizon, (B) CB-270 years E horizon, (C) CB-335 years E horizon, (D) CB-335 years Bh
horizon. Spacings of major reflections are in nanometers.
of chlorite (no peak at 1.40 nm after K treatments), and increase of the
relative abundance of smectite compared to vermiculite (increase of the
peak at 1.60–1.70 nm and almost no peak at 1.40 nm after the Mg–eg
treatment). In the Bh horizon relative to the E horizon of CB-335 year
profile, XRD patterns show the presence of kaolinite and chlorite, absence of smectite (no swelling after Mg–eg treatment), increase in the
abundance of vermiculite, and a decrease of the abundance of MLM
(smaller peak at 1.20 nm after Mg 20 °C treatment).
The mineralogy of the Belgian Podzol (Gaume) is compared to the
mineralogy of the Canadian Podzol. The primary minerals of the loess
contain quartz, feldspars, micas and small amounts of trioctahedral
chlorites and amphiboles (Van Ranst et al., 1982). The mineralogy of
the clay fraction in the Belgian Podzol is comprised of vermiculite, smectite, hydroxyl-interlayered vermiculite, chlorite, MLM and kaolinite
(Herbauts, 1982).
As we are not able to precisely quantify each type of 2:1 minerals on
the powder of the clay fraction (chlorite, smectite, vermiculite, illite,
MLM) with Siroquant software, we carried out the clay mineralogy
quantification in the soil chronosequence by separating the minerals
in the clay fraction in 4 groups: quartz, amphiboles, kaolinite
and 2:1 minerals (Fig. 3). Compared to the mineralogy of BC horizon
(CB-120 years) at the initial stage of soil formation (quartz = 15%,
amphiboles = 63%, kaolinite = 5% and 2:1 minerals = 17%), the
quantification of clay mineralogy indicates an increase of the relative
24
J.-T. Cornelis et al. / Geoderma 235–236 (2014) 19–29
(at t = 335 years); and 30ΔSiB–BC varying from − 0.17‰ (at t =
175 years) to −0.32‰ (at t = 335 years).
A comparable depletion/enrichment in light 28Si in the clay fraction
during pedogenesis is found in the Belgian Podzol (30ΔSiE–BC =
+0.29‰; 30ΔSiB–BC = −0.27‰) from a similar temperate climate but
with a different parent material and rainfall conditions (Fig. 4D).
3.3. Geochemical modifications in the clay fraction over time
Fig. 3. Quantitative evolution of the mineralogy in the clay-sized fraction of the Cox Bay
soil chronosequence. The clay-size mineralogy is comprised of primary minerals (quartz
and amphiboles) and pedogenic clay minerals (kaolinite and 2:1 minerals). Chlorite,
vermiculite, smectite, illite and mixed-layer minerals (MLM) are the 2:1 aluminosilicates
encountered in the podzolic chronosequence.
abundance of kaolinite (+14%) and 2:1 minerals (+6%) in E horizon of
the 175-year-old soil. Then we observe a strong decrease of the relative
abundance of kaolinite in older and more weathered E horizons: −12%
in the 270-year-old soil and − 18% in the 335-year-old soil, while the
relative abundance of 2:1 minerals is constant between the two oldest
soils (=18%). In the Bh and Bhs horizons of the 335-year-old soil, we
note an increase of the relative abundance of kaolinite (+ 7% and
+ 12%, respectively) compared to the stronger weathered E horizon
(=1%). The evolution of primary clay-sized minerals is characterized
by a decrease of the relative abundance of amphiboles in the early
stage of soil formation (−21%), then by a relative increase of the abundance (+14 and +19%) in the more weathered E horizons, which is related to the decrease of kaolinite, while quartz content remains quite
constant (=16 ± 2%) during pedogenesis.
3.2. Si isotopic modifications in the clay fraction over time
In our study, pedogenic clay minerals in the clay fraction of BC horizon are considered as “unweathered” secondary minerals compared to
pedogenic clay minerals in more weathered horizon (E, Bh, Bhs and
Bs) since BC horizon is not yet reached by the podzolic weathering
front (Lundström et al., 2000). In the Cox Bay soil chronosequence, we
therefore compare the Si isotopic signatures of the clay fraction in
each soil horizon with those in the “unweathered” clay fraction in the
BC pedogenic horizon.
In the Cox Bay chronosequence, Si in the “unweathered” clay
fraction (BC horizon; δ30Si = − 0.52 ± 0.16‰, 2SD, n = 3) is
isotopically lighter compared to the primary lithogenic minerals in the
parent beach sand material (C material; δ30Si = −0.27 ± 0.10‰, 2SD,
n = 3) (Fig. 4A). In the early phase of soil formation, the difference of
Si isotope signature between the lithogenic primary minerals in the
sand fraction of the C material and the clay fraction in BC material, 30ɛ
is −0.25‰ (min − max = − 0.12 − 0.37‰). This is not the fractionation factor due to precipitation of pedogenic clay minerals as the clay
fraction also comprises lithogenic primary minerals.
Relative to the “unweathered” BC clay fraction (δ30Si = − 0.52 ±
0.16‰, 2SD, n = 3), the clay fraction of the topsoil eluvial E horizons
shows depletion in light 28Si (i.e., less negative δ30Si values: from
−0.33 ± 0.02‰ to −0.10 ± 0.22‰ Fig. 4B, C). The clay fraction in the
subsoil illuvial Bh–Bs horizons is isotopically lighter (i.e., enriched in
light 28Si) than “unweathered” BC clay fraction (δ30Si from − 0.60 ±
0.06‰ to − 0.84 ± 0.08‰ ‰; Fig. 4B, C). The magnitude of light
Si depletion/enrichment in the clay fraction increases with soil age,
with 30ΔSiE–BC varying from + 0.20‰ (at t = 175 years) to + 0.42‰
As the clay fraction becomes relatively more depleted in Si, the clay
fraction becomes more enriched in light 28Si (Fig. 5A). Our results show
that Si isotopic signature of the clay fraction becomes increasingly light
with enrichment in Al (higher Al/Si ratio in the clay fraction) (Fig. 5B).
The enrichment in light 28Si (and the increase of Al/Si ratio)
in the clay fraction also relates to an increase in the proportion of
poorly-crystalline Si components in the clay fraction (estimated by the
Siox/Siclay ratio). As the Si-bearing phases of the clay fraction accumulates poorly-crystalline aluminosilicates, the Si isotopic composition becomes more enriched in light Si isotope (Fig. 5C). We observe also that
the enrichment in light 28Si in the clay fraction is not systematically related to a relative depletion in Ge, i.e. lower Ge/Si ratio (Fig. 5D).
4. Discussion
4.1. Evolution of clay-sized mineralogy
Different processes, such as transformation and neoformation, modify the chemical composition of the clay mineral within soil profile and
control the clay content and mineralogy during pedogenesis. As water
acts to mediate chemical reactions and to transport reactants and products from topsoil (Chadwick and Chorover, 2001), we observe the
highest content of pedogenic subproducts (clay-sized minerals) in the
top- and subsoils (0–24 cm). The depth where clay-sized minerals concentrate (~3%) increases over time, which highlights the deepening of
the weathering front: 0–3 cm after 175 years, 0–7 cm after 270 years,
and 0–24 cm after 335 years. We show that the chemical modifications
of clay mineral structure in the podzolic weathering front mobilize Al
(and Fe) and Si from secondary minerals over time. The evolution of
Al/Si in the clay fraction substantiates the preferential mobilization of
Al, relative to Si, during the dissolution of secondary clay minerals, in
particular in the presence of organic acids with high complexing capacities, such as those encountered in Podzols (Sokolova, 2013; Stumm,
1992). The clay mineralogy evolution (Fig. 2) in Podzols studied here
under maritime temperate climate is very similar to the ones observed
from postglacial moraines (Righi et al., 1999) and tills (Egli et al.,
2002) in Switzerland. The aluminization of primary clay minerals,
such as chlorites, leads to formation of irregularly-interstratified minerals in the moderately acid B horizons. In the stronger weathering E
system, Al-removal from interlayers by organic complexing agents
leads to the formation of vermiculite. Further alteration induces the formation of smectite-like minerals in the E eluvial horizon. Finally, the Siox
content (Table 1) confirms that the formation of poorly-crystalline aluminosilicates (ITM) occurs when the concentration of organic acids is
sufficiently low to allow the precipitation of Al with Si, as suggested
by Ugolini and Dahlgren (1987) in the fulvate bicarbonate theory of
podzolization. The clay mineralogy evolves with increasing weathering
in the age sequence and formation of typical podzolic soil horizons
(E, Bh, Bhs, Bs, Bw), which is in good agreement with the formation of
two geochemical compartments during podzolization (Ugolini and
Sletten, 1991). The upper E-Bh compartment is controlled by organic
acids as major proton donors and complexing metals, which leads to
dissolution of primary and secondary minerals. In the lower Bhs-Bs
compartment, the absence of organic acids leads to a less aggressive
weathering system mainly controlled by inorganic acids (carbonic and
nitric acids).
J.-T. Cornelis et al. / Geoderma 235–236 (2014) 19–29
25
Fig. 4. Silicon isotopic signature (δ30Si ‰; mean values ± standard deviation represented by error bars) in the clay-sized fraction depending on soil ages and in primary lithogenic minerals
in the beach sand parental material. (A): 0- and 120-year-old soil fraction (δ30Si of primary minerals in beach sand in black and δ30Si of the clay fraction of the 120-year-old BC horizon in
blue), (B): 175- and 270-year-old clay fractions (175 years = red Δ; 270 years = green ◊), (C): 335-year-old clay fraction (purple ○), and (D): clay fraction in an undated Belgian Podzol
(brown □). After only 175 years (B), we observe the depletion in light 28Si in the clay fraction of the eluvial E horizon and enrichment in light 28Si in the clay fraction of deeper illuvial soil
horizon; respectively, relative depletion in light 28Si (+0.20‰) and relative enrichment in light 28Si (−0.17‰) compared to the original Si isotopic signature of the unweathered clay fraction in the BC horizon. The isotopic fractionation increases over time with an enrichment in heavy 30Si of +0.42‰ in the clay fraction of the E horizon and a concomitant enrichment in light
28
Si of −0.32‰ in the clay fraction of the Bhs horizon (after 335 years). We observe exactly the same tendency in the Belgian Podzol with enrichment in light 28Si in the clay fraction of the
Bhs horizon of −0.27‰ compared to the unweathered clay fraction in BC horizon. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)
Four important mineralogical evolutions are observed in the Cox Bay
soil chronosequence, as a result of podzolization: (i) the neoformation
of kaolinite, illite and chlorite from dissolution of primary minerals at
the very beginning of soil formation, (ii) the disappearance of kaolinite
in the strongest weathered E horizon, then (iii) the increase of relative
abundance of kaolinite in Bh and Bhs horizons compared to E horizon
(Fig. 3), and finally (iv) the accumulation of imogolite-type materials
in Bhs and Bs horizons (Table 1).
4.2. Dissolution and re-precipitation of pedogenic clay minerals during
podzolization
Since the clay fraction of soils comprises aluminosilicates and Fe-,
and Al-oxyhydroxides, Si in the clay fraction includes Si incorporated
in primary minerals (quartz and amphiboles), secondary minerals
(kaolinite and 2:1 minerals) and Si adsorbed onto Fe oxyhydroxides.
In the Bhs horizon of the 335-year-old soil, the high content of free Fe
(Fedcb = 44 g·kg− 1) is in the same order of magnitude than in a
weathering sequence in Cameroon (20–85 g·kg− 1) (Opfergelt et al.,
2009), for which the variations of δ30Si values in the clay fraction due
to adsorption onto Fe oxides are known (Opfergelt et al., 2010). We
have to take into account the pool of Si adsorbed onto Fe oxides in the
clay fraction as this Si pool significantly influences the enrichment in
light 28Si in the clay fraction: the difference of the Si isotope signature
in the clay fraction of B horizons before and after dithionite-treatment
(i.e., after the release of Si from the surface of Fe oxides) in the
Cameroon weathering sequence varies between 0.08 and 0.45‰
(Opfergelt et al., 2010). However, all of the Fe in the Cameroon
weathering sequence is in the clay fraction, while in the temperate
soils of the Cox Bay chronosequence, only 20% of the bulk Fe content
is in the clay fraction for Bhs horizon (=8.8 g·kg−1), where we observe
the largest enrichment in light Si isotope. In eluvial E horizons, we
observe the largest depletion in light 28Si while the Fe content in
the clay fraction represents between 70 and 100% of the total Fe
concentration in bulk soil (until 2.8 g·kg−1). The ratio of Fe oxides in
the clay fraction to Si content in the clay fraction is similar between
Bhs (14%) and E (13%) horizons, while the Si isotope composition in
the clay fraction follows opposite trends in these two horizons. As a
consequence, we assume that the δ30Si values of the clay fraction of
Belgian and Canadian temperate soils can be considered representative
of the Si isotopic composition of the primary and secondary silicates,
and not significantly influenced by the fractionation of Si isotopes
through adsorption onto Fe oxides. The role played by the Si adsorption
onto Fe oxides on Si isotope compositions of the clay fraction must
however be further investigated.
It is well established that the preferential incorporation of light 28Si
during neoformation of secondary pedogenic minerals accounts for
their isotopically lighter signature relative to primary lithogenic minerals (Georg et al., 2009; Opfergelt et al., 2010; Ziegler et al., 2005).
The Si isotope composition of the soil clay fraction depends on the degree of soil weathering and the evolution of the clay mineralogy
(Opfergelt et al., 2010, 2012; Ziegler et al., 2005).
26
J.-T. Cornelis et al. / Geoderma 235–236 (2014) 19–29
Fig. 5. Evolution of Si isotope composition with elemental composition (Si, Al, Ge) and the proportion of poorly crystalline Si (Siox/Siclay) in the clay fraction for the Cox Bay soil
chronosequence (175-year-old soil = red Δ; 270-year-old soil = green ◊; 335-year-old soil = purple ○) and for the Gaume Podzol (brown □). (For interpretation of the references to
color in this figure legend, the reader is referred to the web version of this article.)
Using the quantification of primary minerals (quartz and amphiboles) and secondary minerals (kaolinite and 2:1 minerals) and the Si
isotope signature of lithogenic primary minerals (− 0.27‰), we can
compute δ30Si value of “unweathered” secondary clay minerals in the
clay fraction of BC horizon (− 1.40‰; Table 2). The isotopic fractionation factor between primary lithogenic minerals and secondary
pedogenic minerals (30ε = δ30Simin I − δ30Simin II) is therefore
−1.13‰. The mass balance approach (Table 2) shows also a progressive
depletion in light 28Si in secondary minerals of the E horizon (from
−0.51‰ to 0.64‰) and an enrichment in light 28Si in secondary minerals of the illuvial horizons (until − 2.31‰). In identical bio-geoclimatic conditions, the Si isotopic fractionation associated with the
Table 2
Quantification of primary and secondary minerals in the clay-sized fraction of the Cox Bay soil chronosequence. The clay-sized quantification is then used for the Si isotopic mass balance
approach.
Measured data
BC horizon
(120 years)
E horizon
(175 years)
E horizon
(270 years)
E horizon
(335 years)
Bh horizon
(335 years)
Bhs horizon
(335 years)
Computed
Primary minerals
(% in the clay fraction)a
Secondary minerals
(% in the clay fraction)
δ30Si(‰)
in the clay fraction
δ30Si (‰)
of pedogenic clay mineralsb
Δ30SiBC–x (‰)c
78
22
−0.52
−1.40
–
59
42
−0.32
−0.39
+1.01
75
25
−0.33
−0.51
+0.89
81
19
−0.10
+0.64
+2.04
72
28
−0.45
−0.92
+0.48
70
30
−0.84
−2.31
−0.91
a
Mineralogy of the clay fraction quantified using the Siroquant software V4.0; primary minerals = quartz + amphiboles; secondary minerals = kaolinite + 2:1 minerals
(vermiculite, smectite, illite, chlorite, mixed-layers minerals).
b
The δ30Si of secondary minerals present in the clay fraction is computed as follows: δ30Simin II = ((δ30Siclay fraction − % min I ∗ δ30Simin I) / % min II), where min I = primary minerals,
min II = secondary minerals and δ30Simin I = −0.27‰.
c
Si isotope discrimination between “unweathered” clay minerals in BC horizon and pedogenic clay minerals in the “x” horizon of interest (x = E, Bh or Bhs horizons): δ30SiE–δ30SiBC or
δ30SiBC–δ30SiBh/Bhs.
J.-T. Cornelis et al. / Geoderma 235–236 (2014) 19–29
dissolution of primary lithogenic minerals and neoformation of secondary pedogenic minerals should generate comparable Si isotopic signatures in the clay fraction in the entire soil profile with no evolution
over time given identical fractionation factor between the primary and
secondary Si pools. Here, we show that that the signature of secondary
minerals varies in the soil profile and the relative depletion/enrichment
in E and B horizons increases with time in the Cox Bay chronosequence.
The dissolution of primary minerals and precipitation of secondary minerals therefore cannot explain the increasing depletion/enrichment in
light 28Si in the clay fraction over time and with depth. This highlights
that the evolution of δ30Si values in the clay fraction of the soil profiles
observed here rules out the weathering of primary minerals (lithogenic
Si pool) as the sole source for the neoformation of secondary minerals in
the clay fraction.
Germanium (Ge), a chemical analog of Si, generally follows similar
inorganic geochemical pathways than Si (Froelich and Andreae, 1981).
However, secondary pedogenic (clay) and biogenic (phytoliths) minerals display contrasting Ge/Si ratios: neoformed clay minerals are
enriched in Ge (higher Ge/Si) while biogenic silica polymerized in
plants as phytoliths is depleted in Ge (lower Ge/Si) (Derry et al., 2005;
Kurtz et al., 2002). Although there is a negative relationship between
Ge/Si ratios and δ30Si in the youngest soils (Cambisols) of the Canadian
soil chronosequence, the absence of a relationship between Ge/Si and
δ30Si ratios in the oldest soil (Podzol) of the Canadian chronosequence
and in the Belgian Podzol (Fig. 5D) allows us to dismiss the dissolution
of phytoliths (biogenic Si pool) as a major source of Si for clay
neoformation. This process would be characterized by enrichment in
light 28Si and depletion in Ge in secondary clay minerals relative to
beach sand parent material, as phytoliths are Ge-depleted (low Ge/Si
ratio) relative to primary minerals (Derry et al., 2005).
The mass balance approach (Table 2) shows that the enrichment in
light 28Si of secondary minerals of Bhs horizon (−2.31‰) compared to
the “unweathered” secondary minerals in the BC horizon (− 1.40‰)
partly explains the depletion in light 28Si of secondary minerals in the
clay fraction of E horizon (+ 0.64‰) for the oldest soil (Podzol CB335 years). Our data highlight that the isotopic fractionation due to
preferential release of light 28Si during dissolution of secondary
minerals in the E horizon (Δ30SiE–BC = + 2.04‰) partly accounts
for the enrichment in light 28Si during re-precipitation of new clay
minerals in Bhs horizon (Δ30SiBhs–BC = −0.91‰). This combined with
the fact that kaolinite is progressively dissolved in the E horizon and
is almost completely dissolved in the strongly weathered E horizon
(CB-335 years) (Fig. 3), highlights that 28Si is redistributed in the soil
profile through re-precipitation of new pedogenic clay minerals deeper
in the soil profile and leaching. As a part of Si precipitating during the
neoformation comes from the dissolution of secondary clay minerals,
we name those new clay minerals as “tertiary minerals”. This implies
that the preferential lessivage of clay particles enriched in light 28Si
and the resulting relative accumulation of primary clay-sized minerals
in topsoil cannot be responsible for the on-going enrichment in light
28
Si in the clay fraction. Indeed, the increasing enrichment in light 28Si
in new tertiary minerals (tertiary kaolinite) in B horizons can only be
related to a Si source progressively enriched in light 28Si over time.
Kaolinite seems to play a key role in the successive formation of clay
minerals as the content of 2:1 clay minerals is quite constant during pedogenesis in the soil chronosequence (Fig. 3).
The preferential release and incorporation of light 28Si during dissolution and re-precipitation of clay minerals in the pedogenic Si pool account for the Si isotopic depletion/enrichment in the clay fraction over
time in the podzolic chronosequence. The preferential incorporation of
light 28Si during precipitation of Si released from the dissolution of pedogenic clay minerals (in E and Bh horizons) explains the increasing enrichment in light 28Si in newly-formed clay minerals (tertiary clay
minerals in Bhs horizon) during podzolization. This is confirmed by
the fact that pedogenic clay minerals in E horizons are increasingly
heavier over time (Table 2), showing that the dissolution of pedogenic
27
clay minerals discriminate against the release of heavy 30Si as already
demonstrated for diatoms (Demarest et al., 2009) and crystalline basalt
(Ziegler et al., 2005). Besides the lithogenic and biogenic Si pools, we
provide evidence that pedogenic Si pool is therefore involved in the
neoformation of pedogenic clay minerals and as such in the evolution
of their Si isotope signatures (Fig. 6).
4.3. Implications for podzolization theory
For the first time, we document enrichment in light 28Si in secondary
clay minerals over time in a podzolic soil chronosequence. The highest
enrichment in light 28Si and oxalate-extractable Siox in Bhs/Bs horizons
relative to E/Bh horizons (Fig. 5C; Table 2) highlights that the dissolution of secondary aluminosilicates in E/Bh horizons acts as a Si source
for formation of poorly-crystalline aluminosilicates (imogolite-type materials ITM) in Bhs/Bs horizons. The release of Si from the dissolution of
primary and secondary clay minerals and precipitation of dissolved Si
with Al released by microbial decomposition from the organic ligands
(Lundström et al., 1995) can explain the formation of ITM in Bhs/Bs horizons (Ugolini and Dahlgren, 1987). During podzol development, ITM
undergo additional dissolution for the re-precipitating Si as crystalline
tertiary clay minerals in Bhs horizon. The evolution of Si isotopic signature in pedogenic clay minerals of the podzolic soil chronosequence
therefore corroborates the process of dissolution and re-precipitation
of aluminosilicate phases during podzolization (fulvate bicarbonate theory; Ugolini and Dahlgren, 1987). We can infer that low contents of
poorly-crystalline ITM in the Bhs/Bs horizons play a key role in the evolution of Podzols and the progressive enrichment in light 28Si in pedogenic clay minerals. The absence of ITM in the Bh horizon and the
lighter δ30Si in Bhs/Bs indicates their high reactivity during podzolization, dissolving as organic-rich Bh horizon forms and precipitating as
Fe-, Si-, and Al-enriched Bhs/Bs horizons form. This is confirmed by
the high reactivity of ITM also reflected in Ge/Si and δ30Si patterns in
soil solutions of the Santa Cruz soil chronosequence, which indicates
seasonal precipitation and dissolution of hydroxyaluminosilicates such
as allophane (White et al., 2012). The positive correlation between
Siox/Siclay and δ30Si values in the clay fraction (Fig. 5C) highlights that
during podzolization, pedogenic clay minerals become enriched in
light 28Si together with Al in the poorly-crystalline part of the clay fraction. Based on these findings, poorly-crystalline aluminosilicates can be
regarded as a temporary reactive reservoir of light 28Si in Bs horizon.
This reservoir acts as a source of light 28Si in tertiary crystalline clay minerals, such as tertiary kaolinite, in Bhs horizon that will develop in the
current Bs horizon during podzolization. The dissolution and reprecipitation of pedogenic clay minerals are therefore an important
podzol-forming process (Fig. 6).
4.4. Implications for tracing the effects of environmental changes on soils
In the Cox Bay soil chronosequence, we show that the production of
acidity (protons and complexing organic acids) in temperate forests and
the subsequent Podzol formation imply heavy 30Si enrichment in pedogenic clay minerals of E horizons relative to the “unweathered” clay
minerals in BC horizon; Δ30SiE–BC increasing from + 1.01 to + 2.04‰
in ca. 200 years (Table 2). The preferential loss of light 28Si in weathered
clay minerals in E horizons compared to the “unweathered” clay minerals in BC horizon is recorded in the Si isotope signature of pedogenic
clay minerals on very short time-scale. Moreover, the Si isotope fractionation between the “unweathered” clay minerals in BC and pedogenic clay minerals precipitating in Bhs (Δ30SiBC–Bhs) of −0.91‰ highlights
that a part of light 28Si released in topsoil is used for re-precipitation in
the subsoil (Table 2). As a consequence, Si isotope signatures in the clay
fraction of soils should be tested in other systems to trace the modifications of pedogenic clay minerals insoil–plant systems, such as developed in highly weathered tropical and subtropical environments
(Ferralsols, Lixisols, Nitisols, …), in frozen soils (Cryosols), in soils
28
J.-T. Cornelis et al. / Geoderma 235–236 (2014) 19–29
Fig. 6. Conceptual representation of the contribution of Si released from the dissolution of primary and secondary Si pools (lithogenic, biogenic and pedogenic) to the re-precipitation of
new “tertiary” clay minerals during podzolization. Phase I → phase II (C → BC) = transition from the parent C material to the pedogenic BC horizon with neoformation of secondary clay
minerals. Phase II = formation of typical podzolic soil horizons: E, Bh, Bhs, Bs and Bw. Phase II → phase III = transition from young to older Podzol characterized by (i) the deepening of
the E horizon where secondary clay minerals are weathered and enriched in heavy 30Si (EII → EIII), and (ii) B horizons tertiary clay minerals re-precipitate and are enriched in light 28Si (BCII
→ BhsIII).
characterized by illuviation of clay minerals (Luvisols), in young soils
(Cambisols) and in soils with high biological activity (Chernozems). Si
isotope composition of pedogenic clay minerals can be useful to trace
and quantify the impact of environmental changes (temperature, rainfall, acid deposition, land use …) on pedogenic clay evolution. This is
central to a better understanding of soil development and associated
terrestrial biogeochemical processes.
5. Conclusions
The process of dissolution of pedogenic clay minerals during podzolization is confirmed by the Si isotopic signature of the clay fraction in a
podzolic soil chronosequence (Cox Bay, Vancouver Island). Our dataset
shows Si isotopic, geochemical and mineralogical trends with depth and
as a function of pedogenic time, providing an orthogonal dataset which
sheds light on the origin and evolution of pedogenic clay minerals in the
clay fraction. The depletion in light 28Si in pedogenic clay minerals in
topsoil increases over time (from + 1.01 to + 2.04‰) and a part of
light 28Si released accounts for the relative enrichment in light 28Si in
pedogenic clay minerals in subsoil (−0.91‰). This highlights that Si released from the partial dissolution of secondary clay minerals in topsoil
contributes to the neoformation of tertiary clay minerals in subsoil. Clay
mineral dissolution has often been regarded as an irreversible process,
while the increase of 28Si enrichment over time in the clay fraction documented in this study indicates successive formation of clay minerals,
which depends on the downward movement of the weathering front
in the soil. The continuous weathering of pedogenic clay minerals is
an important process in the formation of Podzols as we show that the
Si released in soil solution contributes to the reformation of clay minerals deeper in soils over very short time scales (ca. 300 years). The recording of Si isotopic ratios in the clay fraction as a function of the age of
soil formation is therefore an untapped resource for tracing pedogenic
processes controlling the Si incorporation in pedogenic clay minerals
during podzolization, and offering new perspectives for unraveling the
genesis of pedogenic subproducts in various soil types. This has important implications as the process of dissolution and re-precipitation of
pedogenic clay minerals would play a major role in several soil biogeochemical processes such as the retention of plant nutrients, the preservation of organic carbon from microbial decomposition, and the transfer
of elements and pollutants from land to ocean. Further investigations
are needed for quantifying the contribution of pedogenic Si pool to
newly-formed clay minerals (tertiary, quaternary …) compared to the
contribution of lithogenic and biogenic Si pools. Our dataset shows
that the Si isotope compositions of soils are influenced not only by biogenic (phytolith formation/dissolution) and litho-, pedo-genic processes (primary mineral dissolution and secondary mineral precipitation)
but also by a more advanced weathering process, i.e. successive formation of pedogenic clay minerals. This should be taken into account when
δ30Si values of the bulk soil and soil solutions are used for studying soil
weathering degree and tracing dissolved and particulate Si transferred
from soil–plant systems to the hydrosphere.
Acknowledgments
We thank A. Iserentant, C. Givron, P. Populaire, A. Lannoye, I. Caignet,
P. Sonnet, M. Detienne (UCL), H. Schreier, S. Smukler, B. Kieffer (UBC),
as well F. Talbot and A. Cornelis for field and laboratory assistance,
V. Lai and M. Soon (UBC) for assistance in element analysis and
K. Gordon (UBC) for assistance in Si isotopic analysis. We thank
M. Brzezinski (University of California Santa Barbara) for providing us
diatomite. J-T.C. is supported by “Fonds National de la Recherche
Scientifique” of Belgium (FNRS; Postdoctoral Researcher Grant). This research was also supported by the “Fonds Spécial de Recherche” of the
UCL and by D.W. NSERC Discovery Grant.
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