1
Evaporation fractionation in a peatland drainage network
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affects stream water isotope composition
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Matthias Sprenger, Doerthe Tetzlaff, Claire Tunaley, Jonathan Dick and Chris
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Soulsby.
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Northern Rivers Institute, School of Geosciences, University of Aberdeen, UK
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Corresponding author: Matthias Sprenger, email:
[email protected], phone: +44
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(0)1224 27 2321, address:
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School of Geosciences, University of Aberdeen, Elphinstone Road, Aberdeen, AB24 3UF, UK
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Key points:
evapotranspiration and negatively with discharge.
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The interplay between the precipitation, evaporation, hydrologic connectivity and groundwater explains the isotopic fractionation signal.
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Kinetic isotopic fractionation in peatland drainage correlates positively with potential
The isotopic enrichment within the peatland drainage network is imprinted on the stream water.
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Keywords: 1804 Catchment, 1890 Wetlands, 1818 Evapotranspiration, 1850 Overland flow,
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1041 Stable isotope geochemistry
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Abstract
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There is increasing interest in improving understanding of evaporation within a catchment for
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an enhanced representation of dominant processes in hydrological models. We used a dual-
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isotope approach within a nested experimental design in a boreal catchment in the Scottish
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Highlands (Bruntland Burn) to quantify the spatio-temporal dynamics of evaporation
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fractionation in a peatland drainage network and its effect on stream water isotopes. We
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conducted spatially distributed water sampling within the saturated peatland under different
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wetness conditions. We used the lc-excess – which describes the offset of a water sample from
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the local meteoric water line in the dual-isotope space - to understand the development of kinetic 1
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fractionation during runoff in a peatland network. The evaporation fractionation signal
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correlated positively with the potential evapotranspiration and negatively with the discharge.
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The variability of the isotopic enrichment within the peatland drainage network was higher with
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higher potential evapotranspiration and lower with higher discharge. We found an increased
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evaporation fractionation towards the center of the peatland, while groundwater seepage from
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minerogenic soils influenced the isotopic signal at the edge of the peatland. The evaporation
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signal was imprinted on the stream water, as the discharge from a peatland dominated sub-
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catchment showed a more intense deviation from the local meteoric water line than the
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discharge from the Bruntland Burn. The findings underline that evaporation fractionation
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within peatland drainage networks affects the isotopic signal of headwater catchments, which
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questions the common assumption in hydrological modelling that the isotopic composition of
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stream waters did not undergo fractionation processes.
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1. Introduction
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The temporal and spatial assessment of evaporation within different landscape units remains
44
unresolved, but knowledge about such process variation within catchments has been shown to
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help calibrate and benchmark rainfall-runoff models [Birkel et al., 2014; Anderton et al., 2002].
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In particular, where models also incorporate solute transport and mixing, constraining and
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calibrating evaporation estimates and their consequent influence on tracer dynamics is
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especially challenging. Evaporation alters the isotopic composition in surface waters, which are
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often used to infer mixing and derive travel time estimates [McGuire and McDonnell, 2006].
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However, for travel time studies, evapotranspiration losses may be crucial to assess the
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temporal variability of hydrological fluxes [Soulsby et al., 2016; Sprenger et al., 2016b]. The
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age distribution of evapotranspiration (the “evapotranspiration time” [Botter et al., 2010]) is
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increasingly considered [van Huijgevoort et al., 2016; Harman, 2015; Soulsby et al., 2015;
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Queloz et al., 2015]. Evaporation induces kinetic fractionation of the composition of stable
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isotopes of the residual water [Craig et al., 1963], while transpiration does not change the
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isotope signal of the water remaining in the catchment [e.g, Zimmermann et al., 1967; Allison
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et al., 1984]. Thus, for waters that experienced evaporation losses, the relation between 18O
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and 2H deviates from the original isotopic composition of precipitation. The original isotopic
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composition is characterized for surface waters by the local meteoric water line (LMWL),
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which describes the 18O - 2H ratio for the rainfall. Water that experienced evaporation
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fractionation will plot below the LMWL; as, for example, observed in streams draining lakes
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[e.g., Gibson and Reid, 2010; Gibson and Edwards, 2002; Isokangas et al., 2015] or beaver 2
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ponds [Burns and McDonnell, 1998]. However, a recent global analysis of the isotopic signal
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in the discharge of headwater catchments in tropical, sub-tropical and temperate regions
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Evaristo et al. [2015] found that the stream water does usually not show signs of evaporation
66
fractionation. As such, traditional assumptions that catchment outflows do not reflect isotopic
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fractionation seem to hold for many of the headwater catchments in these regions. However,
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sites from northern latitudes were not included in Evaristo et al. [2015]. It is known that
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peatlands in these areas, defined as rain-fed (ombrotrophic) bog, groundwater-fed
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(minerotrophic) fen, or a combination of both (bog-fen complex) [Gore, 1983], generally
71
increase the evaporation losses [Bullock and Acreman, 2003]. Consequently, pronounced
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evaporation fractionation signals have been found in pools of peatland drainage networks in
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boreal headwater catchments [Carrer et al., 2016; Gibson et al., 2000]. Nevertheless, the
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temporal variation of these evaporation fractionation processes in headwater catchments with
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peatland drainage networks and the environmental drivers of these processes remain poorly
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understood. Especially, if and when the fractionation is passed on to the stream network is yet
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unknown. A consideration of these evaporation fractionation processes in peatland drainage
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networks is a critical prerequisite for an accurate representation of the hydrological fluxes and
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to infer catchment travel times.
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For the northern upland catchment of the Bruntland Burn in Scotland, Birkel et al. [2011]
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showed that adding evaporation fractionation in the water-saturated areas into their
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hydrological model based on the Craig-Gordon Model [Craig and Gordon, 1965] improved the
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simulation of the isotopic signal in the stream water. However, even the additional inclusion of
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fractionation in the unsaturated soil led to an underestimation of the isotopic fractionation
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during summer periods [Birkel et al., 2014]. Therefore, we know that evaporation fractionation
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plays a role in the isotopic composition of runoff generated from the Bruntland catchment, but
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we do not have an understanding of when, where and in which way the isotopically enriched
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water feeds the stream.
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Thus, the objectives of our study are i.) to investigate stable isotope dynamics in a northern
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upland catchment, and ii.) to quantify spatio-temporal dynamics of evaporation fractionation in
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a peatland drainage network, and iii.) to assess the effect of spatially distributed fractionation
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on the stream water signature.
3
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2. Study site
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The study was conducted in the 3.2 km2 Bruntland Burn (BB) experimental catchment, a
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tributary of the Girnock Burn, in the Scottish Highlands (57°02'18.4"N 3°07'52.0"W). The
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catchment has been the focus of numerous hydrological studies to which the reader is referred
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to for a detailed description [e.g., Blumstock et al., 2015; Tetzlaff et al., 2014; Soulsby et al.,
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2015]. The temperate/boreal oceanic climate is characterized by a mean annual air temperature
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of 7°C and 1000 mm year-1 of mean annual precipitation. Usually, less than 5% and maximum
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of 10% of the precipitation falls as snow [Soulsby et al., 2015]. While the temperature varies
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seasonally with daily means of 2°C in January and 13 °C in July, the rainfall shows a less
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marked seasonal pattern and is fairly evenly distributed throughout the year. The annual
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evapotranspiration (ET) is about 400 mm year-1, mostly occurring between May and August.
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The geology of the BB is dominated by granitic bedrock in the north and metamorphic bedrock
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in the south. Above the low permeability bedrock, glacial deposits and till of varying depths
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from 5 m on the hillslopes to 40 m in the valley bottom, cover up to 70 % of the catchment.
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These drift deposits have a relatively low hydraulic conductivity (10-6 to 10-4 m s-1) and form
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the main source of groundwater storage [Soulsby et al., 2015].
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The soils that have developed on the drift deposits differ according to the topography. On the
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hillslopes, freely draining shallow (1 m) rich in organic material classified as peats and peaty gleys prevail
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(Figure 1). The peat covers about 9 % of the BB, mostly in riparian areas fringing the main
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stream channel. Peaty gleys are also found in the riparian zone, but where slopes are slightly
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higher, and cover 10 % of the total catchment area. The peat is permanently at, or close to,
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saturation and the groundwater table is always within 0.25 m below the surface, while the
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groundwater level can drop deeper than this (to ~0.4m) during very dry conditions in the peaty
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gleys [Blumstock et al., 2016]. The extension of fully saturated soils in the BB was found to be
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2 to 40 % depending on antecedent wetness conditions during the field surveys [Birkel et al.,
119
2010]. The peat soils are characterized by a complex peatland drainage network on the surface
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resulting in fill-and-spill runoff mechanisms [Spence and Woo, 2003]: Due to the low hydraulic
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conductivity in the deeper, more humified peat, water movement is mostly lateral in the upper
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0.2 m of poorly humified peat (acrotelm). Drainage from this horizon in the larger soil pores
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collects in small zero order channels which often drain through networks of small
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interconnected ponds [Tetzlaff et al., 2014; Lessels et al., 2016]. The peatland drainage network
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is also fed by minerogenic groundwater seepage at the edges. Towards the center, where the
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peat accumulates to a raised bog, ombotrophic (rain-fed) conditions prevail. This way, the 4
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peatland drainage network represents as a bog-fen complex a continuum integrating
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groundwater and soil water, then flowing along shallow water tracks and pools.
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For our study, we used a nested experimental design, where we monitored a 0.73 km2 headwater
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catchment (HW1) of the BB (green boundaries in Figure 1). HW1 is characterized by an
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extensive raised peat bog [Blumstock et al., 2015; Dick et al., 2015]. The bog is surrounded by
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a groundwater-fed fen area, receiving groundwater seepage from the surrounding steep
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hillslopes [Lessels et al., 2016]. Peat soils cover 14 % of HW1. Overall, we could distinguish
134
four different landscape units (Figure 1): the total BB, the sub-catchment HW1, the peatland
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drainage network, defined as the extended raised peatland (bog) with dynamically
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connected/disconnected drainage networks, and the riparian zone of the BB, where the drainage
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network is more continuously connected and the contribution of minerogenic drainage from the
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hillslopes is higher. The locations of the water sampling for each unit are shown in Figure 1.
139
The vegetation on the peatland drainage network is dominated by Spagnum spp. mosses along
140
the water tracks and pools. In the riparian zone of the BB, minerogenic drainage from the
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hillslopes allow Molinia caerulea. grasses to compete with Sphagnum Mosses. On the rankers
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and podzols, heather (Calluna vulgaris and Erica spp.) is the dominant shrub, while trees
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(mainly Pinus sylvestris) are mostly limited to the hillslopes covering only 10 % of the BB.
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However, heavy grazing of Red Deer (Cervus elaphus) prevents the development of a forest
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cover on the steep hillslopes.
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3. Data and Methods
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The study spanned the period from 2011-06-01 to 2015-12-31. Meteorological data
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(precipitation, net radiation, air temperature, relative humidity) were collected at 15 minute
149
intervals 1 km away from the BB outlet by Marine Scotland Science. Based on this data, daily
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potential evaporation (PET) was estimated using the Penman-Monteith equation. To account
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for the local vegetation characteristics, the aerodynamic and canopy roughness parameters were
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adjusted to the local conditions in the Scottish Highlands as suggested by Dunn and Mackay
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[1995]. While the absolute values will vary depending on the aerodynamic and canopy
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roughness parameters, the temporal dynamics, which are of relevance for our analysis, are not
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affected by the parameter choice. In our analysis, we use PET as a measure for the potential
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evaporation without an arbitrary partitioning between evaporation and transpiration, since the
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evaporation and the transpiration is found to have a linear relation in temperate climates [Renner
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et al., 2016; Schwärzel et al., 2009]. The daily discharge at the BB outlet was derived from 15 5
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min stage height records (Odyssey capacitance probe, Christchurch, New Zealand) in a stable,
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rated stream section. Water levels were translated into discharge by a rating curve (y = 4E-
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10x5.2119) based on at least monthly gaugings in the years 2011 and 2012 and event based
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gauging since then to address bias towards base flows. Stream water at the outlets of both HW1
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(from May 2014 onwards) and BB as well as rainfall were sampled on daily basis using ISCO
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3700 (Teledyne Isco, Lincoln, USA) automatic water samplers (locations indicated in Figure
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1). To avoid sample evaporation, paraffin was added to the autosampler bottles and the bottles
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were emptied in at least a fortnightly frequency. Soil water was extracted weekly until
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November 2013 from a peat soil (for location see Figure 1) at -10 and -30 cm depth by applying
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a vacuum to Rhizon soil moisture samplers (Rhizosphere Research Products, Wageningen,
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Netherlands) [Tetzlaff et al., 2014]. Suction lysimeters, like the Rhizon sampler, were shown to
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sample the mobile soil water in a comparison with other soil water isotope methods [Sprenger
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et al., 2015]. Two groundwater springs draining deeper groundwater stores (>5 m depth) were
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sampled fortnightly until May 2012 (GWa and GWb in Figure 1) [Birkel et al., 2011].
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The drainage network of channels and pools in the valley bottom of the BB was sampled at up
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to 94 locations on six sampling days between September 2011 and September 2013. The sample
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number varied between 35 and 94, since the sampling days cover different catchment wetness
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states.
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connected/disconnected peatland drainage network (green crosses in Figure 1) and more
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perennially connected waters of the riparian zone (purple crosses in Figure 1). In parallel to the
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spatially distributed sampling, stream water was sampled at 10 different locations along the
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perennial main stream (grey crosses in Figure 1).
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All samples were stored in air tight vials and kept refrigerated until they were analyzed for their
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water stable isotopic composition with a Los Gatos DLT-100 laser isotope analyzer (Los Gatos
183
Research, Inc., San Jose, USA). The precision of the measurements is given as ±0.1 ‰ for
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oxygen-18 ( 18O) and ±0.4 ‰ for deuterium ( 2H). The isotopic composition is given in the
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delta notation (in ‰) describing the relative difference in the ratio of heavy to light isotopes of
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a water sample to the Vienna Standard Mean Ocean Water (VSMOW).
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The isotopic composition of precipitation is characterized by equilibrium fractionation
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processes leading to a strong correlation between 18O and 2H in rainfall water (Global
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Meteoric Water Line, GMWL, Dansgaard [1964]). This relation is described locally by the
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local meteoric water line (LMWL) representing the regression line, characterized by a slope
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and a 2H-axis intercept, between 18O and 2H in a dual isotope plot (LMWL for the BB:
These
spatially
distributed
samples
6
were
divided
into
a
dynamically
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2H=7.7x18O+5.1). However, since the heavier 1H218O isotopologue (molecular weight =
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20.015 g mol-1, Horita et al. [2008]) is less likely to change from the liquid phase to the gaseous
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phase than the lighter 1H2H16O isotopologue (molecular weight = 19.017 g mol-1, Horita et al.
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[2008]), the ratio between 18O and 2H will change during evaporation processes, called
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kinetic fraction processes [Craig et al., 1963]. As a result, water samples that experienced
197
evaporation will plot below the LMWL. As more water evaporates, the remaining water from
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which the evaporation takes place gets more kinetically fractionated. These samples will
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increasingly deviate from the LMWL and their regression line in the dual-isotope plot will have
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a slope that is lower than the LMWL. This regression line is usually called evaporation water
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line (EWL). The resulting deviation of the water sample from the LMWL is described as the
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line-conditioned excess (lc-excess) as defined by Landwehr and Coplen [2006]:
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𝑙𝑐 − 𝑒𝑥𝑐𝑒𝑠𝑠 = 𝛿 2𝐻 − 𝑎 × 𝛿 18𝑂 − 𝑏
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with a and b representing the slope and intercept of the LMWL (for BB: a=7.7; b=5.1‰). While
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the widely applied deuterium-excess (d-excess, Dansgaard [1964]) describes the offset
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between the 2H of a water sample on the GMWL to a water sample in the dual-isotope space
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at a given 18O, the lc-excess uses the LMWL as reference. We derived for our water samples
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the lc-excess and their uncertainty was estimated based on the precision of the isotope analysis
209
and the slope of the LMWL to be ±1.17 ‰.
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We computed 14-days mean lc-excess of P (lc-excessP14) and daily PET (PET14days) to account
211
for the antecedent conditions of the water sampling.
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For description of the flow regime, seasonal coefficients of variations of the discharge (CVQ)
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were computed in accordance to Botter et al. [2013]. Differences between the lc-excess of the
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water samples taken from the peatland drainage network, stream, and riparian zone were tested
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with the Kruskal-Wallis rank sum test (kruskal.test in R) followed by a post-hoc test after
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Nemenyi (posthoc.kruskal.nemenyi.test in R). We tested the data for normality with the
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Shapiro-Wilk test. We applied the Spearman rank-order correlation (coefficient give as ρ) to
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test relationships that are not normally distributed (e.g., monthly Q, monthly P, monthly PET,
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Soil water lc-excess, distance to stream and altitude of sampling locations of the water samples
220
in the extended peatland drainage network). For data that was normally distributed, we applied
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the Pearson correlation (e.g., average lc-excess and SD of lc-excess in the extended peatland
222
drainage network, Q and PET14days for the days of sampling in the extended drainage network
223
(correlation coefficient given as r). The significance level was defined to be 0.05 for all
224
statistical tests including. When we defined meteorological seasons, we refer to spring (April,
(Eq. 1),
7
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March, May), summer (June, July, August), autumn (September, October, November), and
226
winter (December, January, February).
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4. Results 4.1. Hydroclimatic dynamics
228 229
As for most years in the BB, precipitation (P) was quite evenly distributed throughout the study
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period (2/3 of all days were rainy days) with no strong seasonality (Figure 2a). Annual P ranged
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between 930 (2015) and 1200 mm year-1 (2014). Median monthly P (± standard deviation (SD))
232
was 78 mm/month (±53). Precipitation occurred mainly at low intensities, with 90% of the rainy
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days having < 10 mm d-1 (summing up to 50% of the total precipitation amount) and 1/3 of the
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rainy days with less than 2 mm d-1 of P.
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Annual potential evapotranspiration (PET) ranged between 460 (in 2012) and 520 mm year-1
236
(in 2013). PET showed a strong seasonal variation with average PET rates of 2.3 to 2.7 mm d-
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1
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maximum PET rates were between 3.5 to 6.5 mm d-1 from May to August.
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Annual discharge (Q) varied between 440 (in 2015) and 900 mm year-1 (in 2014, Figure 2e).
240
Monthly sums of Q were positively correlated with monthly P (ρ = 0.54, p < 0.001) and
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negatively correlated with PET (ρ = -0.57, p < 0.001). On 90% and 40% of the days, Q was
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below 4 mm d-1 and 1 mm d-1, respectively. However, during P events, the BB responded
243
quickly and flow rates increased within few hours. Q was on average higher between November
244
and February with mean values between 2.2 and 3.3 mm d-1. From March to October, the
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average Q was between 1.7 and 0.8 mm d-1. With minimum Q that usually did not drop below
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1 mm d-1, stream flows remained higher during winter than the base flow rates from spring to
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autumn. Hence, the flow regime, as defined by Botter et al. [2013], was more erratic during
248
summer (CVQ = 1.31) and autumn (CVQ = 1.37), while more persistent and indifferent during
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winter (CVQ = 0.91) and spring (CVQ = 1.05), respectively. The variation of the seasonal CVQ
250
was positively correlated to mean seasonal P (ρ = 0.52, p = 0.027) and not correlated to the
251
mean seasonal PET (ρ = 0.24, p = 0.34).
252
from May to August and 0.3 to 0.7 mm d-1 from November to February (Figure 2b). However,
4.2. Stable isotope dynamics
253
The isotopic signal of P generally followed the seasonal variation of air temperature (T), with
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monthly lc-excess values being negatively correlated with T (ρ = -0.63, p < 0.001). The
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correlation of P lc-excess to PET was weaker (ρ = -0.55, p < 0.001). The lc-excess varied 8
256
throughout the year, with lower monthly weighted average values (ranging between -2.7 to -1.4
257
‰) between May and September and higher values (+0.9 to +2.7 ‰) between November and
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February (Figure 2a). However, the long term median of P lc-excess was 0.5 ‰ (Table 1,
259
Supplementary Figure 1). The local meteoric water line (LMWL; 2H=7.7x18O+5.1) was
260
close to the GMWL.
261
The lc-excess signal of the BB stream water followed the seasonal variation of the P input, but
262
its variation was only one third of the variation in the P lc-excess (see SD in Table 1). However,
263
the stream water lc-excess showed a lower correlation with PET (ρ = -0.48, p = 0.001) than P
264
lc-excess (ρ = 0.60, p < 0.001). The lc-excess at the BB outlet was on average more depleted in
265
heavy isotopes than the P input and most of the time greater than 0 (Table 1, Figure 2a,
266
Supplementary Figure 1). The regression between 18O and 2H (the evaporation water line,
267
EWL) had with 5.3 and -12.5 ‰ a lower slope and intercept than the LMWL.
268
The lc-excess of the mobile soil water sampled with the Rhizon samplers at -10 cm of the peat
269
soil followed the P input signal (Figure 2d) and was correlated to the lc-excessP14 (ρ = 0.45, p
270
< 0.001). However, the variation of the lc-excess in the topsoil was damped compared to P by
271
about 60% (Table 1). The mobile soil water was isotopically enriched from June to September
272
with average monthly lc-excess values ranging between +1.4 and +1.8 ‰ and more depleted
273
values (monthly averaged lc-excess > +3.1) from November to March. The correlation of the
274
monthly averages of the topsoil water lc-excess with PET had a Spearman's rank correlation
275
coefficient of ρ = -0.58 (p < 0.001), which was stronger than for the precipitation or the stream
276
water. The soil water isotopic signal at -30 cm was more damped than at -10 cm, since the
277
standard deviation (SD) was only 20% of the SD of the P lc-excess (Table 1). The correlation
278
between the monthly average lc-excess at -30 cm and the monthly average PET was relatively
279
low (ρ = -0.37, p = 0.04) and no influence of the lc-excessP14 was detected (ρ = 0.09, p = 0.53).
280
Generally, the soil water samples had mostly an lc-excess >0 (above or on the LMWL), but
281
dropped below 0 during the summer 2013 (Figure 2d). The EWL was for the mobile soil water
282
with 5.3 and -12.4 ‰ for slope and intercept, respectively, very similar to the stream water
283
EWL. The lc-excess of the mobile soil water in the peat soil generally followed the stream water
284
signal at -10 cm (ρ = 0.44, p < 0.001) and -30 cm (ρ = 0.32, p < 0.001).
285
The groundwater had the most depleted isotope signals with an average lc-excess of +4.5 ‰
286
and generally very little variation throughout the year (Table 1, Figure 2e). In contrast to the
287
precipitation, stream, and mobile soil water, the groundwater lc-excess showed no correlation
288
with the monthly mean T (for GWa: = -0.28, p= 0.40; for GWb: = 0.43, p= 0.19). All the 9
289
groundwater samples plotted above the LMWL in a dual isotope plot (lc-excess >0) and the
290
slope and intercept of the EWL was 4.2 and -22.4 ‰, respectively.
291
4.3. Spatio-temporal dynamics of evaporation fractionation
292
The spatially distributed sampling of surface waters from water tracks and pools in the valley
293
bottom of the BB showed that the water in the peatland drainage network deviated considerably
294
from the LMWL, while the water from the riparian zone and the BB stream stayed usually
295
above and generally close to the LMWL (Figure 3). The slope of the EWL for the peatland
296
drainage network samples ranged accordingly between 3.9 (in September 2011) and 5.4 (in
297
February 2013) (Table 2). The lower slopes of the EWLs show that the isotopic enrichment in
298
the peatland drainage network was most pronounced for the sampling campaigns between May
299
and September, while the winter samples followed the LMWL, indicating no or little
300
evaporation fractionation (Figure 3a). The stream water and riparian zone in BB followed the
301
temporal variation of the P input with more enriched isotopic signals during summer and
302
depleted values during winter (Figure 3b and c). The water samples of the riparian zone did not
303
differ significantly from the stream water samples in terms of their median lc-excess values at
304
any sampling date (indicated with lower case letters in Figure 4a). In contrast, the lc-excess of
305
the peatland drainage network was significantly different from the water in the stream and
306
riparian zone for half of the sampling campaigns (Sept. 2011, April 2012, and May 2013). For
307
the peatland drainage network, the variation of the lc-excess was lowest in February 2013,
308
where a significantly different lc-excess was found compared to the other sampling campaigns
309
(indicated by upper case letters in Figure 4b). The spring (in April 2012 and May 2013) and
310
summer/fall sampling (in September 2011, August 2012) of the peatland drainage network
311
formed two different groups that differed significantly from each other in terms of their lc-
312
excess (Figure 4b).
313
There was a pronounced spatial variability in the lc-excess pattern in the peatland drainage
314
network. The average lc-excess was lower (indicating higher evaporation fractionation) in the
315
center of the bog-fen complexes (green open symbols in Figure 5a). Towards the edges, the lc-
316
excess increased, which indicates increasing groundwater influence. In fact, as more the
317
sampling location was located within the area covered by peat soil (dark brown in Figure 1),
318
the higher was the evaporation signal (= -0.70, p < 0.01). Towards the minerogenic soils, the
319
groundwater seepage into the peatland drainage network increases. The sample locations
320
outside the peat soil cover had an almost constant average lc-excess signal (range between +2
321
and +4 ‰) and did not show a dependency on the distance to the peat soils (= -0.18, p = 0.40). 10
322
Both altitude and distance to the stream of the sampling locations showed a significant relation
323
to the average lc-excess of the individual sampling locations, but they were relatively weak
324
predictors of the lc-excess pattern (= -0.23 and = -0.36, respectively). Thus, we conclude
325
that in our environment the pedogenic landscape characteristics are more important than
326
topographic conditions.
327
The variability of the lc-excess in space, given as the SD over all six sampling campaigns, was
328
highly correlated with the mean lc-excess (=-0.76, p
