Evaporation fractionation in a peatland drainage

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fractionation in a peatland drainage network and its effect on stream water ... composition is characterized for surface waters by the local meteoric water line ...
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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

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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

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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

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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.

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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.,

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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

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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.

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The vegetation on the peatland drainage network is dominated by Spagnum spp. mosses along

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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

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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

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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

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were

divided

into

a

dynamically

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2H=7.7x18O+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

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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

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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

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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

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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

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drainage network, Q and PET14days for the days of sampling in the extended drainage network

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(correlation coefficient given as r). The significance level was defined to be 0.05 for all

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statistical tests including. When we defined meteorological seasons, we refer to spring (April,

(Eq. 1),

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March, May), summer (June, July, August), autumn (September, October, November), and

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winter (December, January, February).

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4. Results 4.1. Hydroclimatic dynamics

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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))

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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

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(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).

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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

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quickly and flow rates increased within few hours. Q was on average higher between November

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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

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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

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was positively correlated to mean seasonal P (ρ = 0.52, p = 0.027) and not correlated to the

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mean seasonal PET (ρ = 0.24, p = 0.34).

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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

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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

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throughout the year, with lower monthly weighted average values (ranging between -2.7 to -1.4

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‰) 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,

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Supplementary Figure 1). The local meteoric water line (LMWL; 2H=7.7x18O+5.1) was

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close to the GMWL.

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The lc-excess signal of the BB stream water followed the seasonal variation of the P input, but

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its variation was only one third of the variation in the P lc-excess (see SD in Table 1). However,

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the stream water lc-excess showed a lower correlation with PET (ρ = -0.48, p = 0.001) than P

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lc-excess (ρ = 0.60, p < 0.001). The lc-excess at the BB outlet was on average more depleted in

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heavy isotopes than the P input and most of the time greater than 0 (Table 1, Figure 2a,

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Supplementary Figure 1). The regression between 18O and 2H (the evaporation water line,

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EWL) had with 5.3 and -12.5 ‰ a lower slope and intercept than the LMWL.

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The lc-excess of the mobile soil water sampled with the Rhizon samplers at -10 cm of the peat

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soil followed the P input signal (Figure 2d) and was correlated to the lc-excessP14 (ρ = 0.45, p

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< 0.001). However, the variation of the lc-excess in the topsoil was damped compared to P by

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about 60% (Table 1). The mobile soil water was isotopically enriched from June to September

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with average monthly lc-excess values ranging between +1.4 and +1.8 ‰ and more depleted

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values (monthly averaged lc-excess > +3.1) from November to March. The correlation of the

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monthly averages of the topsoil water lc-excess with PET had a Spearman's rank correlation

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coefficient of ρ = -0.58 (p < 0.001), which was stronger than for the precipitation or the stream

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water. The soil water isotopic signal at -30 cm was more damped than at -10 cm, since the

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standard deviation (SD) was only 20% of the SD of the P lc-excess (Table 1). The correlation

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between the monthly average lc-excess at -30 cm and the monthly average PET was relatively

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low (ρ = -0.37, p = 0.04) and no influence of the lc-excessP14 was detected (ρ = 0.09, p = 0.53).

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Generally, the soil water samples had mostly an lc-excess >0 (above or on the LMWL), but

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dropped below 0 during the summer 2013 (Figure 2d). The EWL was for the mobile soil water

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with 5.3 and -12.4 ‰ for slope and intercept, respectively, very similar to the stream water

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EWL. The lc-excess of the mobile soil water in the peat soil generally followed the stream water

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signal at -10 cm (ρ = 0.44, p < 0.001) and -30 cm (ρ = 0.32, p < 0.001).

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The groundwater had the most depleted isotope signals with an average lc-excess of +4.5 ‰

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and generally very little variation throughout the year (Table 1, Figure 2e). In contrast to the

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precipitation, stream, and mobile soil water, the groundwater lc-excess showed no correlation

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with the monthly mean T (for GWa:  = -0.28, p= 0.40; for GWb:  = 0.43, p= 0.19). All the 9

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groundwater samples plotted above the LMWL in a dual isotope plot (lc-excess >0) and the

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slope and intercept of the EWL was 4.2 and -22.4 ‰, respectively.

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4.3. Spatio-temporal dynamics of evaporation fractionation

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The spatially distributed sampling of surface waters from water tracks and pools in the valley

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bottom of the BB showed that the water in the peatland drainage network deviated considerably

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from the LMWL, while the water from the riparian zone and the BB stream stayed usually

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above and generally close to the LMWL (Figure 3). The slope of the EWL for the peatland

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drainage network samples ranged accordingly between 3.9 (in September 2011) and 5.4 (in

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February 2013) (Table 2). The lower slopes of the EWLs show that the isotopic enrichment in

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the peatland drainage network was most pronounced for the sampling campaigns between May

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and September, while the winter samples followed the LMWL, indicating no or little

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evaporation fractionation (Figure 3a). The stream water and riparian zone in BB followed the

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temporal variation of the P input with more enriched isotopic signals during summer and

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depleted values during winter (Figure 3b and c). The water samples of the riparian zone did not

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differ significantly from the stream water samples in terms of their median lc-excess values at

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any sampling date (indicated with lower case letters in Figure 4a). In contrast, the lc-excess of

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the peatland drainage network was significantly different from the water in the stream and

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riparian zone for half of the sampling campaigns (Sept. 2011, April 2012, and May 2013). For

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the peatland drainage network, the variation of the lc-excess was lowest in February 2013,

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where a significantly different lc-excess was found compared to the other sampling campaigns

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(indicated by upper case letters in Figure 4b). The spring (in April 2012 and May 2013) and

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summer/fall sampling (in September 2011, August 2012) of the peatland drainage network

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formed two different groups that differed significantly from each other in terms of their lc-

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excess (Figure 4b).

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There was a pronounced spatial variability in the lc-excess pattern in the peatland drainage

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network. The average lc-excess was lower (indicating higher evaporation fractionation) in the

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center of the bog-fen complexes (green open symbols in Figure 5a). Towards the edges, the lc-

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excess increased, which indicates increasing groundwater influence. In fact, as more the

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sampling location was located within the area covered by peat soil (dark brown in Figure 1),

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the higher was the evaporation signal (= -0.70, p < 0.01). Towards the minerogenic soils, the

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groundwater seepage into the peatland drainage network increases. The sample locations

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outside the peat soil cover had an almost constant average lc-excess signal (range between +2

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and +4 ‰) and did not show a dependency on the distance to the peat soils (= -0.18, p = 0.40). 10

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Both altitude and distance to the stream of the sampling locations showed a significant relation

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to the average lc-excess of the individual sampling locations, but they were relatively weak

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predictors of the lc-excess pattern (= -0.23 and = -0.36, respectively). Thus, we conclude

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that in our environment the pedogenic landscape characteristics are more important than

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topographic conditions.

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The variability of the lc-excess in space, given as the SD over all six sampling campaigns, was

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highly correlated with the mean lc-excess (=-0.76, p