Lateral carbon flux through river networks is an
important and poorly understood component of the global carbon budget. This
work investigates how temperature and hydrology control the production and
export of dissolved organic carbon (DOC) in the Susquehanna Shale Hills
Critical Zone Observatory in Pennsylvania, USA. Using field measurements of
daily stream discharge, evapotranspiration, and stream DOC concentration, we
calibrated the catchment-scale biogeochemical reactive transport model
BioRT-Flux-PIHM (Biogeochemical Reactive
Transport–Flux–Penn State Integrated Hydrologic Model, BFP), which met the satisfactory standard of a Nash–Sutcliffe
efficiency (NSE) value greater than 0.5. We used the calibrated model to estimate
and compare the daily DOC production rates (Rp; the sum of the local DOC
production rates in individual grid cells) and export rate (Re; the
product of the concentration and discharge at the stream outlet, or load).
Results showed that daily Rp varied by less than an order of magnitude, primarily depending on seasonal temperature. In contrast, daily Re
varied by more than 3 orders of magnitude and was strongly associated with
variation in discharge and hydrological connectivity. In summer, high
temperature and evapotranspiration dried and disconnected hillslopes from
the stream, driving Rp to its maximum but Re to its minimum. During
this period, the stream only exported DOC from the organic-poor groundwater
and from organic-rich soil water in the swales bordering the stream.
The DOC produced accumulated in hillslopes and was later flushed out during the
wet and cold period (winter and spring) when Re peaked as the stream
reconnected with uphill and Rp reached its minimum.
The model reproduced the observed concentration–discharge (C–Q) relationship
characterized by an unusual flushing–dilution pattern with maximum
concentrations at intermediate discharge, indicating three end-members of source waters. A sensitivity analysis indicated
that this nonlinearity was caused by shifts in the relative contribution of
different source waters to the stream under different flow conditions. At low
discharge, stream water reflected the chemistry of organic-poor groundwater;
at intermediate discharge, stream water was dominated by the organic-rich
soil water from swales; at high discharge, the stream reflected uphill soil
water with an intermediate DOC concentration. This pattern persisted regardless
of the DOC production rate as long as the contribution of deeper groundwater
flow remained low (<18 % of the streamflow). When groundwater
flow increased above 18 %, comparable amounts of
groundwater and swale soil water mixed in the stream and masked the high DOC concentration from swales. In that case, the C–Q patterns switched to a
flushing-only pattern with increasing DOC concentration at high discharge.
These results depict a conceptual model that the catchment serves as a
producer and storage reservoir for DOC under hot and dry conditions and
transitions into a DOC exporter under wet and cold conditions. This study
also illustrates how different controls on DOC production and export –
temperature and hydrological flow paths, respectively – can create temporal
asynchrony at the catchment scale. Future warming and increasing
hydrological extremes could accentuate this asynchrony, with DOC production
occurring primarily during dry periods and lateral export of DOC dominating
in major storm events.
Introduction
Soil organic carbon (SOC) is the largest terrestrial stock of organic
carbon, containing approximately 4 times more carbon than the atmosphere
(Stockmann et al., 2013; Hugelius et al., 2014). Understanding the SOC
balance requires the consideration of lateral fluxes in water, including
dissolved organic and inorganic carbon (DOC and DIC, respectively), and vertical fluxes of
gases such as CO2 and CH4 (Chapin et
al., 2006). Both lateral and vertical fluxes influence SOC mineralization to
the atmosphere (Campeau et al., 2019), although lateral
fluxes are arguably less understood and integrated into Earth system models (Aufdenkampe et al., 2011; Raymond et al., 2016). Lateral fluxes from
terrestrial to aquatic ecosystems are similar in magnitude to net vertical
fluxes (Regnier et al., 2013; Battin et al.,
2009), highlighting the importance of quantifying the controls on the lateral
carbon (C) flux. In addition to its role in the global C cycle, DOC is an
important water quality parameter that can mobilize metals and contaminants
as well as imposing challenges for water treatment (Sadiq and Rodriguez, 2004; Bolan et al., 2011). DOC also regulates
food web structures by acting as an energy source for heterotrophic
organisms and interacts with other biogeochemical cycles (Malone et al.,
2018; Abbott et al., 2016).
SOC decomposition and DOC production have been studied extensively
(Abbott et al., 2015; Hale et al., 2015; Humbert et
al., 2015; Lambert et al., 2013; Neff and Asner, 2001), yet the interactions
between SOC and DOC and their response to climate change at catchments or
larger scales remain unresolved (Laudon et al., 2012; Clark et al., 2010). Some
regions have experienced long-term increases in DOC, potentially due to
recovery from acid rain or climate-induced changes in temperature (T) and
hydrological flow (Laudon et al., 2012; Perdrial et al., 2014; Evans et
al., 2012; Monteith et al., 2007), whereas others have observed decreases or no
change (Skjelkvale et al., 2005; Worrall et al., 2018). Generally, the
linkages among SOC processing, hydrological conditions, and DOC export or
concentration remain poorly understood. Recent analyses indicate that the
relationship between DOC concentration and discharge (C–Q) at stream outlets
is primarily positive (Moatar et al., 2017; Zarnetske et al., 2018).
Approximately 80 % of watersheds in the USA and France show a flushing
C–Q pattern (i.e., the stream DOC concentration increases with discharge), whereas
the rest shows dilution (decreasing DOC with discharge) or chemostatic
behavior (negligible concentration change with discharge). These C–Q patterns
generally correlate with catchment characteristics, including topography,
wetland area, and climate characteristics, but it remains uncertain how hydrological and biogeochemical
processes regulate SOC decomposition, DOC production, and DOC export
(Jennings et al., 2010; Worrall et al., 2018). This gap in process
understanding limits the integration of lateral carbon dynamics into
projections of future ecosystem response to environmental change.
Attributes of the Susquehanna Shale Hills Critical Zone
Observatory (SSHCZO): (a) surface elevation, (b) soil depth, and (c) soil
organic carbon (SOC). The surface elevation was generated from lidar topographic
data (criticalzone.org/shale-hills/data), whereas soil depths and SOC were
interpolated using ordinary kriging based on field surveys with 77 and 56
sampling locations, respectively (Andrews et al., 2011; Lin, 2006). The
SOC distribution in panel (c) is further simplified using the high, uniform
SOC (5 % v/v) in swales and valley soils based on field survey information
(Andrews et al., 2011). Swales and valley floor areas were defined
based on surface elevation via field survey and a 10 m resolution
digital elevation model (Lin, 2006). Additional sampling
instrumentation is shown in panel (b), including six soil water sites (circles)
and three soil Tsites (squares).
Stream DOC can be influenced by a variety of factors that control SOC
decomposition and DOC production rates. DOC production generally
increases as T increases; however, there may be multiple thermal optima, and the
local rates can vary with SOC characteristics, soil type, and soil biota
(Davidson and Janssens, 2006; Jarvis and Linder, 2000; Yan et al., 2018;
Zarnetske et al., 2018). DOC production rates can exhibit low temperature
sensitivity in highly weathered soils with a high clay content (Davidson and Janssens, 2006). They have also been shown to increase with
soil water content in sandy loam soils (Yuste et al.,
2007) and to have an optimum with a volumetric water content of approximately
0.75 in fine sands (Skopp et al., 1990). Because DOC export (or load) is the product of discharge and DOC concentration, it
may differ from local DOC production rates in complex ways. For example,
high T can produce a peak soil water DOC concentration but not necessarily
stream concentration or export, due to temporal or spatial mismatches (D'Amore et al., 2015). These confounding factors present
significant challenges to quantify the predominant mechanisms that regulate
DOC production and export under varying environmental conditions.
One approach to understanding DOC production and export is the use of reactive
transport models (RTM). These models integrate multiple production,
consumption, and export processes, enabling the differentiation of individual
and coupled processes (Steefel et al., 2015; Li, 2019; Li et al., 2017b). The use of RTMs
complements statistical tools for the identification of influential
factors (Correll et al., 2001; Herndon et al., 2015; Zarnetske et al.,
2018). Historically, RTMs have been used in groundwater systems, where
direct observations are particularly challenging (Kolbe et al., 2019; Li
et al., 2009; Wen and Li, 2018; Wen et al., 2018). At the catchment scale,
biogeochemical modules have been developed as add-ons to hydrological
models. For example, a DOC production module was coupled to the HBV
hydrological model using a static SOC pool that emphasized the influence of
active-layer dynamics and slope aspect (Lessels et al., 2015). The INCA-C
(Futter et al., 2007) and extended LPJ-GUESS (Tang
et al., 2018) models have investigated the importance of land cover in
determining DOC terrestrial routing and lateral transport. Terrestrial and
aquatic carbon processes have also been integrated into the Soil and Water
Assessment Tool (SWAT) to simulate aquatic DOC dynamics (Du et al., 2019). These modules typically simulate
individual reactions without considering multicomponent reaction thermodynamics and
kinetics.
In this context, the recently developed BioRT-Flux-PIHM model (BFP, Biogeochemical Reactive
Transport–Flux–Penn State Integrated Hydrologic Model) fills an important gap by incorporating coupled elemental
cycling, stoichiometry, and rigorous thermodynamics and kinetics (Bao et
al., 2017; Zhi et al., 2019). We used the BFP to address the following question: how do hydrology and T interact to determine rates of DOC production and export at the catchment scale? We
applied the BFP to a temperate forest catchment in the Susquehanna Shale
Hills Critical Zone Observatory (SSHCZO). This small
catchment (<0.1 km2) has gentle topography with a network of
shallow depressions or swales that have high SOC and deep soils (detailed in
Sect. 2). It is underlain by one type of lithology (shale) and land use
(forest), providing a useful test bed to evaluate biogeochemical and
hydrological functions (Brantley et al.,
2018). Previous lab and field work have identified non-chemostatic C–Q
patterns of DOC at SSHCZO that are attributable to differences in the
hydrologic connectivity of organic-rich soils during different flow
conditions (Andrews et al., 2011; Herndon et al., 2015). SSHCZO has
spatially extensive and high-frequency measurements of soil properties,
hydrology, and biogeochemistry (Brantley et
al., 2018). These data facilitate detailed benchmarking of the BFP model and
evaluation of processes controlling DOC production and export. We expected
that T and soil moisture would drive DOC production in the soil, whereas DOC
export and, thus, C–Q patterns would be most related to hydrological
connectivity. Therefore, we predicted that DOC production and export might
be asynchronous (i.e., not occurring at the same time) because they respond
differently to changes in T and hydrology. Although soil respiration is an
important process, this study focuses on the net production and export of
DOC.
MethodsStudy site: a small catchment with an intermittent stream
The Shale Hills catchment is a 0.08 km2, V-shaped, first-order
watershed with an intermittent stream in central Pennsylvania. It is
forested with coniferous trees and is situated on the Rose Hill shale
formation. The annual mean air T is 9.8±1.9∘C (± SD)
and the annual mean precipitation is 1029±270 mm over the past
decade. The watershed is characterized by large areas of swales and valley
floors with deep and wet soils (Fig. 1b). These lowland soils contain more
SOC (∼5 % v/v) than the hillslopes and uplands
(∼1 % v/v; Fig. 1c).
A schematic representation of major processes in the catchment
reactive transport model BFP (BioRT-Flux-PIHM). Stream discharge Q includes
surface runoff QS, soil water interflow (lateral flow) QL, and
groundwater flow QG. In the vertical direction, soil pores are not
saturated with water in the shallow unsaturated zone and water flows
vertically until it reaches the saturated zone where water forms interflows
and moves laterally to the stream. Soil water total storage ST is the
sum of water in the unsaturated (Su) and saturated zones (Ss). Some
water also recharges further into deeper groundwater. Within the soil zone,
SOC decomposes and releases DOC, which also sorbs onto the soil surface to
become ≡XDOC.
Soil water DOC samples were collected using lysimeters with a diameter of 5 cm installed at 10 or 20 cm intervals from the soil surface to a depth of
hand-auger refusal, which varied from 30 to 160 cm depending on soil
thickness. There were a total of six sampling locations (Fig. 1b),
including three at the south planar sites – valley floor (SPVF), mid-slope
(SPMS), and ridgetop (SPRT) – and three at the swale sites – valley floor
(SSVF), mid-slope (SSMS), and ridgetop (SSRT). No soil water DOC samples were
collected on the north side of the catchment. Stream water DOC samples were
collected daily in glass bottles at the stream outlet weir. All soil water
and stream water DOC samples were filtered to 0.45 µm using Nylon
syringe filters and were analyzed with a Shimadzu TOC-5000A analyzer
(detailed in Andrews et al., 2011). Real-time soil T (every 10 min) was measured at the ridgetop, mid-slope, and valley floor (squares in
Fig. 1b) using automatic monitoring stations at depths of ∼0.10, 0.20, 0.40, 0.70, 0.90, 1.00, and 1.30 m (Lin and Zhou,
2008).
The BioRT-Flux-PIHM (BFP) model
BFP is the catchment reactive transport model of the general PIHM (Penn
State Integrated Hydrologic Modeling System) family of code (Duffy et
al., 2014). The code includes three modules (Fig. 2): the surface
hydrological module PIHM, the land surface module Flux, and the
multicomponent reactive transport module BioRT (Biogeochemical Reactive
Transport). The code has been applied to simulate conservative solute
transport, chemical weathering, surface complexation, and biogeochemical
reactions at the catchment scale (Bao et al., 2017; Zhi et al., 2019; Li,
2019). Here, we only introduce the salient features that are relevant to this
study; readers are referred to earlier publications for further details.
Flux-PIHM separates the subsurface flow into active interflow in shallow
soil zones and groundwater flow deeper than the soil-weathered rock
interface. Note that this “deeper groundwater” is the groundwater that
actively interacts with the stream and shallow layer, not necessarily the
water in the deep groundwater aquifer. The PIHM module simulates
hydrological processes including precipitation, infiltration, surface runoff
QS, soil water interflow (lateral flow) QL, and discharge Q (Fig. 2). The Flux module simulates processes including solar radiation and
evapotranspiration. Flux-PIHM calculates water variables (e.g., water
storage, soil moisture, and water table depth) in unsaturated and saturated
zones and assumes a no-flow boundary at the soil–bedrock interface with high
permeability contrast. In this version of Flux-PIHM, the deeper groundwater
flow QG is a separate input to the stream and is decoupled from the
shallow soil water. This is supported by field data that show negligible
seasonal variation in groundwater chemistry (Jin et al., 2014; Thomas et
al., 2013; Kim et al., 2018). The QG is estimated using conductivity
mass-balance hydrograph separation (Lim et al., 2005).
The BioRT module takes in water calculated at each time step to simulate
reactive transport processes. BFM discretizes the domain into prismatic
elements and uses a finite volume approach based on mass conservation. The mass conservation
governing equation for the reactive transport of a single solute m is as
follows:
Vid(Sw,iθiCm,i)dt=∑j=Ni,1Ni,xAijDijCm,j-Cm,ilij-qijCm,j+rm,i,m=1,np,
where i and j represent the grid block i and the neighboring grid j; the
subscript x distinguishes between flow in the unsaturated zone (infiltration
and recharge) and saturated zone (recharge and lateral flow); V is the total
bulk volume (m3) of each grid block; Sw is the soil moisture
(m3 water m-3 pore volume); θ is porosity; C is the aqueous
species concentration (mol m-3 water); t is time (s); N is the index of
elements sharing surfaces; A is the grid interface area (m2); D is the
diffusion/dispersion coefficient (m2 s-1); l is the distance (m) between
the center of two neighboring grid blocks; q is the flow rate (m3 s-1);
rm is the kinetically controlled reaction rates (mol s-1) involving
species m, which is the DOC production rate from SOC decomposition at the
grid block i; and np is the total number of independent solutes.
DOC production and sorption
In the model, DOC is produced
by the decomposition of SOC via the kinetically controlled reaction
SOC(s)→DOC. With abundant SOC and O2 in soils serving as electron
donors and acceptors, a typical dual Monod kinetics can be simplified into
zero-order kinetics with additional temperature and soil moisture
dependence:
rp=kAfTfSw,
where rp is the local DOC production rate in individual grids (rm in
Eq. 1, m is DOC); k is the kinetic rate constant of net DOC production
(10-10 mol m-2 s-1) (Zhi et al., 2019; Wieder et al., 2014);
and A is a lumped “surface area” (m2, (2.5×10-3 m2 g-1) × (g of SOC mass)) that quantifies SOC content and biomass
abundance (Chiou et al., 1990; Kaiser and Guggenberger, 2003; Zhi et al.,
2019). The functions f(T) and f(Sw) describe the rate
dependence on soil T and moisture, respectively. f(T) follows a widely used
Q10-based formation: f(T)=Q10T-10/10, where Q10 quantifies the rate increases with T, with the
superscript 10 referring to a T value of 10 ∘C (Davidson and Janssens, 2006). Q10 in the base case is set at
2.0, within the typical range of 1.2–3.8 for forest ecosystems (Liu et
al., 2017). The f(Sw) has the form f(Sw)=(Sw)n
in the base case, where n is the saturation exponent with a value of 1.0,
which is within the typical range of 0.75–3.0 for most soils (Yan et al.,
2018; Hamamoto et al., 2010). The dependence of production rates on soil T
and moisture have been described using multiple forms in existing studies
(Davidson and Janssens, 2006; Yan et al., 2018) and will be further
explored via sensitivity analysis, as detailed in Sect. 2.6. The SOC
content typically decreases with depth (Billings et al., 2018; Bishop et al.,
2004), although the specific pattern may vary with soil texture, landscape
position, vegetation, and climate (Jobbagy and Jackson, 2000).
The depth function of SOC at Shale Hills has been observed to be exponential
(Andrews et al., 2011), which is typical of many soils
(Billings et al., 2018; Currie et al., 1996). To take this into account,
we use the equation Cd(z)=C0exp-zbm, where Cd is SOC at depth z below the surface; C0
is the SOC level at the ground surface, and bm quantified the decline
rate with depth, which is set here to a value of 0.3 (Weiler and
McDonnell, 2006).
DOC produced from SOC can also sorb on soils via the following reaction: ≡X+DOC↔≡XDOC, where ≡X and ≡XDOC
represent the functional group without and with sorbed DOC, respectively (Rasmussen et al., 2018). This reaction is considered
fast and is thermodynamically controlled with an equilibrium constant
Keq that links the activity (here approximated by concentrations) of the
three chemicals via Keq=≡XDOC≡XDOC. The DOC concentrations calculated from Eq. (1) were used to establish the concentrations of ≡X and ≡XDOC. The Keq value represents the thermodynamic limit of the sorption,
i.e., the sorption affinity of the soil for DOC. It depends on temperature
but also on soil properties such as the clay content and the abundance of iron
oxides (Kaiser et al., 2001; Conant et al., 2011). A Keq value of
100.2 was obtained by fitting the stream and soil water DOC data
(detailed in Sect. 2.4). The sum of [≡X] and [≡XDOC]
represents the soil sorption capacity. A value ranging from 4.0×10-5 to 6.0×10-5 mol g-1 soil was used for Shale Hills
(Jin et al., 2010; Li et al., 2017a) depending on the mineralogy in
different zones of the catchment.
Domain setup
BFP is a model with full discretization in the horizontal direction and
partial discretization in the vertical direction with three layers: ground
surface, unsaturated, and saturated zones. Although a new version of BFP
explicitly includes a groundwater zone, it was not released in time for this
work, so the groundwater fluxes were estimated separately. The study
watershed was discretized into 535 prismatic land elements and 20 stream
segments using PIHMgis (http://www.pihm.psu.edu/pihmgis_home.html, last access: 11 February 2020), a GIS interface for BFP. The land elements are unstructured
triangles with mesh sizes varying from 10 to 100 m. The simulation domain
was set up using national datasets: the USGS National Elevation
Dataset for topography; the National Land cover Database for vegetation distribution; the National Hydrography Dataset for water drainage network; the North American Land Data Assimilation Systems Phase 2 (NLDAS-2) for hourly meteorological forcing; and the Moderate Resolution
Imaging Spectroradiometer (MODIS) for leaf area index. In addition,
extensive characterization and measurement data at Shale Hills were
used to define soil depth and soil mineralogical properties such as surface
area and ion exchange capacity that are heterogeneously distributed across
the catchment (Andrews et al., 2011; Lin, 2006; Jin and Brantley, 2011;
Jin et al., 2010; Shi et al., 2013) (http://criticalzone.org/shale-hills/data/, last access: 11 February 2020).
Other soil matrix properties include conductivity, porosity, and van
Genuchten parameters. Soil macropores such as cracks, fractures, and roots
can generate preferential flows. Their properties are represented using the
area macropore fraction, depth, and conductivities. They are parameterized
based on values quantified in previous studies at Shale Hills (Shi et
al., 2013; Lin, 2006), as shown in Fig. S1 and Table S1 in the Supplement.
Based on field measurements, the SOC content in swales and valley areas is relatively
high (Andrews et al., 2011) and was set at 5 % (v/v solid
phase) compared with 1 % in the rest of the catchment (Fig. 1c). The clay
minerals were set at 23 % (v/v solid phase) along the ridgetop and 33 %
at the valley floor (Jin et al., 2010; Li et al., 2017a). The input DOC
concentrations in rainfall and groundwater (below soils) were set at reported
medians of 0.6 and 1.2 mg L-1, respectively (Andrews et al., 2011;
Iavorivska et al., 2016), as high-frequency DOC observations were not
available. The initial DOC concentration in soil water was set at 2.0 mg L-1, which was
the average concentration from the six field sampling locations in Fig. 1.
Model calibration
We used stream (daily) and soil pore water (biweekly) DOC concentration data
from April to October 2009 for model calibration and the year 2008 as spin-up
until a “steady state” for both water and DOC was reached. The “steady
state” here refers to a state where the inter-annual difference between
stored mass within the catchment is less than 5 % of the total mass. The
water input is precipitation, and its output is ET and discharge. The DOC
mass input is from rainfall, groundwater, and production, and the DOC output
is the export load at the stream outlet. The model performance was evaluated
using the monthly Nash–Sutcliffe efficiency (NSE) (Nash and
Sutcliffe, 1970) that quantified the residual variance of modeling output
compared to measurements. The general satisfactory range for monthly average
outputs for hydrological models is NSE >0.5
(Moriasi et al., 2007), and we used similar standards for
biogeochemical solutes (Li et al., 2017a). To reproduce the DOC
data, we first set the SOC surface area A using a literature range of
10-3–100 m2 g-1 (Zhi et al., 2019; Chiou et al., 1990;
Kaiser and Guggenberger, 2003). We also set Keq using a literature range
of 100–101 (Oren and Chefetz, 2012; Ling et al., 2006). Once
the simulated output captured the temporal trend of data, we refined
QG based on the estimation from hydrograph separation (Fig. S2) to
capture the peaks of stream DOC concentration, especially under low-discharge periods. Because not all soils are in contact with water, the
calibrated surface area represents the effective solid–water contact area,
and is orders of magnitude lower than the reported SOC surface areas from
laboratory experiments (Kaiser and Guggenberger, 2003). The
calibrated hydrological parameters are mostly from Shi et al. (2013), except
groundwater estimation. Groundwater estimates were based on Li et al. (2017a)
and further refined using conductivity mass-balance hydrograph separation
(Lim et al., 2005) and then by reproducing the stream
DOC concentration. In other words, stream and groundwater chemistry data
together helped constrain the groundwater flow.
Temporal dynamics of (a) daily precipitation, stream discharge Q, and evapotranspiration ET on the arithmetic scale; (b) stream discharge Q, soil water interflow QL, and groundwater QG on a logarithmic scale with soil T on an arithmetic axis (right); (c) soil water storage
ST (unsaturated water storage Su+ saturated water storage
Ss) and hydrological connectivity Ics/Width. The yellow dots in panel (b)
represent the average soil T from three sampling locations (square symbols in
Fig. 1b) with the shading reflecting variation in measurement. Q was highly
responsive to intense precipitation events in spring and winter. Note high
soil T, high ET, low Ss, and lowIcs/Width during July–August 2009. Stream
discharge was primarily comprised of QL, except in July–October when the
relative contribution of QG increased.
Quantification of water and DOC dynamicsHydrological connectivity
Saturated soil water storage calculated
from the model was used to quantify hydrological connectivity
Ics/Width. With “Width” defined as the average width of catchment in the direction
perpendicular to the stream (230 m), the term Ics/Width quantifies the
average proportional width of the catchment connected to the stream (e.g.,
Ics/Width=0.10, 0.35, and 0.70 in Fig. S3). Depending on the catchment
geometry and the extent of connectivity, Ics/Width may vary from 0 to 1.0. A high
Ics/Width value (i.e., high hydrological connectivity) indicates that a large
catchment area is connected to the stream. To determine whether two grids
are hydrologically connected, the spatial distribution of saturated water
storage was used to calculate connectivity following the equation
Ics=∫0∞τ(h)dh and an algorithm in the
literature (Allard, 1994; Western et al., 2001; Xiao et al., 2019). Here
τ(h) is the probability of two grid blocks being connected at a
separation distance of h. Two grids are considered “connected” if they are
joined by a continuous flow path and have saturated storages exceeding the
threshold of the 75th percentile of saturated storage (over the whole
catchment). Note that Ics/Width here only quantifies the hydrological
connectivity in soils and does not reflect the groundwater in shallow
aquifers below the soil–bedrock interface.
DOC concentration–discharge relationships
At the catchment scale, we differentiate
the DOC production rates and export rates. The production rate Rp is the
sum of the local DOC production rate rp in individual grid blocks (Eq. 2) across the whole catchment. The export rate Re is the product of
discharge and the DOC concentration at the stream outlet. Total stored DOC is
the difference between stream output and input from production, rainfall,
and groundwater. The DOC input from the rainfall Rr (mg d-1) is the
precipitation rate (m d-1) times the rainfall DOC concentration (6.0×10-4 mg m-3= 0.6 mg L-1×10-3 L m-3) and the
catchment drainage area (m2). The DOC input from groundwater Rg
(mg d-1) is the total groundwater influx (flow rate) times the groundwater DOC
concentration (1.2 mg L-1).
C–Q patterns were quantified using two complementary approaches: the power
law equation C=aQb (Godsey et al., 2009) and the ratio of the
coefficients of variation of the DOC concentration and discharge
CV[DOC]CVQ (Musolff et al., 2015). The slope of
the power law equation b does not account for the goodness of fit of the C–Q
pattern itself. For example, a slope of b=0 would be considered
chemostatic (i.e., relatively small variation of concentration compared with
discharge), although high variability in solute concentrations would in fact
reflect a chemodynamic behavior (i.e., solute concentrations are sensitive
to changes in discharge). We considered two
general categories based on these metrics (Musolff et al., 2015): if b values fell between -0.2 and 0.2 and
CV[DOC]CVQ≪1, C–Q patterns were
considered chemostatic; values of |b|>0.2 or
CV[DOC]CVQ≥1, indicated a chemodynamic behavior.
In the chemodynamic category, values of b>0.2 indicate flushing,
whereas values of b<-0.2 indicate dilution. We used the MATLAB
curve-fitting toolbox to obtain the best fit model parameters.
Sensitivity analysis
We used a sensitivity analysis to explore the influence of soil T and
moisture in the DOC production kinetics. The Q10 in f(T)=Q10T-10/10 was explored using a minimum value
of 1.0 (i.e., no dependence on T) and a maximum value of 4.0 (Davidson
and Janssens, 2006) (Fig. S4a), i.e., f(T)=1 and f(T)=4T-10/10. The rate dependence on soil moisture
was explored using the base case f1(Sw)=(Sw)n
(increase behavior), and three additional functions (f2, f3, and
f4) representing the most commonly observed forms (Fig. S4b),
including decrease behavior, constant behavior, and threshold behavior
(Gomez et al., 2012; Yan et al., 2018):
the decrease behavior function was
f2Sw=1-Sw0.60.77,
the constant behavior function was
f3Sw=0.65,
and the threshold behavior function was
f4Sw=Sw0.71.5Sw≤0.71-Sw1-0.71.5Sw>0.7.
The constants in Eqs. (3)–(5) were selected to ensure similar averages of
f(Sw) across the whole Sw range such that trajectories rather than
absolute values of f(Sw) were compared (Fig. S4b). The sensitivity of DOC
sorption onto soils was tested using Keq values of 0 (no sorption),
100.5 and 101.0.
The sensitivity of C–Q patterns and Re to changes in groundwater was
also tested with groundwater flow contribution and DOC concentration. The
groundwater flow rates were varied from negligible (QG=0) to
2.5 times those of the base case (QG=3.3×10-4 and 1.0×10-4 m d-1 for the wet and dry periods,
respectively). The corresponding fractions (QG/Q) of groundwater flow to
the total annual discharge for the two cases were 0 % and 18.8 %,
respectively. The groundwater DOC concentration (DOCGW) was varied by
2 orders of magnitude (0.12 and 12.0 mg L-1). Results from these
analyses were compared with the base case, in which the groundwater
contributed to 7.5 % of the total annual streamflow at 1.2 mg L-1.
(a) Temporal dynamics of measured and simulated stream DOC
concentrations as well as groundwater and soil water DOC. The stream DOC (bright
blue line) was from the soil water (light blue line) and groundwater
QG (dark blue line). Under low-discharge conditions (e.g.,
July–September), QG contributed a larger proportion of discharge and
stream DOC was more similar to groundwater DOC. Under wet conditions, stream
DOC resembled soil water DOC from QL. (b–g) The local soil water DOC
concentration for the six sampling locations shown in Fig. 1b, including three
planar (panels b–d) and three swale locations (panels e–g). The mean ±SD for each location was calculated based on measurements at
different depths with 10 or 20 cm intervals from the soil surface down to a
depth of hand-auger refusal.
ResultsWater dynamics
The total precipitation from 1 April 2009 to 31 March 2010 was 1130 mm. Stream discharge was highly responsive to intense
precipitation events and was high (∼10-2 m d-1) in
spring and fall compared with summer with high soil T and high ET
(∼10-5 m d-1). The model captured the temporal dynamics
of daily discharge, ET, and soil T with NSE values of 0.68, 0.72, and 0.62,
respectively (Fig. 3a, b). The model estimated that 47.5 % of annual
precipitation contributed to discharge, whereas the rest contributed to ET. The stream
discharge has three components: surface runoff QS, soil water
interflow QL (lateral flow), and groundwater flow QG from the shallow
subsurface that interacts with the stream (Fig. 2). On average, lateral
flow QL is about 90.2 % and surface runoff QS is about 2.3 %.
Following the conductivity mass-balance hydrograph separation (Lim et al., 2005), QG was estimated to be
1.3×10-4 and 4.0×10-5 m d-1 for the wet
and dry periods (August–September), which is equivalent to 6.9 % and 42.2 % of
average stream discharge at the corresponding times, respectively. Overall
QG accounted for ∼7.5 % of the annual Q, similar to
previously reported values (Li et al., 2017a; Hoagland et al., 2017). In
the dry months from August to September, the stream was almost dry with no
visible flow, and the relative contribution of groundwater to discharge was
comparable to that of QL (Fig. 3b). The unsaturated water storage
Su was typically more than 10 times larger than the saturated storage
Ss such that the ST and Su curves almost overlapped (Fig. 3c).
Ss was negligible in the dry period (close to 0 m), contributing
negligibly to the stream. Hydrological connectivity (Ics/Width) covaried with
Ss but showed significant temporal fluctuations. High summer ET drove
the catchment to drier conditions, thereby decreasing the connectivity to
the stream.
Temporal patterns of DOC concentrations
The model captured the
general trend of stream DOC (NSE of 0.55 for the monthly DOC concentration;
Fig. 4). Under dry conditions (e.g., Q<1.0×10-4 m d-1), QG contributed substantially to Q (∼32 %–71 %;
Fig. 3), and the stream DOC concentration reflected the mixing of groundwater
and soil water (Fig. 4a), with a contribution from groundwater DOC of
7 %–17 %. Under wet conditions, the stream DOC concentration overlapped with the
soil water DOC concentration (light blue line in Fig. 4). Only
∼1 %–8 % of stream DOC was sourced from groundwater at these
times.
Spatial profiles in May (wet), August (dry), and October (wet
after dry) of 2009: (a) soil T, (b) soil moisture, (c) hydrologically
connected zones, (d) local DOC production rates rp, and (e) soil water
DOC concentration. The soil DOC and rp were high in swales and the valley that had a relatively high soil water and SOC content (Fig. 1c).
Although water content in August was relatively low compared with May and
October, high soil T led to high rp, with most DOC production and
accumulation in zones that were disconnected from the stream.
The temporal dynamics of soil water data showed relatively small temporal
variation compared with stream DOC (Fig. 4b, c, d, e, f, g), and local soil pools were
not always hydrologically connected to the stream. The simulated soil water
DOC captured this small-variation trend with acceptable overall model
performance (i.e., NSE >0.5), although the goodness of fit was
lower in some locations, e.g., a NSE value of 0.36 (SPRT), 0.42 (SPMS), 0.60
(SPVF), 0.46 (SSRT), 0.40 (SSMS), and 0.51 (SSVF). The variation in model
performance at different locations may arise from the lack of detailed
information on local soil properties and organic carbon content. Although
the model explicitly considered spatial heterogeneities such as topography
and soil properties, averaged values represented grid sizes from 10 to 100 m, and this local scale was large compared with the field sampling size (e.g.,
lysimeters with a diameter of 5 cm). Geochemical processes are sensitive to
local properties, including SOC %, SOC surface area, and sorption sites,
and the representation of these properties was based on a few measurements
that were only coarsely defined as ridgetop, mid-slope, and valley floor.
Temporal dynamics of DOC storage, influent rate (rainfall
Rr, groundwater Rg, production Rp), and outflow rate (effluent Re) at the catchment scale. The stored DOC mass (dark red line) was calculated as follows: (DOC influent rate - outflow rate) × time. The temporal Re dynamics mostly followed the trend of discharge (black line, top panel), whereas Rp mostly followed the trend of soil T (orange line, top
panel).
(a) Relationship of daily discharge (Q) with stream DOC
concentration: open circles are simulations and filled circles with a black
outline are data. (b) Relationship of daily discharge (Q) with soil water storage ST, connectivity
(Ics/Width), the catchment-scale DOC export rate Re, and the DOC production
rate Rp. At low Q, the stream water transitioned from organic-poor
groundwater to organic-rich water from the valley floor and swales, leading to a
flushing (positive) pattern. At higher Q, the stream water shifted from
organic-rich soil water from swales and valley areas to lower DOC water from
planar hillslopes and uplands, decreasing the stream DOC concentration and
resulting in a dilution C–Q pattern. Re increased by 2 orders of
magnitude with increasing Q, whereas Rp varied within an order of
magnitude.
The catchment-scale DOC production rate Rp and export rate
Re as a function of (a) soil T, (b) soil water storage ST, and (c) hydrological connectivity (Ics/Width). Cross symbols are daily values in the
base case. Rp increased with soil T and decreased slightly with ST
and connectivity. In contrast, Re increased with ST and connectivity
but decreased with soil T. Re tended to decrease with soil T in the hot,
dry summer due to low discharge during that period.
Spatial patterns and mass balance
Spatial patterns vary between
May (wet), August (dry), and October (wet after dry) (Fig. 5). In May, the
average soil T was around 12 ∘C with small spatial variations
(<3∘C). Most flow-convergent areas (valley areas and swales)
were well connected to the stream and had a high water content (Fig. 5b, c).
The distribution resembles that of SOC (Fig. 1c) and water content (Fig. 5b), with a high rp and soil water DOC concentration in swales and valley.
Low rp in relatively dry planar hillslopes and uplands led to a low soil
water DOC concentration. In August, the average soil T increased to around
20 ∘C. The hydrologically connected zones shrank to the immediate
vicinity of the stream, but rp increased 2-fold from May. The simulated
soil water DOC concentration increased by a factor of 2 across the whole
catchment, especially in hillslope and uplands on the north side, because
the DOC produced was trapped in low soil moisture areas that were not
hydrologically connected to the stream. This indicates that DOC samples
collected on the south side may not represent the DOC dynamics of the entire
catchment, especially in the summer and fall dry months. In October,
rp decreased as soil cooled down, but increased precipitation and
decreased ET expanded the hydrologically connected zones beyond swales and
valley areas (Fig. 5c), promoting the desorption and the flushing of stored DOC.
The soil water DOC concentration, however, remained high because of the large
store of sorbed DOC produced during the antecedent dry times.
Figure 6 shows the catchment-scale DOC production and export rates and mass
balance. Generally, the daily Rp (5.1×105 mg d-1) was
greater than the daily Rr from rainfall (1.6×105 mg d-1) or groundwater Rg (1.2×104 mg d-1). During storm
events, Rr occasionally exceeded Rp. Rp was generally high in
summer, despite low water storage. The export rate Re did not follow the
temporal patterns of the total input rate
(Rp+Rr+Rg) or Rp. Instead, it primarily followed the
discharge patterns: large rainfall events exported disproportionally high
DOC, plummeting the DOC mass within the catchment. From the wet to dry period,
as water levels dropped, DOC accumulated within the catchment (Fig. 5e,
May to August). During the dry-to-wet transition, as the catchment became
wetter, the contributing areas expanded to the uplands and the DOC was flushed
out, reducing the overall DOC soil pool to much lower values (Fig. 5e,
August–October). The DOC mass storage increased by 1.8×106 mg
over the year, which was about 1.0 % of the overall DOC production, indicating a
general mass balance at the catchment scale.
Sensitivity analysis of temporal DOC rates for (a) soil
temperature f(T) and (b) soil moisture f(Sw). A varying Q10 value in
f(T) had a larger influence on Rp than varying f(Sw). Neither f(T)
nor f(Sw) had a significant influence on Re. Instead, Re mostly
followed the temporal trend of discharge, indicating the predominant control
of hydrological conditions.
Sensitivity analysis of the sorption equilibrium constant Keq on
(a)Rp and Re and on (b) DOC sorbed on soils averaged at the catchment
scale. High Keq led to more DOC sorbed on soils and, therefore, lower
Re. However, Re showed similar temporal patterns regardless of
Keq.
C–Q patterns and rate dependence
The C–Q relationships showed a
slightly positive correlation at low Q followed by a negative correlation at
higher Q (Fig. 7a). The simulated C–Q relationship captured this trend but
overestimated the positive relationship at low Q. The simulated C–Q
relationships showed a general dilution behavior with the C–Q slope b=-0.23 and CV[DOC]CVQ=0.22, which was consistent with the
general pattern exhibited in the field data (Fig. 7a). This C–Q
pattern can be explained by the dynamics of different water sources with
different DOC contributing to the stream. At low discharges (<1.8×10-4 m d-1) with small water storage (0.25–0.28 m) and
connectivity (Ics/Width<0.1) (Fig. 7b), the stream DOC was a mix
of organic-poor groundwater and organic-rich swales and valley floor zones.
As connectivity and discharge increased and the stream expanded, the
contribution of organic-rich swales increased, elevating the DOC concentration
to its maximum. Under even wetter conditions with connectivity exceeding 0.1,
the contribution from planar hillslopes and uplands with a lower DOC
concentration increased, diluting the organic-rich DOC from swales and
valley areas. Daily Re correlated positively with ST, hydrological
connectivity, and Q, and increased by 2 orders of magnitude as Q rose by
3 order of magnitude. The variation of daily Rp with Q was small
(105–106 mg d-1) compared with that of Re (Fig. 7b). Values of
Rp depended more on soil T than on soil water storage and hydrological
connectivity (Ics/Width) (Fig. 8). In contrast, Re increased with soil
water storage ST but notably decreased with soil T (>17∘C) due to the low discharge during the hot and dry summer.
C–Q relationships under different (a)T, (b)f(Sw), and
(c) sorption equilibrium constants Keq for the two extreme cases. The
C–Q patterns were similar in all cases, although the extent of dilution
slightly changed. This indicates potential factors other than reaction
kinetics and thermodynamics that regulate C–Q patterns.
Sensitivity analysis of groundwater on rates (Rp and
Re) and C–Q relationships: (a) scenarios with a different groundwater
volume contribution (%) to stream discharge and (b) scenarios with a
different groundwater DOC concentration (DOCGW). DOCGW and GW
(QG/Q) in the base case were 1.2 mg L-1 and 7.5 %, respectively. “2.5 GW”
in panel (a) represents the case with 2.5 times QG compared with
the base case. Increases in the relative groundwater contribution lowered
Re and shifted the C–Q pattern from an overall dilution pattern to an
overall flushing pattern; changing DOCGW had negligible influence on the
DOC rates and C–Q patterns.
Sensitivity analysisControl of temperature, soil moisture, and sorption
Higher Q10 values in f(T) led to more pronounced
seasonality in Rp (Fig. 9a). The Rp for Q10=4.0 was more
than 4 times higher than that of Q10=1.0 in summer, and much lower in
winter with low soil T (<10∘C). In contrast, the
temporal pattern of Re almost overlapped at different Q10 values,
and it mostly followed the discharge dynamics (black line in Fig. 9). Daily
Rp varied only slightly (within 15 %) with different f(Sw) (Fig. S4b), while Re showed very little change (Fig. 9b). Although we varied
Q10 from 1.0 to 4.0 in f(T), it is worth noting that varying the kinetic rate
constant, SOC surface area, volume fraction, and biomass amount could have
similar effects (not shown here) because they are all multiplied in Eq. (2).
Simulations showed that strong DOC sorption (Keq=101.0) did not
change Rp but lowered the stream DOC concentration and resulted in smaller
Re (Fig. 10a). DOC sorption had little impact on Rp, but strong
sorption decreased the magnitude of Re by 10 %–69 %. The sorbed DOC
concentration differed by more than a factor of 3, with more sorbed DOC with
larger Keq values (Fig. 10b). Large amounts of sorbed DOC persisted
until early fall, when large rainfall events flushed out sorbed DOC and
reduced DOC storage (Fig. 6). This means that catchments can store large
quantities of DOC, although the specific amount of DOC stored depends on sorption
capacity.
Varying DOC production kinetics did not change the overall C–Q patterns,
although the magnitude of overall dilution changed slightly in cases with
different f(T) and Keq (Fig. 11). High Q10 values in f(T) led to less
dilution, due to more accumulated soil DOC in the dry period (low discharge)
and, thus, more DOC flushing as discharge increased in the dry-to-wet period.
High Keq resulted in less dilution as the higher sorption capacity acts as a
stronger buffer to compensate for the concentration variations.
Groundwater control on DOC export
As shown in Fig. 12, changing the
groundwater volume contribution to stream (GW) had more significant impacts
than changing the groundwater DOC concentration (DOCGW), especially at low
discharges (Q<1.8×10-4 m d-1). Increasing the GW
contribution from no GW to 2.5 GW (i.e., 18.8 %) lowered stream DOC at low
discharges, shifting the C–Q pattern from overall dilution (or a chevron
pattern) to overall flushing (or flushing until stable). More specifically,
the threshold that separated distinct phases of these segmented C–Q
responses (Fig. 12a2) shifted from Q=1.8×10-4 to
about 1.0×10-3 m d-1. This reflects the relative groundwater
contribution to streamflow for each case. In contrast, varying the groundwater
DOC concentration (DOCGW) by 2 orders of magnitude while keeping the
same groundwater contribution (GW) did not change C–Q pattern.
Total annual Rp (red) and Re (blue) under two
groundwater volume contribution conditions (QG/Q=7.5 % and
18.8 %) for three different variables: (a) soil T, (b) soil moisture, and
(c) sorption equilibrium Keq. Rp was not influenced by a deeper
groundwater contribution, so there is only one curve for each variable.
Rp has higher dependence on temperature than on soil moisture function
form and sorption capacity.
Figure 13 summarizes the annual total Rp and Re in all sensitivity
test scenarios. Annual Rp was more sensitive to T than to Sw or
sorption thermodynamics. Annual Re was less sensitive to T variation,
although it also increased with Q10 because a higher production led to
higher DOC export. Annual Rp also depended on f(Sw), with the
threshold function f4(Sw) (Sect. 2.6) having the highest production
rates. However, Re did not follow the trend of Rp (Fig. 13b).
Generally, under the same hydrological conditions, a doubling of Rp only
led to about a 50 % increase in Re. Higher sorption affinity (higher
Keq) did not change production rates but could reduce DOC export by
about 30 % due to high DOC storage in soils. High relative
groundwater inputs (18.8 % versus 7.5 %) lowered Re in all scenarios
because more water came from deeper organic-poor groundwater.
Discussion
This study revealed that DOC production was primarily regulated by temperature, but the
lateral export of DOC was controlled by hydrological conditions.
This work contributes to the growing body of research concluding that lateral carbon
flux is determined by water routing and hydrological connectivity
and only secondarily by biological activity (Zarnetske et al., 2018). Although soil respiration and
vertical CO2 fluxes are closely related, this work focuses on
the net production and export of DOC because it has been studied and
understood to a much lesser extent than soil respiration (Tank
et al., 2018). To better appreciate the relative importance of
land–water–atmosphere carbon fluxes, future research needs to fully
integrate lateral DOC fluxes in concert with vertical fluxes of CO2
across terrestrial and freshwater ecosystems.
DOC production
The DOC production rate Rp depends more on T than on
water storage or soil moisture. This finding is expected, as DOC production
is biologically mediated and, thus, influenced by temperature and the
metabolic dependence on temperature (Gillooly et al.,
2001). Although the local-scale soil moisture varies from 0.40 at the ridgetop to 0.70 in swales and riparian zones (Fig. 5b), the averaged
catchment-scale soil moisture has relatively small variations across
different seasons in this temperate humid catchment (0.46 to 0.56 on average
over the whole catchment), especially compared with places where water is
limited and soil moisture can drop below 0.15 (Korres
et al., 2015). This small variation is due to the capability of the
shale-derived, clay-rich soils at Shale Hills to hold water (Xiao et
al., 2019). The influence of soil moisture on DOC production is likely higher in catchments
with more pronounced seasonal changes and more fluctuations in soil
moisture.
This work also suggests that catchment-scale (Rp) and local-scale
(rp) production rates have different controls. The rate law used at the
local scale is measured at relatively small scales, i.e., 0.1–2.0 m in soil
pedons (Bauer et al., 2008; Yan et al., 2016). Our results
show that even when the rate law with an optimum soil moisture was used at
the local scale (f4(Sw) in Fig. S4b), the catchment-scale rates do
not exhibit maximum rates at an “optimal” soil moisture (Fig. 8),
indicating different controls at the local scale versus the catchment scale. In
addition, due to the spatial heterogeneities of T, soil moisture, and SOC
content, the temporal variations of Rp and rp may be not consistent.
The daily Rp spanned less than an order of magnitude with its maximum in
the dry, hot summer and its minimum in the wet, cold winter and spring (Fig. 6). Local-scale rp exhibited similar temporal dynamics but varied by
more than 2 orders of magnitude, with rapid production mostly in “hot
spots” (swales and riparian zones) with persistently high water and SOC content (Fig. 5). Note that the local-scale rate laws are
often used directly at the catchment scale or at even larger scales (Crowther et al.,
2016; Conant et al., 2011; Fissore et al., 2009; Moyano et al., 2012). This
work suggests that although local-scale rate laws have been developed
extensively, direct extrapolation of rates from local to catchment scales
can be misleading. This speaks to the importance of understanding controls
on biogeochemical transformation rates and developing reaction rate theories
at the catchment scale for regional-scale and global-scale simulations.
The simulations here suggest that DOC storage depends on the sorbing capacity of soils such that clay content and the presence of organo-mineral aggregates
might play a role in mediating DOC dynamics (Lehmann et al., 2007;
Cincotta et al., 2019).
DOC production depends on catchment size, hydrogeologic structure,
vegetation, and climatic setting. Geomorphological and ecological processes
have been shown to co-generate systematic differences in the vertical and
lateral distribution of SOC and plant biomass, with a greater concentration of
organic carbon in the valley floor than in hillslopes in some catchments
(Piney et al., 2018; Temnerud et al., 2016; Campeau et al., 2019; Thomas
et al., 2016) and enriched SOC in the uplands in other catchments
(Herndon et al., 2015). These differences may explain the large variation of stream DOC in catchments
under diverse climate conditions (Moatar et al., 2017).
The median stream DOC at Shale Hills is relatively high (10.0 mg L-1),
compared with 3.0 mg L-1 in temperate humid catchments in Germany (Musolff et al., 2018), 4.1 mg L-1 in some UK catchments
with oceanic climate (Monteith et al., 2015), and 4.5 mg L-1 in
France (Moatar et al., 2017). It is also close to 9.5 mg L-1
in boreal catchments in Sweden (Winterdahl et al.,
2014), and 8.1 mg L-1 measured in boreal wet and 10.5 mg L-1 in boreal dry catchments in
Norway and Finland (de Wit et al., 2016). These differences
suggest that climate, vegetation, and landscape heterogeneity may together
shape how much DOC can be produced as well as when, where, and to what degree a hill
is connected to a stream and the export of DOC at different times.
Temporal asynchrony of DOC production and export
The contrasting
temporal patterns of simulated DOC production and export reflect the
asynchronous nature of the two processes at the catchment scale. Local DOC
production is influenced by the seasonal pattern of soil T, whereas the export
is predominantly controlled by precipitation events and antecedent
conditions that modulate the degree to which DOC production zones are
hydrologically connected to the stream. This differs from studies showing
the synchronization of DOC production and export in temperate
climatic regions at soil pedons (Michalzik et al., 2001). This may be due to the relatively short water residence time and high connectivity at
the pedon scale. The temporal asynchrony between DOC production and
export is therefore strongly influenced by the seasonality of temperature
and precipitation, which is shaped by local climate. At Shale Hills the wet winter and spring
happens to be the cold season, whereas the dry summer is hot. In the summer,
the catchment essentially produces and stores DOC in soil water and soil
surfaces and waits for the arrival of the next storm to export. In other words, low hydrological connectivity in the summer imposes a
lag period with respect to DOC export such that the DOC we see today is
often the DOC produced a while ago. As such, the catchment acts as a
DOC producer in the summer and a DOC exporter in spring and winter
when the soil is wetter.
These findings indicate a strong climatic control over DOC production and
export. In places where climate seasonality is not as strong and catchments
are hydrologically connected to streams throughout the year, we can expect to see
DOC export all year long and, therefore, much less asynchrony. In
places with strong seasonality, a few high-flow events can dominate the DOC
export of the whole year. Under the Mediterranean climate with strong
seasonality, for example, antecedent moisture conditions have been observed
to be essential for understanding the temporal pattern of DOC and nutrient
(N) export (Bernal et al., 2005, 2002). Hydrological connectivity and
water flow paths become dominant as subsurface saturation expands across
valley floors and into hillslopes (Covino, 2017; Abbott et al., 2016).
Implications for vertical and lateral carbon
fluxes
This work focuses on DOC lateral fluxes and does not simulate the
carbon loss through soil respiration and associated vertical CO2
fluxes, which has been the focus of some previous work (Brantley et al., 2018; Hasenmueller et al., 2015). Soil respiration is an
important pathway of carbon flux that, similar to DOC production, can be
shaped by soil temperature and moisture. Generally, warm temperature and
medium soil moisture provide optimal conditions for microbial respiration,
leading to significant vertical losses of carbon during summer months
(Perdrial et al., 2018; Stielstra et al., 2015). In contrast, low
temperature and high soil moisture can hinder aerobic respiration and
associated carbon losses as CO2 (Smith et al.,
2018), effectively accumulating DOC until the arrival of large storms. This pattern is consistent with
observations that the total CO2 release and DOC production are positively
correlated (Neff and Hooper, 2002). The dependence of DOC
production and export on soil T and soil moisture might also hold true for
soil respiration. Conversely, as part of the sorbed DOC may be respired
by microbes into CO2, our model might overestimate the DOC accumulation
in the catchment, especially in summer.
This work does not consider the transport of particulate organic carbon (POC)
in soil water and stream water, although POC can play an important role in the
carbon budget and biogeochemical cycles in some cases (Ludwig et al.,
1996; Diem et al., 2013). In Shale Hills, DOC
comprises a major fraction (between 70 % and 80 %) of the total organic
carbon export (Jordan et al., 1997). Similar patterns have been
reported for organic carbon export at the global scale (Alvarez-Cobelas et al., 2012). However, POC export can be
significant in anthropogenically impacted areas (Correll et al., 2001; Mattsson et
al., 2005) with significant disturbance (Abbott et al., 2016). They follow a different
temporal pattern from DOC, due to different sources, transport
mechanisms, and sensitivity to hydrologic variations (Dhillon and
Inamdar, 2014; Alvarez-Cobelas et al., 2012).
Regulation of C–Q patterns
During dry periods, stream water mostly
reflects the carbon-poor groundwater. As precipitation wets the catchment,
the valley floor area that is characterized by a high SOC and DOC concentration is
connected to the stream (Figs. 5, 7), elevating the stream DOC.
This increase in DOC concentration continues until the catchment becomes
wetter and expands the connected zones to the whole valley and swales. Under
these conditions, the influence of high DOC in the vicinity of the stream fades
and the DOC concentration decreases. This occurs at a threshold connectivity
of about ∼0.1 (≈ the valley width divided by the catchment
width). In other words, during wet periods when the whole catchment is
hydrologically connected to the stream, the stream DOC reflects the “average”
concentration across the catchment (∼2.5 mg L-1). The increase
and subsequent decrease pattern (or chevron pattern) therefore indicates the
presence of three end-members from different sources: the groundwater with a
very low DOC concentration, the soil water at stream beds and in
organic-rich swales with the highest DOC content, and the uphill soil water with a
DOC level in between these two.
The overall dilution (or chevron) C–Q pattern observed here with a maximum
at a mid-range discharge contrasts the commonly observed flushing
pattern for DOC (Moatar et al., 2017). In fact, it more closely
resembles the hysteresis behavior often observed in storm and
snowmelt events for metals and nutrients (Zhi et al., 2019; Duncan et
al., 2017). Previous field studies have illustrated that the hydrological
connectivity to the stream versus the distribution of SOC ultimately
dictates the spatial and temporal dynamics of the DOC concentration in soil and
stream water, leading to different C–Q relationships (dilution versus
flushing) (Bernhardt et al., 2017; Bernal and Sabater, 2012; Covino,
2017). This has been illustrated by different C–Q relationships at Shale Hills (USA) and Plynlimon (UK) (Herndon et al., 2015). Stream water at Shale Hills is
derived from SOC-rich swales with a high DOC concentration at low flow and
from both swales and hillslopes with a low DOC concentration when discharge
increases. Conversely, at Plynlimon, SOC is enriched in uplands; therefore,
concentrations are high at high flow when water flows connect
SOC-rich uplands. Our reactive transport modeling provides a quantitative
and mechanistic approach to explain the overall C–Q
patterns, which have generally been interpreted as a production/source
limitation (Covino, 2017; Zarnetske et al., 2018). Our results suggest
that the spatial distribution of source zones and the degree of their
connectivity to the stream determine when they are flushed out. Modeling
approaches such as the one presented here can help us understand the
mechanisms underlying C–Q patterns, and, thus, improve our ability to predict
the evolution of C–Q trajectories under changing conditions.
C–Q patterns also relate to the mixture of different sources of water in the
stream, which is composed of the time-varying relative contribution from the shallow soil
water and relatively deep groundwater. Their DOC contribution can
be affected by the vertical distribution of reacting materials (Musolff
et al., 2017; Bishop et al., 2004; Seibert et al., 2009; Winterdahl et al.,
2016) and the relative volume contribution of source water (soil water versus
groundwater below the soil-weathered rock interface) to the stream (Zhi
et al., 2019; Radke et al., 2019; Weigand et al., 2017). Within the shale
bedrock, the groundwater contribution to the stream is relatively small
(∼7.5 %) at Shale Hills. Soil water (although from a very
limited swale area) dominates inflow to the stream even during the summer
dry period. When the groundwater volume input increases to about 18.8 % of
the streamflow by volume (2.5 times the actual case; Fig. 12), the C–Q relationships shift to an overall
flushing pattern. This may provide a potential explanation for the DOC C–Q
flushing pattern at sandstone-dominant Garner Run (a neighboring catchment of
Shale Hills), where the groundwater contributions to the stream are
typically higher (Hoagland et al., 2017; Li et al., 2018). More
interestingly, when the groundwater contribution is “sufficiently” high,
it can mask the signature of the swale-derived soil water such that the
three-end-member chevron C–Q pattern become a two-end-member pattern with
monotonic flushing pattern that is similar to the observation in Coal Creek where groundwater
contributes about 20 % annually (Zhi et al., 2019). C–Q relationships
have been categorized into nine patterns, including three monotonic and six segmented
types (Moatar et al., 2017; Underwood et al., 2017). The shifting
threshold that separates segments of C–Q responses by the relative
groundwater contribution in this work (Fig. 12) suggests that the relative
contribution of groundwater to streamflow may play a pivotal role in shaping
the C–Q patterns. This threshold value can potentially provide a rough
estimation for the relative contribution of different end-members to the
stream.
The mechanisms that regulate DOC C–Q patterns – seasonally variable
hydrological connectivity and groundwater contribution – are consistent with
previous literature on geogenic species (Mn, Fe), isotopes, and particle
fluxes at Shale Hills (Herndon et al., 2018; Kim et al., 2018; Sullivan
et al., 2016; Thomas et al., 2013). For example, Mn is associated with DOC
via biotic cycling and storage in plant species, and Fe is associated with
DOC via aqueous complexation. Therefore, both solutes are more abundant in
shallow soils. The C–Q pattern of Fe and Mn shows a dilution pattern with
concentrations decreasing as discharge increases (Herndon et al., 2015;
2018). In the dry summer, stream water derives from rich-organic swales and
riparian zones with high concentrations of soluble Fe and Mn (Herndon et al., 2018), leading to corresponding high
stream concentrations. At high flows, these solutes are diluted by the
influx of uphill soil water without as much DOC. This again emphasizes the
key role of solute sources and hydrological dynamics in controlling stream
chemistry.
Conclusions
The production and export of DOC remain central uncertainties in determining
the ecosystem-level carbon balance. These uncertainties persist because of complex interacting controls on DOC production and
export. Indeed, few studies have quantitatively addressed the linkages
between SOC processing, hydrological conditions, and corresponding DOC
processing and export at the catchment scale. We found that DOC
production was the major DOC source at Shale Hills (75 % compared with
23 % from precipitation and 2 % from groundwater). The simulations
showed that the temporal dynamics of DOC export rates (Re) were more
linked to hydrological flow paths and precipitation events. A sensitivity analysis further confirmed that the DOC production rates Rp were
primarily controlled by temperature: changing the temperature dependence altered DOC concentrations significantly, whereas the effects of
changing soil moisture dependence were relatively small. Conversely,
DOC export was most sensitive to changes in hydrology, rendering more than
2 orders of magnitude differences in dry and wet periods. This difference
in environmental drivers led to an asynchrony between DOC production and export. During the wet
period (spring and winter), the catchment was well connected and DOC
production and export occurred simultaneously. During summer, DOC
accumulated (often in sorbed form) in soils disconnected from the stream, and
DOC export was limited and constrained to the near-stream areas. In other
words, the catchment serves as a DOC producer in the dry and hot summer but as an exporter in the wet and cold winter.
This work quantitatively demonstrates the key role of hydrological flow
paths and the degree of connectivity in determining the C–Q patterns
exhibited at the catchment outlet. At low discharges where connectivity is
limited (Ics/Width< 0.1), stream DOC was mainly sourced from groundwater or from the
valley floor and swales with enriched SOC. At higher discharges, an increasing contribution of soil lateral flow from planar hillslopes and uplands with low soil water DOC decreased the stream DOC concentration, ultimately rendering a dilution C–Q pattern. Although changing DOC reaction characteristics
alters the soil water DOC concentration, there is little change in the overall
C–Q patterns. However, when groundwater contributes 18.8 % of total annual
discharge, the stream DOC concentration increases with discharge and flushing
patterns emerge. This underscores the significance of the relative contribution and chemical signature of
different water sources in shaping DOC export patterns. This study provides
new insights into how DOC production and export interact at multiple scales
and emphasizes the importance of considering different constraints when
projecting the response of lateral and vertical carbon fluxes to climate
changes.
Data availability
The field data have been digitized and are
accessible from the national CZO data portal
(http://criticalzone.org/shale-hills/data/datasets/; Brantley and Duffy, 2012). The source code of BFP
(BioRT-Flux-PIHM) and the input files necessary to reproduce the results are
available from the authors upon request (lili@engr.psu.edu).
The supplement related to this article is available online at: https://doi.org/10.5194/hess-24-945-2020-supplement.
Author contributions
HW, LL, and the co-authors conceived
the idea and designed the numerical experiments based on ideas generated
from a workshop and monthly discussions. HW ran the simulations, analyzed the
simulation results, and wrote the first draft of the paper. All
co-authors participated in editing the paper.
Competing interests
The authors declare that they have no conflict of interest.
Acknowledgements
We acknowledge financial support from the US
National Science Foundation Geobiology and Low-Temperature Geochemistry
program (grant EAR-1724171). We appreciate data from the Susquehanna
Shale Hills Critical Zone Observatory (SSHCZO), which is supported by the National Science
Foundation (grant no. EAR-0725019: principal investigator Christopher J. Duffy; grant no. EAR-1239285: principal investigator Susan L. Brantley;
and grant no. EAR-1331726: principal investigator Susan L. Brantley). Data were collected in Penn State's Stone
Valley Forest, which is funded by the Penn State College of Agricultural
Sciences, Department of Ecosystem Science and Management and managed by the
staff of the Forestlands Management Office. We thank the ISU Center for
Ecological Research and Education and EPSCoR (grant no. IIA 1301792) that
stimulated ideas in this paper. Susana Bernal's work was funded by CANTERA
(grant no. RTI-2018-094521-B-101) and a Ramón y Cajal fellow (grant no. RYC-2017-22643) from
the Spanish Ministry of Science, Innovation, and Universities.
Financial support
This research has been supported by the US National Science Foundation Geobiology and Low-Temperature Geochemistry program (grant no. EAR-1724171).
Review statement
This paper was edited by Christian Stamm and reviewed by two anonymous referees.
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