ESurfEarth Surface DynamicsESurfEarth Surf. Dynam.2196-632XCopernicus PublicationsGöttingen, Germany10.5194/esurf-6-723-2018Late Holocene channel pattern change from laterally stable to meandering –
a palaeohydrological reconstructionLate Holocene channel pattern change from laterally stable to
meanderingCandelJasper H. J.jasper.candel@wur.nlKleinhansMaarten G.https://orcid.org/0000-0002-9484-1673MakaskeBartHoekWim Z.QuikCindyWallingaJakobSoil Geography and Landscape Group, Wageningen University &
Research, Wageningen, P.O. Box 47, 6700AA, the NetherlandsDepartment of Physical Geography, Utrecht University, Utrecht, P.O. Box 80125, 3508TC, the NetherlandsJasper H. J. Candel (jasper.candel@wur.nl)31August2018637237413April201815May201819July201817August2018This work is licensed under the Creative Commons Attribution 4.0 International License. To view a copy of this licence, visit https://creativecommons.org/licenses/by/4.0/This article is available from https://esurf.copernicus.org/articles/6/723/2018/esurf-6-723-2018.htmlThe full text article is available as a PDF file from https://esurf.copernicus.org/articles/6/723/2018/esurf-6-723-2018.pdf
River channel patterns may alter due to changes in hydrological regime
related to changes in climate and/or land cover. Such changes are well
documented for transitions between meandering and braiding rivers, whereas
channel pattern changes between laterally stable and meandering rivers are
poorly documented and understood. We hypothesize that many low-energy
meandering rivers had relatively low peak discharges and were laterally
stable during most of the Holocene, when climate was relatively stable and
human impact was limited. Our objectives in this work are to identify a Late
Holocene channel pattern change for the low-energy Overijsselse Vecht river,
to develop and apply a novel methodology to reconstruct discharge as a
function of time following a stochastic approach, and to relate this channel
pattern change to reconstructed hydrological changes. We established that the
Overijsselse Vecht was laterally virtually stable throughout the Holocene
until the Late Middle Ages, after which large meanders formed at lateral
migration rates of about 2 m yr-1. The lateral stability before the
Late Middle Ages was constrained using a combination of coring information,
ground-penetrating radar (GPR), radiocarbon (14C) dating, and
optically stimulated luminescence (OSL) dating. We quantified bankfull
palaeodischarge as a function of time based on channel dimensions that were
reconstructed from the scroll bar sequence and channel cut-offs using coring
information and GPR data, combined with chronological constraints from
historical maps and OSL dating. We found that the bankfull discharge was
significantly greater during the meandering phase compared to the laterally
stable phase. Empirical channel and bar pattern models showed that this
increase can explain the channel pattern change. The bankfull discharge
increase likely reflects climate changes related to the Little Ice Age and/or
land use changes in the catchment, in particular as a result of peat
reclamation and exploitation.
Introduction
Channel patterns describe the planform of a river, which reflects the
interaction of the river channel with its floodplain. Channel patterns are
classically distinguished: laterally inactive channels consist of straight
and sinuous stable planforms, whereas laterally active channels consist of
meandering and braiding planforms (Leopold and Wolman, 1957; Nanson and
Knighton, 1996). Flume experiments and field data have shown that the channel
pattern depends on several variables (Kleinhans, 2010). The first is the
potential specific stream power, which is the product of the channel-forming
discharge and valley slope (Kleinhans and Van den Berg, 2011; Nanson and
Croke, 1992). The second is the bank erodibility (Ferguson, 1987; Friedkin,
1945), which is determined by the presence of bedrock in the valley side
(Turowski et al., 2008), the bank cohesiveness (Peakall et al., 2007), and
vegetation (Gurnell, 2014; Millar, 2000). The third is the type and amount of
sediment supply (Gibling and Davies, 2012; Nanson and Croke, 1992).
Channel patterns can change in response to environmental variations
(Ferguson, 1987). Many examples of channel pattern changes from braiding
to meandering and vice versa are known to be associated with
glacial–interglacial oscillations (Vandenberghe, 1995, 2002). Studies on the last glacial–interglacial
transition have especially shown the simultaneous occurrence of channel pattern changes
with a changing climate (Kasse et al., 2016; Vandenberghe et al.,
1994). Climate change affects the vegetation, sediment availability, and
discharge regime and consequently the bank stability, sediment transport,
and potential specific stream power, resulting in different channel patterns.
Within the Holocene, several examples of channel pattern changes are
documented from braiding to meandering rivers and vice versa (Brewer and
Lewin, 1998; Lewin et al., 1977; Passmore et al., 1993; Słowik, 2015).
However, channel pattern changes between laterally stable and meandering
rivers have rarely been reported (Lewin and Macklin, 2010),
except where human intervention transforms meandering rivers into heavily
regulated and laterally stable rivers by introducing weirs, dams, groynes,
and bank protection measures (Hesselink et al., 2003; Hobo et al., 2014;
Słowik, 2013; Surian and Rinaldi, 2003). The partial abandonment of
former meandering valleys may also result in underfit, laterally stable rivers
like the former Rhine branches in the Niers and Oude IJssel valley
(Janssens et al., 2012; Kasse et al., 2005).
Many studies have reported increased fluvial activity (e.g. increased
discharge, sediment transport and deposition, and bank erosion rates) in
relation to human, environmental, and climatic pressures during the Holocene
(e.g. Hoffmann et al., 2008; Lespez et al., 2015; Macklin et al., 2010;
Notebaert et al., 2018; Notebaert and Verstraeten, 2010). An example of
increased fluvial activity is known from the Pine Creek (Idaho, USA), where
mining and deforestation combined with intensive grazing resulted in an
increase in discharge and sediment input, followed by river widening and an
increase in bank erosion (Kondolf et al., 2002). The reverse
change has been observed in settings as a result of afforestation
(Kondolf et al., 2002; Liébault and Piégay, 2001) or increases in
riparian vegetation fixing the channel banks (Eekhout et al., 2014;
Vargas-Luna et al., 2016). The increase in fluvial activity during the
Holocene was corroborated by an extensive review of existing studies
concerning sediment accumulation in west and central European river
floodplains by Notebaert and Verstraeten (2010). They concluded that
sedimentation rates increased during the Middle and Late Holocene due to
environmental changes. However, it is unknown whether the channel pattern
changed simultaneously with the floodplain because no Early Holocene
channel deposits representing a stable phase were identified. De Moor et
al. (2008) hypothesized that the Geul River in southern Netherlands may have
been relatively laterally stable during the Early and Middle Holocene, while
it was actively meandering during the past 2000 years. Most of the
floodplain deposits from the laterally stable phase have not been preserved,
but De Moor et al. (2008) were able to reconstruct the bankfull depth
for both periods. They estimated the bankfull depth to be a factor of 2 to
3 higher during the Late Middle Ages compared to the Early and Middle
Holocene, and related this change to human and climate impact.
We conjecture that the change from laterally stable to meandering has
occurred in some rivers for which increased Holocene fluvial activity was
reported. The fact that such changes were not reported in the literature
may either mean that critical conditions for channel pattern change were not
reached or that evidence of such transitions is poorly preserved or left
unnoticed. Both laterally stable and meandering rivers may display sinuous
planforms, but the geomorphic processes in both rivers are different.
Laterally stable channels are rivers without meandering processes, i.e.
helicoidal flows causing bar formation and bank erosion at a significant
rate (Kleinhans and Van den Berg, 2011; Nanson and Knighton, 1996;
Seminara, 2006). In fact, the bends and channel cut-offs in laterally stable
rivers may be the result of random and local perturbations (e.g. falling
trees, beavers, bank collapse after heavy rainfall, etc.) leading to very
limited and local displacement of the channel. Meandering and laterally
stable rivers should therefore be distinguished by their different patterns
of bar and floodplain formation, rather than merely by planform
(Candel et al., 2017; Kleinhans and Van den Berg, 2011). We
suggest that identifying channel pattern changes requires more detailed
historic accounts or a much higher resolution of subsurface data than
usually gathered because palaeochannels of laterally stable channels poorly
preserve in the fluvial archive of meandering channel belts
(Van de Lageweg et al., 2016), except when they have been cut
off by random and local perturbations prior to the meandering phase. Using
numeric (e.g. Oorschot et al., 2016) or scaled (e.g. Van
Dijk et al., 2012) river simulation models is problematic for testing these
ideas because these have not yet been capable of reproducing channel
pattern changes. This reflects the lack of understanding of river processes
and patterns (Kleinhans, 2010) and the need to gather such information
from field studies.
This research entails a case study focussing on a river in which lateral
activity during the past 500 to 600 years caused spectacular meandering: the
Overijsselse Vecht in the Netherlands (Fig. 1). Previous work on this system
has identified a transition from braiding to meandering during the
Late Glacial (Huisink, 2000), while the meandering pattern remained throughout
the Holocene until the river was channelized between 1896 and 1914 CE
(Huisink, 2000; Neefjes et al., 2011). However, Quik and Wallinga (2018)
found that the meanders were relatively young, with the oldest scroll bars
dating from ca. 1400 to 1500 CE, by reconstructing meander formation using a
combination of optically stimulated luminescence (OSL) dating of scroll bars
and planform reconstruction based on historical maps. No fluvial deposits
were found dating from before this period, except from a Holocene
palaeochannel (here referred to as “Palaeochannel Q”) in a
ground-penetrating radar (GPR) profile recorded by Huisink (2000, p. 123)
13 km upstream near Hardenberg (Figs. 1b and 2). Palaeochannel Q is
relatively small compared to the meandering channel, seems to lack scroll
bars, and was already cut off on the historical map of 1720 CE. Therefore, it
is questionable whether the Overijsselse Vecht meandered prior to ca.
1400 CE. Alternatively, the river changed from a laterally stable into a
meandering river during the Late Middle Ages. Our aims are (1) to identify
whether a channel pattern change has occurred from laterally stable to
meandering by collecting and combining detailed subsurface and
geochronological data of the river from the pronounced meandering phase and
the preceding phase; (2) to develop a methodology to reconstruct bankfull
discharge as a function of time using the scroll bar deposits and channel
remnants as a geological archive of the former channel dimensions; (3) to
test whether palaeohydrological changes may explain the potential channel
pattern change; and (4) to elaborate on the potential causes for changes in
the discharge and channel pattern.
Maps of the Overijsselse Vecht. (a) Map showing the
location of the Overijsselse Vecht catchment and the location of the study
site. (b) Digital elevation map (DEM; Actueel Hoogtebestand
Nederland, 0.5×0.5m) (Van Heerd and Van't Zand, 1999) of the
downstream section of the Overijsselse Vecht river, indicating both study
sites: Junnerkoeland and Prathoek. DEM of the Junnerkoeland bend
(c) and Prathoek bend (d), including locations of cores,
OSL samples by Quik and Wallinga (2018), the OSL and 14C samples
from this study, the GPR transects, the grain size samples, and inflection
points. The possible historical course of Palaeochannel X according to
Maas (1995) is indicated. (e) Zoomed-in figure of Palaeochannel X.
(f) Topographical military map (TMK) dating from 1851 CE (CC-BY
Kadaster, 2018; Van der Linden, 1973), showing the Overijsselse Vecht during
its meandering phase.
Interpretation by Huisink (2000) of subsurface strata from GPR data
collected near Hardenberg 13 km upstream of Junnerkoeland (see location in
Fig. 1b). Horizontal strata of cover sand deposits (A) on top of the channel
deposits of an interpreted braiding system (B). A relatively small,
symmetrical palaeochannel is present (C) within the Late Glacial deposits,
hereafter referred to as “Palaeochannel Q”. Figure adapted after
Huisink (2000).
Study area
The Overijsselse Vecht (Fig. 1) is a low-energy, sand-bed river flowing from
Germany into the Netherlands, with an average annual discharge (Qm) of
22.8 m3 s-1 and a mean annual flood discharge (Qmaf)
of 160 m3 s-1 derived from the gauging station in Mariënberg
for the period 1995 to 2015 (see location in Fig. 1b). The river has a length
of 167 km, its catchment covers 3785 km2 with the highest point
+110m above sea level (m a.s.l.), and a relatively uniform valley
slope of 1.42×10-4 to 1.7×10-4 in the Dutch part of its
trajectory (TAUW, 1992; Wolfert and Maas, 2007). The Overijsselse Vecht
incised its current valley during the Late Glacial within fluvioperiglacial
sands, locally covered by aeolian cover sands (Huisink, 2000; Ter Wee, 1966;
Wolfert and Maas, 2007). During the Late Holocene, aeolian drift sands formed
along the Overijsselse Vecht as a result of agricultural overexploitation
(Van Beek and Groenewoudt, 2011). The Overijsselse Vecht was an actively
meandering river until 1896, when weirs were constructed and parts of the
river were channelized. The river was completely channelized after 1914 CE,
with five weirs controlling the water levels. Recently, sinuous side channels
bypassing the weirs have been created as part of river restoration aiming to
restore past physical and ecological characteristics of the river.
At present the topography of the meandering phase is partly still intact in
the floodplain (Maas, 1995). Wolfert and Maas (2007) reconstructed the
pre-channelization planform from historical maps of 1720, 1850, and 1890 CE.
Large differences in meander development and lateral migration rates were
found between different river reaches. In particular in areas where
non-cohesive aeolian sands formed the channel banks, large meanders formed
and lateral migration reached rates up to 3 m yr-1. In this research
we will study two of the large meanders, named Prathoek and Junnerkoeland
(Fig. 1), for which Quik and Wallinga (2018) reconstructed the scroll bar
development using OSL dating in combination with historical maps.
Here we take advantage of the preservation of a palaeochannel (here referred
to as “Palaeochannel X”) with comparable dimensions as Palaeochannel Q
(Huisink, 2000, p. 123) preserved in the Junnerkoeland as a sharp bend
(Fig. 1c). Maas (1995) interpreted Palaeochannel X to be connected to the
oldest swale of the Junnerkoeland scroll bar deposits (Fig. 1c).
Palaeochannel X, however, was likely abandoned before the scroll bar
formation because large differences in dimensions exist between
Palaeochannel X and the meander bend, but the well-preserved nature
suggests that Palaeochannel X is relatively young. The small dimensions of
both Palaeochannels X and Q would suggest that the river had
comparatively less energy and may have been relatively laterally stable
prior to the meandering phase.
MethodsLithological description
Cores were performed in a transect perpendicular to the scroll bars of both
meander bends (Fig. 1c, d). An additional transect was cored perpendicular to
Palaeochannel X (Fig. 1e). In the case that the deposit consisted of peat we used a
gouge auger (Ø: 3 cm); in the case of unsaturated sand we used an Edelman
auger, and in the case of saturated sand we used a Van der Staay suction corer
(Van de Meene et al., 1979). In total, 68 cores were performed to a maximum
depth of 7.3 m. The surface elevation of each coring site was either
determined using a GPS combined with a DEM (Van Heerd and Van't Zand, 1999)
or with a global navigation satellite system (GNSS) device. A standard method
was used to describe the sediment cores in 10 cm thick intervals using the
Dutch texture classification scheme, which approximately matches the USDA
terminology (Berendsen and Stouthamer, 2001; De Bakker and Schelling, 1966).
The median sediment grain size (D50, m) of non-organic, sandy samples
was visually checked in the field by comparison with a sand ruler. Grain size
analysis was used to estimate a D50 for the entire scroll bar deposit
(Sect. 3.3). In addition, the plant macro-remains, any visible bedding, and
colour were described. The percentage of gravel (> 2 mm) was
estimated in the field using sieves. The lithogenesis was inferred from the
lithological properties, facies geometries, and DEM topography,
distinguishing fluvial, fluvioperiglacial, cover sand, drift sand, and
residual channel-fill deposits (Huisink, 2000; Ter Wee, 1966).
Ground-penetrating radar
Ground-penetrating radar (GPR) was used to reconstruct the channel dimensions
of the scroll bars. GPR measurements were conducted with a pulseEKKO PRO
250 Hz with a SmartTow configuration. The GPR transects were placed along the
centreline of the meander bends, perpendicular to the ridge and swale
morphology (Fig. 1c, d). The electromagnetic-wave velocity was
0.060 m ns-1, derived by using isolated reflector points (Neal, 2004;
Van Heteren et al., 1998) and by comparing depths of recognizable layers with
the coring data.
Grain size analysis
In total 33 samples for grain size analysis were taken from the scroll bar
deposits and three samples were taken from Palaeochannel X. The samples of
the scroll bar deposits were taken from each 0.5 m interval from the channel
lag up to the swale surface at three locations in Junnerkoeland and two
locations in Prathoek (Fig. 1c–e). The samples of Palaeochannel X were
taken from three locations below the residual channel-fill from the former
riverbed. Grain size samples were analysed in a laboratory with a Beckman
Coulter LS230 laser particle sizer. This instrument has a measurement range
of 0.1 to 2000 µm. Samples were sieved with a 2 mm sieve and
prepared with HCl (1 M) and H2O2 (30 %). All data
were processed using a Fraunhofer.rfd optical model because of the low
clay–silt content (Agrawal et al., 1991). Finally, the average and standard
deviation were calculated for both the scroll bar deposits and Palaeochannel
X and used in the palaeohydrological calculations.
OSL dating
We used the modelled age–distance relationships determined by Quik and
Wallinga (2018) in our calculations. Their obtained OSL ages from the scroll
bar deposits were used as priors and combined with historical map data in a
Bayesian sequence model using the OxCal software (Bronk Ramsey, 2009). For
details on the method see Quik and Wallinga (2018). In this study, we took
four additional samples for OSL dating on the inner and outer bank of
Palaeochannel X. These samples were collected in an opaque PVC tube (Ø
4.5 cm) mounted on a hand auger, allowing for sampling without light exposure.
The analysis in the laboratory followed the same procedure as in Quik and
Wallinga (2018). The OSL age was determined at the Netherlands Centre for
Luminescence dating, with equivalent doses measured on small aliquots of
quartz using the SAR protocol (Murray and Wintle, 2003) and dose rates
determined from activity concentrations measured using gamma-ray
spectrometry. A bootstrapped version of the minimum age model (Cunningham and
Wallinga, 2012) was used to derive the best estimate of the burial dose and
deposition age. Given the limited amount of samples associated with
Palaeochannel X and the absence of additional age constraints from historical
maps, no Bayesian analysis was performed for these samples.
14C dating
A sample was taken in the deepest part of Palaeochannel X, at the sand–peat
interface, using a piston corer (Ø: 6 cm). Macro-remains and leaf
fragments from terrestrial species were selected from 1 cm intervals in the
laboratory using a light microscope. Samples were stored in diluted
HCl (4 %) at 5 ∘C. The sand content was measured for each
interval to precisely determine the position of the sand–peat interface.
Material with volumetric sand percentages lower than 10 % to 20 % was
considered as peat (Bos et al., 2012). The macro-remains from the centimetre
above this interface were selected for the 14C analysis, providing a
terminus ante quem date for the abandonment of the channel. The
14C age was determined by accelerator mass spectrometry (AMS) at
the Centre for Isotope Research (Groningen University). For calibration, the
IntCal13 curve was used in the OxCal 4.2.4 software (Bronk Ramsey, 2009;
Reimer et al., 2013).
Channel dimensions
The channel dimensions of Palaeochannel X were determined from the
lithological cross section. The residual channel-fill was delineated along
the sand–peat interface. Bankfull depth (Hbf) was defined from
the bottom of the palaeochannel up to the first clear knick point on the
bank, which was mapped with a GNSS device such that the width–depth ratio
was minimal (Williams, 1986). Relative error of Hbf was assumed
to be similar to the relative error of Hbf during the meandering
phase (ca. 10 %) and used in the calculations (see details below)
because both Hbf values were determined by using coring data.
Additional dimensions were measured from the delineated channel, involving
the bankfull width (W), cross-sectional area (A), and wetted perimeter
(P). These channel dimensions were also measured for Palaeochannel Q from
the GPR profile recorded by Huisink (2000, p. 123) (Fig. 2). Here we assumed
a similar relative error of W, A, and P as was taken for
Hbf.
Sketch of the cross-sectional flow area of a meandering channel used
for the bankfull palaeodischarge calculations (Allen, 1965; Hobo, 2015;
Leeder, 1973).
The river channel was assumed to have the channel dimensions as shown in
Fig. 3 during the meandering phase. This sketch is based on Allen (1965),
Leeder (1973), and Hobo (2015). The bankfull depth (Hbf) was
estimated from the coring data taken from the bottom of the channel lag up
to the surface elevation in the swales (Fig. 4). Small elevation differences
were expected to result from local variation rather than real changes in
Hbf, and therefore the average Hbf was calculated from
the smoothed bottom and surface elevation. The standard deviation of
Hbf was calculated from the actual variable bottom elevation over
the length of the scroll bar. The transverse bed slope (α) of the
inner bend was determined based on the GPR transects (Fig. 5), in which
lateral accretion surfaces could be distinguished. The angle was measured on
the steepest (middle) parts of the identified lateral accretion surfaces. The
average and standard deviation of α were calculated and used in the
calculations. The calculations of the channel dimensions follow from Fig. 3.
The bankfull width (W, m) and cross-sectional area (A, m2) were
determined by Eqs. (1) and (2).
W=1.5Hbftan(α)A=WHavgHbf is the bankfull depth (m), and Havg=7Hbf12 approximates the average water depth (m). The
wetted perimeter (P, m) was calculated from the assumed channel geometry
(Fig. 3) following Eq. (3).
P=Hbfsin(α)+W6+Hbf2+W62
The hydraulic radius (R, m) was calculated by Eq. (4).
R=AP
For each swale visible on the DEM the sinuosity (s, -), radius of curvature
(Rcurv, m) and scroll bar surface area (SBsurf,
m2) were measured. The former channel sinuosity was estimated by the use
of the DEM, measuring the distance along the swales relative to the distance
along the valley between the inflection points (Fig. 1c, d). The sinuosity of
Palaeochannel X was measured using the same approach (Fig. 1c). The channel
slope (Sc, –) was calculated from the sinuosity and valley slope
(Sv, –) determined by TAUW (1992) and Wolfert and Maas (2007)
following Eq. (5).
Sc=Sv/s
The volumetric rate of scroll bar growth (SBvol,
m3 yr-1) was determined from scroll bar surface area
(SBsurf, m2 yr-1) and the thickness between each swale
and interpolated time interval following Eq. (6):
SBvol=SBsurfHbf(1-φ)Δage
where φ is the porosity (here 0.3 to 0.35 volume fraction) (Nimmo,
2004), which was included to compare the SBvol with the sediment
transport, and Δage is the age difference between the scroll
bars (yr) based on the datings by Quik and Wallinga (2018). Although scroll
bar deposits were absent, following Eq. (6) we also calculated the volumetric
sediment transport for the fluvial deposits on the inside of Palaeochannel
X.
Stratigraphic cross sections of the study sites (for location see
Fig. 1). Lithological cross sections of Junnerkoeland (a) and
Prathoek (b). Lithogenetic cross sections of Junnerkoeland
(c) and Prathoek (d) including the OSL samples by Quik and
Wallinga (2018) and OSL and 14C dating results from this study. The
surface and erosive base elevation are indicated with dashed lines, resulting
in the inferred bankfull channel depth (Hbf).
(e) Zoomed-in lithogenic cross section of Palaeochannel X. The
thick dashed line indicates the bankfull level of the palaeochannel.
Palaeodischarge
The channel dimensions were used to calculate the bankfull discharge
(Qbf, m3 s-1). Bankfull discharge is
commonly considered an approximation of the channel-forming discharge with a
recurrence interval of 1 to 2 years (Dury, 1973; Wolman and Miller, 1960). We
assumed that the bankfull discharge was similar for both Junnerkoeland and
Prathoek, regarding the short distance between these river sections
(Fig. 1b). Hence the bankfull discharge was presented by combining the
bankfull discharges for both meander bends. The bankfull discharge was
estimated by applying the Chézy equation, following Eq. (7):
Qbf=CARSc,
where Qbf is the bankfull discharge (m3 s-1), and C
is the Chézy coefficient (m0.5 s-1). The Chézy
coefficient, i.e. flow resistance, was estimated by substituting Eq. (8) in
Eq. (7). Equation (8) is an empirical relation (Brownlie, 1983):
Qbf=R(0.3724Sc-0.2542σs0.105D50)1.529Wg0.5D501.5,
where σs is the sorting of the bed material grain size (–)
derived from the grain size analysis (Sect. 3.3) and approximated by
0.5(D50D16+D84D50), D16 and D84
are the 16th and 84th percentile sediment grain size (m), respectively,
and g is the gravitational acceleration (m2 s-1). As a
validation, the calculated Chézy coefficient was compared with average
Chézy coefficients of 12 comparable low-energy, sand-bed rivers with
scroll bars (Sv < 0.001,
90 < Qbf < 320 m3 s-1)
calculated from a large river dataset (Kleinhans and Van den Berg, 2011; Van
den Berg, 1995). The cross-sectional-averaged flow velocity (ubf,
m s-1) was calculated by following Eq. (9).
ubf=QbfA
Sediment transport
The sediment transport was calculated to compare with the SBvol
(Eq. 6). Sediment transport was calculated in two different ways. The first
method was the slightly modified Engelund and Hansen (1967) relation
following Eq. (10):
Qs,bf=0.05u5Wti(ρsρ-1)2g0.5D50C3(1-φ),
where Qs,bf is the yearly sediment transport derived from the
bankfull discharge (m3 yr-1), t is the number of seconds in a
year, i is the intermittency assumed to be 0.03 to 0.07 (Parker, 2008),
ρs is the sediment density (kg m-3), ρ is the
water density (kg m-3), and φ is the porosity assumed to be
0.3 to 0.35 (Nimmo, 2004). The relation of Engelund and Hansen was used
because the relation is suitable for sand-bed rivers with relatively low flow
velocities (Van den Berg and Van Gelder, 1993), and the input variables
required were available.
In the second method the sediment transport was determined for each discharge
magnitude and related frequency (Qs,freq) (Wolman and Miller,
1960) from present-day flow conditions by assuming that the current discharge
frequency distribution also applied to the meandering phase. We used the
hourly discharge data from 1995 to 2015 of the gauging station in
Mariënberg (Fig. 1b). This gauging station is close to the study location
and has the lowest amount of data gaps compared to the other stations. The
flow duration was calculated for intervals of 10 m3 s-1, and for
each discharge interval the sediment transport was calculated using Eq. (10),
excluding the intermittency factor. If the discharge was above bankfull, the flow would go across the floodplain.
The Chézy coefficient for the floodplain was assumed to be half the
Chézy coefficient in the channel because of the higher roughness of the
floodplain compared to the channel. We assumed that the floodplain width was
350 m for the start of the meandering phase, which was estimated from the
DEM (Fig. 1c), and that the width would increase proportionally with the
lateral migration rate for each time step during the meandering phase.
Potential specific stream power
The potential specific stream power was calculated to plot into a stability
diagram. Kleinhans and Van den Berg (2011) distinguished four different
stability fields, further building on Van den Berg (1995) and Makaske et
al. (2009): rivers with laterally stable channels, meandering rivers with
scroll bars, meandering rivers with scroll and chute bars as well as
moderately braided rivers, and braided rivers. In this research, only the
first two stability fields are relevant. These stability fields are separated
by a discriminator that represents the theoretical minimum energy needed for
the channel pattern to occur (Kleinhans and Van den Berg, 2011). The
potential specific stream power was calculated by applying the relationship
presented by Kleinhans and Van den Berg (2011) following Eq. (11):
ωpv=ρgQbfSvε,
where ωpv is the potential specific stream power
(W m-2) and ε=4.7sm-1 for
sand-bed rivers (Van den Berg, 1995). The discriminator line was plotted
by applying the relationships presented by Makaske et al. (2009) and Kleinhans
and Van den Berg (2011) following Eq. (12):
ωia=90D500.42,
where the subscript ia refers to the discrimination between laterally stable
and meandering channels with scroll bars.
Bar regime
Bar regime was predicted by applying the relationships of Struiksma et
al. (1985) and Kleinhans and Van den Berg (2011). River bends can be seen as
an example of a perturbation to both the flow and bed sediment, which have
different adaptation lengths over which they return to equilibrium. This
difference in response is expressed by the interaction parameter (IP), which
is the ratio between the adaptation length of bed perturbation and the
adaptation length of flow. The adaptation length of flow was calculated
following Eq. (13),
λw=C2Havg2g,
and the adaptation length of a bed perturbation (m) is calculated following
Eq. (14):
λs=Havgπ2WHavg2f(θ),
where f(θ) is the magnitude of the transverse slope effect (–)
calculated following Eq. (15) (Talmon et al., 1995).
fθ=9D50Havg0.3θ
Here, θ is the dimensionless shear stress (–) calculated following
Eq. (16):
θ=τρs-ρgD50,
where τ is the shear stress (Pa) calculated following Eq. (17).
τ=ρgRSc
The interaction parameter (IP, –) was calculated following Eq. (18) to
determine the bar regime for the historical and prehistorical Overijsselse
Vecht and for comparison with the theoretical thresholds of bar regime
(Crosato and Mosselman, 2009; Struiksma et al., 1985) by
IP=λsλw.
The IP is strongly related to the width–depth ratio and was therefore
separately calculated for the meander bends Junnerkoeland and Prathoek. A low
IP means that when a bar forms in response to a local perturbation, such as
local curvature, the bar disappears within a short distance of the
perturbation (Struiksma et al., 1985). This is called an overdamped regime
and occurs in channels with a low width–depth ratio. The threshold
between overdamped and underdamped can be calculated following Eq. (19):
IP≤2n+1+22n-2,
where n is the degree of non-linearity of sediment transport versus
depth-averaged flow velocity (–). Following Crosato and Mosselman (2009) we
chose n=4, which corresponds to values for a sand-bed river. A higher IP,
and hence a higher width–depth ratio, results in an underdamped regime
associated with bars that also form further downstream of the perturbation.
The thresholds can be calculated following Eq. (20).
2n+1+22n-2<IP<2n-3
Example of a ground-penetrating radar (GPR) profile (250 Hz) in the
Prathoek bend. (a) Original GPR profile and (b) interpreted
GPR profile with lateral accretion surfaces and the channel lag, indicated by
yellow lines.
Errors and uncertainty
The above described calculations (Eqs. 1 to 11 and 13 to 20) were run
10 000 times to take into account the random errors of the input parameters,
following a stochastic approach by using Monte Carlo simulations. The
uncertainty of these parameters was described above, relating to the
transverse bed slope, bankfull depth of the meanders, valley slope, porosity,
grain size, intermittency, and the measured channel dimensions of
Palaeochannels X and Q. Systematic errors were not taken into account
because the palaeohydrological reconstruction was meant to distinguish
relative differences between fluvial phases, rather than reconstructing
absolute hydrological parameters. Results are plotted with average values
from the Monte Carlo simulations when normally distributed, or median values
when not-normally distributed, including the 16th and 84th quantile
representing the uncertainty margin. All formulas and example data used are
made available in the Supplement.
ResultsLithogenetic units
Several lithogenic units were distinguished (Fig. 4), following similar
interpretations of the sedimentary units as Huisink (2000). The descriptions
of the lithogenic units are summarized in Table 1. The cover sand deposits
were sometimes difficult to distinguish in borehole descriptions from the
fluvioperiglacial deposits when the latter had a relatively fine grain size.
Because our interest is in the delineation of the scroll bar and residual
channel-fill deposits, we combined both the fluvioperiglacial and cover sand
deposits into one unit. The fining upward sequence within the scroll bar
deposits (Table 1) can be recognized in the grain size analysis done for the
scroll bar deposits at Junnerkoeland and Prathoek (Fig. 6). The
depth-averaged grain size for both scroll bar complexes is 0.28±0.05mm. Commonly, at the base of the scroll bar deposits, a sharp
transition occurs to the brightly coloured substratum of fluvioperiglacial
deposits below, which lack organic material (Table 1). Cores that did not
reach the fluvioperiglacial deposits below the scroll bar deposits indirectly
indicate the boundary between these units because strongly consolidated
layers are present in the fluvioperiglacial deposits that were difficult to
core into. An example of a consolidated clay layer can be found directly
below the southern part of the scroll bar deposits at Prathoek (Fig. 4b).
Description of lithogenic units.
FluvioperiglacialCover sandOther channelResidual channel-fillScroll barDrift sanddepositsdepositsdepositsdepositsdepositsdepositsLithologyMod. sort. 75–2000 µmWell sort. 75–21 µmMod. sort.Sandy peat or peaty sandMod. sort. 75–600 µmWell sort.Lenses of loam andLoamy sand105–600 µmLenses of sand, silty clayLoamy sand near surface75–210 µmloamy sandloam or clay loamColourLight greyLight grey–brownLight grey–brownDark brown or blackLight brownGreyish brownto brownor whiteto dark greyGravel (%)0–20< 1< 1< 1< 40< 1Plant remainsMostly absentNoneSporadicallyAbundantFragmented andRarenear bottomabundant near bottomThickness (m)> 2< 24–54–54–51–5Width (m)> 1000> 1000< 10020–40> 10010–100Bedscm to dm thickNoneNoneNonecm to dm thickNoneAdditionalPalaeo-podzolSlightly coarserMay be poorlyFining upward, lateralMicro-podzolin topnear bottompreservedaccretion surfaces (GPR)in top
The GPR profiles clearly show the lateral accretion surfaces of the scroll
bar deposits (see example in Fig. 5). The GPR results were poor only where
the scroll bar deposits are relatively loamy or clayey on top (i.e. northern
parts of Prathoek and Junnerkoeland). The bottom of the scroll bar deposits
is mostly unrecognizable because of a low GPR reflection at this depth. In
Fig. 5 the bottom of the scroll bars is visible because this part is located
in the southern part of Prathoek where the above-mentioned clay layer was
present (Fig. 4), which caused a strong reflection of the GPR signal. The
well-preserved Palaeochannel X is a relatively symmetrical palaeochannel
(Fig. 4e) similar to Palaeochannel Q of Huisink (2000) (Fig. 2). The outer
bank consists of Weichselian–Early Holocene deposits (Fig. 4c). The average
grain size of the Palaeochannel X bed sediments is 0.23±0.12mm. No lateral accretion surfaces can be observed in the GPR
profile that was placed along the centreline of the Palaeochannel X bend
(Fig. 1e and Supplement).
Cumulative grain size distributions of the scroll bar deposits in
(a) Junnerkoeland and (b) Prathoek. Three series were made
for Junnerkoeland and two for Prathoek, each indicated by a different line
type. Each sample within a series is indicated by a different grey tone. The
averages of D16, D50, and D84 are plotted. Figure 1c and d
indicate the locations of the grain size samples.
Dating results
The channel deposits on the inside of Palaeochannel X date from 850–320
and 1408–918 BCE. Palaeochannel X was cut off at 739–117 BCE (Figs. 1e,
4c, e and Table 2).
OSL and 14C dating results from Palaeochannel X. Locations are indicated in Fig. 1c, d and Fig. 4c.
The reconstructed transverse bed slopes do not show a spatial trend
(Fig. 7a, b), and hence the mean and standard deviations were used in the
palaeohydrological calculations. The transverse bed slope at Prathoek is
higher (4.5±1.0∘) than at Junnerkoeland (3.3±1.3∘), but much lower than the transverse bed slope of Palaeochannel
X (16.9±1.9∘) and of Palaeochannel Q (28.8±3.8∘). The age as a function of distance of lateral accretion follows
from Fig. 7c, d. This relation was used for the meander and channel geometry
calculations (Fig. 8). The bankfull depths of Palaeochannels X and Q are
comparable to the bankfull depths of the meanders Prathoek and Junnerkoeland
ca. 1500 CE (Fig. 8a, b) (3 to 4 m). The bankfull depths at Junnerkoeland
decreased relatively fast ca. 1800 CE because the erosive base elevation
rises towards the cut-off channel (Fig. 4c). At Prathoek, the bankfull depth
decreased more gradually over time. The reconstructed bankfull width of
Palaeochannels X and Q is much lower compared to the meandering phase
(Fig. 8c, d), resulting in a relatively small cross-sectional area of
Palaeochannels X and Q (Fig. 8e, f).
Transverse bed slope derived from GPR cross sections from the inner
point bar to the outer bend for Junnerkoeland (a, c) and Prathoek
(b, d) as well as lateral migration distance plotted against age for
both bends. Panels (a) and (b) show the transverse bed slope of
lateral accretion surfaces measured in the GPR profile (example in Fig. 4),
including the mean and standard deviation of all measurements. Panels
(c) and (d) show the relation between age and migration
distance of the bends. Shading indicates standard deviation of the Bayesian
deposition model determined by Quik and Wallinga (2018) for the OSL and
historical map dates.
Palaeohydrology
The reconstructed Qbf is 3 to 9 times higher at the start
of the meandering phase (85–194 m3 s-1) compared to the
preceding phase represented by Palaeochannels X and Q
(19–32 m3 s-1) (Fig. 9). The difference in Qbf
between 400 BCE and 1500 CE is significant despite the relatively large
uncertainty. A similar discharge in 400 BCE compared to 1500 CE would
require a cross-sectional area 5 times larger than currently estimated
(Fig. 8e) or a 50 times higher valley slope, which falls outside the
uncertainty ranges of these parameters. The Qbf eventually
declines over time and drops to 32–70 m3 s-1 ca. 1850 CE. The
calculated Chézy coefficients for the meandering phase (47.5±0.9m0.5s-1; Eqs. 7 and 8) were comparable to average
Chézy coefficients derived from 12 low-energy rivers (44.8±13m0.5s-1) from the river dataset by Kleinhans and
Van den Berg (2011).
Combining the frequency of each discharge interval with the sediment
transport rate (Fig. 10a) results in a histogram of the sediment transport
contribution as a function of discharge (Qs,freq, Fig. 10b). The
highest measured discharge at the gauging station Mariënberg between 1995
and 2015 is 185.5 m3 s-1. The most frequent discharge occurring
in the channelized Overijsselse Vecht is 0 to 10 m3 s-1, with a
frequency of 8.2 % (Fig. 10a). When discharge is still below bankfull,
sediment transport increases relatively fast with an increasing discharge.
Above bankfull, additional discharge largely flows across the more
flow-resistant floodplain, and hence the sediment transport rates increase
less. The effective discharge (Qeff) is 29 m3 s-1,
represented by the highest sediment transport contribution (Fig. 10a, b).
Reconstructed meander
and channel geometry over time, assuming the date–distance relations (see
Fig. 7c, d) over the scroll bars. Panels (a) and (b) show
the bankfull depth (Hbf) derived from the coring data taken from
the bottom of the channel lag to the inferred bankfull water surface
(Fig. 4c, d), for both the Junnerkoeland (a, c, e) and Prathoek (b, d, f). Panels (c) and (d) show the bankfull width derived from the bankfull depth and reconstructed transverse bed slope
(Eq. 1). The river width data from Wolfert and Maas (2007) observed on
historical maps and the bankfull river width data from Staring and
Stieltjes (1848) were included for comparison. Panels (e) and
(f) show the cross-sectional area derived from the bankfull width
and water depth (Eq. 2). Shading indicates the 16th and 84th quantile.
X&Q refers to Palaeochannels X and Q.
Calculated sediment transport rates were higher than the inner bank growth or
scroll bar growth, suggesting the channel deposition can be explained
entirely by the reconstructed sediment transport (Fig. 10c). The
Qs,bf of the laterally stable phase was much lower than for the
meandering channels, explaining the large difference between the growth rate
of the channel deposits on the inner bank at Palaeochannel X
(7.0 m3 yr-1) and the scroll bars of Junnerkoeland and Prathoek
at the start of the meandering phase
(1.8×103m3yr-1). Both the sediment
transport and average scroll bar growth decreased during the meandering
phase.
Figure 11a shows that the river theoretically had insufficient stream power
for meandering ca. 400 BCE, and the bar regime was overdamped (Fig. 11b).
The stream power seemed sufficient for meandering ca. 1500 CE, and the bar
regime was underdamped. The potential for meandering gradually decreased
during the meandering phase and became again insufficient when the potential
specific stream power drops relatively fast ca. 1850 CE. The damping regime
also gradually decreased, but remained underdamped ca. 1850 CE.
Bankfull discharge over time, combined for Junnerkoeland and
Prathoek. Shading indicates the 16th and 84th quantile. X&Q refers to
Palaeochannels X and Q.
DiscussionLaterally stable phase
A relatively laterally stable phase existed prior to the meandering phase,
which is corroborated by the geochronological and palaeohydrological
reconstruction. Palaeochannel X formed by extremely slow channel
displacement of ca. 6 cm yr-1, assuming a constant channel
displacement rate, shown by the OSL dates taken from the channel deposits on
the inside of Palaeochannel X (Figs. 1e, 4c and Table 2). The lateral
migration rate of the Junnerkoeland meander bend was ca. 40 times higher
(Quik and Wallinga, 2018; Wolfert and Maas, 2007). The outer bank of
Palaeochannel X consists of Weichselian and Early Holocene deposits. No
Middle Holocene deposits were found in the corings (Fig. 4c, d), reflecting
the stable character of the Overijsselse Vecht during this period.
The preservation potential of deposits associated with the laterally stable
phase is likely very small. Deposits and dimensions of active channel reaches
are not preserved during the stable to meandering transition because
channel belt dimensions increase. Hence, channel reaches are only preserved
when they were cut off prior to the stable–meandering transition, e.g. due to
local perturbations. A channel cut-off probably caused Palaeochannels X and
Q of the laterally stable phase to become disconnected from the main river
before the meandering phase started. In this way these reaches escaped from
later lateral erosion during the meandering phase. Consequently, the lateral
stability of the river is not immediately evident from these preserved
channel reaches because the perturbations led to very slow channel
displacement, as was found for Palaeochannel X. However, scroll bar deposits
did not form and lateral accretion surfaces were lacking (Figs. 2, 4e and
the Supplement), showing that the displacement was not related to meandering in
which helicoidal flows cause bar formation and bank erosion at a significant
rate and all along the channel (Seminara, 2006). The laterally stable phase
lacked the potential to meander given its low position in Fig. 11a and is
characterized by an overdamped regime (Fig. 11b, c) and low sediment transport
(Fig. 10c). Consequently, the formation of bars was suppressed and the inner
bank deposition was small (Fig. 10c).
Sediment transport budgets calculated from present-day flow
conditions and from meander migration. (a) Discharge and sediment
transport characteristics of the Overijsselse Vecht derived from hourly
discharge data from 1995 to 2015 of the gauging station Mariënberg,
including the frequency of each discharge class over a year on a frequency
scale from 0 to 1 and the sediment transport as a function of discharge for a
randomly selected year (1546 CE) in the Junnerkoeland meander bend.
(b) Histogram of the sediment transport contribution as a function of
discharge. (c) The sediment transport and average scroll bar growth
over time (JK: Junnerkoeland, PH: Prathoek, X&Q: Palaeochannels X
and Q). The abbreviations Qs,freq and Qs,bf are
explained in Sect. 3.8. The inner bank growth X refers to the growth rate
of the channel deposits on the inner bank at Palaeochannel X, assuming a
constant lateral migration rate. Shading indicates the 16th and 84th
quantile.
The bend curvature is also an indication for the channel stability.
Palaeochannel X comprises a very sharp bend (RcurvW=1.4±0.2) compared to the meandering phase
(RcurvW=2.1±0.4), which is often found in
low-energy rivers in which lateral migration is limited (Candel et al., 2018;
Hickin and Nanson, 1984). Large similarities exist between the laterally
stable phase reported here and the laterally stable channels in highly
cohesive sediment on the intertidal mudflat, which are mostly laterally
stable except for some sharp bends at which bank failure and flow separation
result in very limited and local channel migration (Kleinhans et al., 2009).
Channel pattern change
The Overijsselse Vecht river changed from a laterally stable into a
meandering river. Differences in palaeohydrological conditions between the
two
phases were large enough to distinguish despite the large uncertainties in
the palaeohydrological reconstruction. Bar regime changed from an overdamped
regime into an underdamped regime (Fig. 11b, c), leading to overdeepening of
the outer bend pool and enhancement of the point bars in the inner bend
(Struiksma et al., 1985; Crosato and Mosselman, 2009; Kleinhans and Van den
Berg, 2011). The significantly higher bankfull discharge (factor of 3 to
9; Fig. 9) explains the potential to meander (Fig. 11a), the high sediment
transport, and the high scroll bar growth (Fig. 10c) at the start of the
meandering phase.
The exact moment of the channel pattern change is between the cut-off of
Palaeochannel X (400±300BCE) and the reconstructed
initiation of scroll bar formation (1504±52CE). Most likely,
the transition occurred shortly before the latter because both point bars
had a relatively similar meander start age (Figs. 2 and 4), the surrounding
floodplain is formed by Late Glacial or Early Holocene deposits (Fig. 4), and
there is no evidence of older scroll bar deposits in the vicinity of the
studied meander bends. Mature meandering river systems would always leave
traces of older scroll bar deposits, channel cut-offs, or meander scars
because these are never completely being removed by the river (Toonen et al.,
2012; Van de Lageweg et al., 2016).
The palaeohydrological reconstruction shows that the increasing bankfull
discharge likely explains the channel pattern change. The increasing bankfull
discharge may reflect an increase in annual discharge, but could also be
related to a more irregular discharge regime because the bankfull discharge
largely represents the higher discharges in a river (Dury, 1973; Wolman and
Miller, 1960). Consequently, the discharge may have been constant over a year
with low peak discharges and a relatively high base flow during the laterally
stable phase, changing into a more peaked discharge regime with a relatively
low base flow at the start of the meandering phase.
The potential for meandering with time. (a) The potential
specific stream power in a stability diagram (Eq. 11). Two discriminators were plotted
for a range of median particle sizes of the bed sediment, which is the range
of particle sizes found in the scroll bars and Palaeochannels X and Q
(Fig. 6). Panels (b) and (c) show the bar regime for both
Junnerkoeland and Prathoek, determined with the interaction parameter (IP)
(Eq. 18) and compared to the thresholds (Eqs. 19 and 20). Shading indicates
the 16th and 84th quantile. X&Q refers to Palaeochannels X and Q.
A potential cause of the discharge regime and channel pattern change may be
the climate change at the start of the Little Ice Age (14th to 19th
century) (Grove, 1988) given the overlap in time with the meandering phase
(Fig. 2c, d). Although geomorphological responses differ for each river
during the Little Ice Age, enhanced lateral migration or incision was
generally observed for most rivers in north-western Europe (Rumsby and
Macklin, 1996). The increased bankfull discharge in the Overijsselse Vecht
may have been caused by higher run-off relative to precipitation due to
reduced evapotranspiration rates and frozen soils (Rumsby and Macklin, 1996;
Van Engelen et al., 2001) and/or a higher snowfall to rainfall ratio due to
lower winter temperatures in the Netherlands and Germany (Behringer, 1999;
Lenke, 1968). Higher snowfall rates were also recorded for the United Kingdom
(Manley, 1969), where they led to more flooding during the snowmelt period
(Archer, 1992). Studies on historical observations of rivers nearby the
Overijsselse Vecht (IJssel, Elbe, Lower Rhine, and Meuse) suggested a
significantly higher flooding rate during the Little Ice Age compared to more
recent flooding rates (Glaser et al., 2010; Glaser and Stangl, 2003; Mudelsee
et al., 2003, 2004).
An additional cause for an increasing bankfull discharge may have been land
use change in the catchment (Kondolf et al., 2002), which affects the
discharge regime due to the direct relation with evapotranspiration (Fohrer
et al., 2001). For the Overijsselse Vecht catchment, peat reclamation started
in the 12th and 13th century (Gerding, 1995; Van Beek et al., 2015b) and
intensified from the 14th century onwards (Borger, 1992; Van Beek et al.,
2015a). Reclamation of peatlands partly comprised digging canals to drain
the land, and although the reclamation was mainly limited to the margins of
peatlands, the hydrological consequences were large. The margins are a
natural seal of the peat bog, with a low hydraulic conductivity compared to
the remainder of the bog, ensuring peat dome growth. Destruction of these
margins will result in drainage of the entire peat bog (Baird et al., 2008;
Van der Schaaf, 1999). After several centuries, focus shifted from peat
reclamation to exploitation by excavating large peatland areas for fuel during
the 17th and 18th century (Gerding, 1995). The largest part of the peat
has currently disappeared. Yearly average discharges in peatlands can
increase by 40 % in the Dutch climatological setting due to
evapotranspiration differences for reclaimed peat areas compared to
undisturbed peat areas (Baden and Eggelsmann, 1964; Streefkerk and Casparie,
1987; Uhden, 1967). This increase cannot fully explain the large increase in
bankfull discharge in the Overijsselse Vecht (factor of 3 to 9) because
peat covered just ca. 27 % of the Overijsselse Vecht catchment area
during the 14th century (Casparie and Streefkerk, 1992; Vos et al., 2011),
and hence the yearly average discharge of the catchment increased by ca. 11 %
due to evapotranspiration differences.
However, several studies have also shown that an increased drainage network
in peatlands resulted in higher discharge peaks with a fast discharge
response to precipitation (Conway and Millar, 1960; Holden et al., 2004,
2006; Streefkerk and Casparie, 1987). For example, the run-off to rainfall
ratio was a factor of 3 higher in a drained Irish peatland compared to an
undrained Irish peatland (Burke, 1975), which is comparable to the observed
bankfull discharge increase in the Overijsselse Vecht. Finally, canals were
not only dug for peat reclamation, but also for shipping and effective
generation of water power starting in the 11th and 12th century (Driessen
et al., 2000), which may have promoted the higher peak flows even more. New
canals resulted in a faster run-off, but also changed the watershed
delineation (Driessen et al., 2000). We conclude that both climatic and land
use changes were likely responsible for an increase in both total discharge
and peak flows, resulting in the transition of a relatively stable river to a
highly dynamic meandering system.
Meandering phase
Interestingly, the bankfull discharge declined during the meandering phase
(Fig. 9), leading to decreasing sediment transport relative to the scroll bar
growth (Fig. 10c) and insufficient potential specific stream power for
meandering after ca. 1850 CE (Fig. 11a). This decline was corroborated by
observations of river width from previous studies, which can be compared to
the reconstructed widths (Fig. 8c, d). These observations included
measurements from historical maps by Wolfert and Maas (2007) and measurements
of the bankfull river width over a large river section in 1848 CE by Staring
and Stieltjes (1848). The river width data from Wolfert and Maas (2007)
largely fall in the range of reconstructed bankfull widths at Junnerkoeland
and show a similar decreasing trend (Fig. 8c). However, the historical maps
used by them may result in large uncertainties because the water stage that
these maps represent is unknown (Quik and Wallinga, 2018). The measured
widths by Staring and Stieltjes (1848) are in line with the predicted width
at Junnerkoeland, falling within the uncertainty range. The predicted width
at Prathoek is underestimated compared to the measured widths by Wolfert and
Maas (2007) and Staring and Stieltjes (1848). This underestimation may
explain the lower cross-sectional area compared to Junnerkoeland (Fig. 8f)
and hence an underestimated bankfull discharge (Fig. 9) and potential
specific stream power (Fig. 11a).
The observed decline of bankfull discharge would suggest that the
hydrological forcing disappeared or diminished and had a temporary
character, which would fit the Little Ice Age that ended in the 19th
century as a potential cause. Consequently, it would be expected that the
channel pattern reorganized and became laterally stable again. However, the
river was still laterally migrating until channelization between 1896 and
1914 CE (Wolfert and Maas, 2007), which may be related to the presence of an
underdamped regime enhancing point bar formation in the inner bend (Fig. 11b,
c). Additionally, historical bank stability changes may have promoted the
river meandering during this period. For example, floodplains were
intensively used for cattle grazing, which may have weakened the banks,
enhancing meandering after 1850 CE (Beschta and Ripple, 2012; Trimble and
Mendel, 1995; Wolfert et al., 1996). Drift sand activity was also initiated
by intensive land use since the Late Middle Ages (Fig. 1c, d) (Koster et al.,
1993), which may have affected the bank stability.
Conclusions
We show that bankfull discharge and the associated river parameters can be
reconstructed by following a stochastic approach and through detailed
geochronological and lithological analysis of scroll bar deposits and
palaeochannels. For the Overijsselse Vecht river we demonstrate that an
increase in bankfull discharge ca. 1400 to 1500 CE resulted in a river
channel pattern change from laterally stable to meandering. Geochronological
data confirmed our hypothesis on the lateral stability of the river prior to
the meandering phase, in contrast to previous assumptions that were made of
continuous meandering during the Holocene. We show that the reconstructed
river parameters are consistent with both the laterally stable and meandering
channel pattern by applying empirical channel and bar pattern models.
Potential causes for the discharge regime changes include climate change
(Little Ice Age) and land use changes (peat reclamation, peat exploitation,
digging of canals). We conjecture that the change from laterally stable to
meandering has occurred in other rivers for which increased Holocene fluvial
activity was reported.
Example data and a calculation spreadsheet following the
presented method in this research are available in the Supplement.
The supplement related to this article is available online at: https://doi.org/10.5194/esurf-6-723-2018-supplement.
The authors contributed in the
following proportions to concept and design, data collection, analysis, and
conclusions and paper preparation: JHJC (40 %, 40 %, 80 %, 60 %), MGK
(20 %, 0 %, 10 %, 5 %), BM (20 %, 0 %, 0 %, 20 %), WZH
(0 %, 40 %, 5 %, 0 %), CQ (0 %, 20 %, 5 %, 0 %), JW (20 %,
0 %, 0 %, 15 %).
The authors declare that they have no conflict of
interest.
Acknowledgements
This research is part of the research programme RiverCare, supported by the
Netherlands Organization for Scientific Research (NWO) and the Dutch
Foundation of Applied Water Research (STOWA), and is partly funded by the
Ministry of Economic Affairs under grant number P12-14 (perspective
programme). Maarten G. Kleinhans was also supported by the NWO (grant Vici
016.140.316/13710). The paper has benefited greatly from reviews by
Peter Houben and two anonymous reviewers. The authors would like to thank the
following persons for their help with the different methods used in this
research: Joep Storms, Gerard Heuvelink, Marijn van der Meij, Wobbe
Schuurmans, Alice Versendaal, Erna Voskuilen, Aldo Bergsma, Marjolein
Gouw-Bouman, and UU students Karianne van der Werf, Sjoukje de Lange, Jip
Zinsmeister, and Pascal Born. We would also like to thank the De Roos family,
Staatsbosbeheer, and Water Board Vechtstromen for access to and inside
knowledge of the field sites. Edited by:
Andreas Lang
Reviewed by: Peter Houben and two anonymous referees
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