Largesample assessment of varying spatial resolution on the streamflow estimates of the wflow_sbm hydrological model
 ^{1}Water Resources Section, Faculty of Civil Engineering and Geosciences, Delft University of Technology, Stevinweg 1, 2628 CN Delft, the Netherlands
 ^{2}The Netherlands eScience Center, Science Park 140, 1098 XG Amsterdam, the Netherlands
 ^{3}Catchment and Urban Hydrology, Department of Inland Water Systems, Deltares, P.O. Box 177, 2600MH Delft, the Netherlands
 ^{4}Operational Water Management, Department of Inland Water Systems, Deltares, P.O. Box 177, 2600MH Delft, the Netherlands
 ^{5}Hydrology and Quantitative Water Management Group, Wageningen University and Research, P.O. Box 47, 6700AA Wageningen, the Netherlands
 ^{6}Applied Research Center, Florida International University, FL 33174 Miami, United States of America
 ^{1}Water Resources Section, Faculty of Civil Engineering and Geosciences, Delft University of Technology, Stevinweg 1, 2628 CN Delft, the Netherlands
 ^{2}The Netherlands eScience Center, Science Park 140, 1098 XG Amsterdam, the Netherlands
 ^{3}Catchment and Urban Hydrology, Department of Inland Water Systems, Deltares, P.O. Box 177, 2600MH Delft, the Netherlands
 ^{4}Operational Water Management, Department of Inland Water Systems, Deltares, P.O. Box 177, 2600MH Delft, the Netherlands
 ^{5}Hydrology and Quantitative Water Management Group, Wageningen University and Research, P.O. Box 47, 6700AA Wageningen, the Netherlands
 ^{6}Applied Research Center, Florida International University, FL 33174 Miami, United States of America
Correspondence: Jerom P. M. Aerts (j.p.m.aerts@tudelft.nl)
Hide author detailsCorrespondence: Jerom P. M. Aerts (j.p.m.aerts@tudelft.nl)
Distributed hydrological modelling moves into the realm of hyperresolution modelling. This results in a plethora of scalingrelated challenges that remain unsolved. To the user, in light of model result interpretation, finerresolution output might imply an increase in understanding of the complex interplay of heterogeneity within the hydrological system. Here we investigate spatial scaling in the form of varying spatial resolution by evaluating the streamflow estimates of the distributed wflow_sbm hydrological model based on 454 basins from the largesample CAMELS data set. Model instances are derived at three spatial resolutions, namely 3 km, 1 km, and 200 m. The results show that a finer spatial resolution does not necessarily lead to better streamflow estimates at the basin outlet. Statistical testing of the objective function distributions (Kling–Gupta efficiency (KGE) score) of the three model instances resulted in only a statistical difference between the 3 km and 200 m streamflow estimates. However, an assessment of sampling uncertainty shows high uncertainties surrounding the KGE score throughout the domain. This makes the conclusion based on the statistical testing inconclusive. The results do indicate strong locality in the differences between model instances expressed by differences in KGE scores of on average 0.22 with values larger than 0.5. The results of this study open up research paths that can investigate the changes in flux and state partitioning due to spatial scaling. This will help to further understand the challenges that need to be resolved for hyperresolution hydrological modelling.
Hydrological model development follows competing model philosophies (Hrachowitz and Clark, 2017). From one end of the spectrum to the other these include highresolution, smallscale, processresolving distributed models and spatially lumped conceptual models. All model structures developed within the competing philosophies have their own limitations and advantages. There are overlapping challenges that all models face including parameter estimation and the representation of spatial heterogeneity (Clark et al., 2017).
The parameter identifiability problem stems from the inability to obtain unique and realistic parameters at the modelling scale due to structural model deficiencies and applied calibration techniques (Sorooshian and Gupta, 1983). Multiple studies have extensively researched calibration techniques to overcome the parameter identifiability problem (e.g. Sorooshian and Gupta, 1983; Vrugt et al., 2002; Guse et al., 2020). The identifiability problem is emphasized in distributed modelling, the focus of this study, by the limitation of parameter measurements not being compatible with the modelling scale (Grayson et al., 1992). This results in the need for transferring parameters in space and time. Multiple studies have looked into parameter transferability (e.g. Finnerty et al., 1997; Haddeland et al., 2006; Wagener and Wheater, 2006). Melsen et al. (2016) discussed that the inadequacy of transferring parameters in space and time may indicate a lack in spatial heterogeneity and temporal variability representation in the models. Methods such as the multiscale parameter regionalization (MPR) technique (Samaniego et al., 2010) emerged to increase the transferability of parameters from the data resolution to the hydrological model resolution. Imhoff et al. (2020) used a different method that applied pedotransfer functions (PTFs) to derive parameters at the finest available data resolution and upscale them to the various spatial model resolutions of the wflow_sbm hydrological model.
The effects of spatial heterogeneity has been studied at a catchment scale using the representative elementary watershed (REW) theory developed by Wood et al. (1988); Reggiani et al. (1998, 1999); Reggiani and Schellekens (2003). In Wood et al. (1988) the basin was divided in subbasins and then aggregated to basin level to study scaling behaviour. This type of research is still very relevant as in recent years, the hydrologic modelling community has surpassed the REW scale threshold of 1 km^{2} with the move towards so called hyperresolution modelling. The discussion following this move revealed that in addition to the many benefits, e.g. applicability for stakeholders, there are multiple challenges to address (Wood et al., 2011; Beven and Cloke, 2012; Bierkens et al., 2015). These challenges include scaling issues (Gupta et al., 1986; Blöschl and Sivapalan, 1995) such as (1) the need to explicitly model processes that are parameterized at coarser resolutions, (2) lateral connections between compartments of the hydrological system that are averaged out or ignored at coarser resolutions, and (3) an increase in uncertainty due to lacking process and parameter knowledge as a result of insufficient data quality at finer resolutions (Bierkens et al., 2015).
The scaling issues arise when the (often unconscious) assumption is made that a hydrological model used at various spatial and temporal resolutions should estimate similar states and fluxes independent of scale. A utopian model has scaleinvariant model parameterization and hydrological process descriptions. The development of scaleinvariant hydrological models is, however, very challenging as most hydrological processes do not scale in a linear manner (e.g. Bras, 2015; Rouholahnejad Freund et al., 2020). Instead, processes at one length scale influence those at other scales (Horritt and Bates, 2001).
Due to the complex nature of scaling issues and a shifting distributed modelling climate towards hyperresolution modelling, it is important to continuously assess the effects of scaling. Without investigating what this move entails, the hydrological modelling community risks communication problems with the users of model results. To the user, in the case of spatial model resolution, the increase in the level of detail in model output might imply an increase in understanding of the complex interplay of heterogeneity within the hydrological system. We can only determine this by continuously assessing how models behave under various spatial (and temporal) resolutions.
Multiple studies have looked into spatial scaling effects by varying spatial model resolution. Booij (2005) found that increasing the spatial resolution of a semilumped HBV hydrological model only marginally increased model performance based on streamflow estimates. The coupled ParFlow–CLM model was evaluated with various gridcell sizes by Shrestha et al. (2015) and they found, among other effects, that soil moisture estimates were dependent on gridcell size. Sutanudjaja et al. (2018) introduced the transition from 30 to 5 arcmin gridcell size simulations of the distributed PCRGLOBWB model. Results showed a general increase in model performance compared to streamflow observations. However, regional scaling issues were present. In some of the basins model performance was lower at a finer spatial model resolution. This study made it apparent that a large sample of hydrological diverse basins should be considered when investigating the effects of varying spatial resolution. To our knowledge there are no studies that have looked into varying spatial resolution within the hyperresolution realm on a large sample of basins. This will be the focus of this research.
The distributed conceptual wflow simple bucket model (wflow_sbm) (Schellekens et al., 2020) utilizes highresolution data sets to derive model instances globally at varying spatial resolution. Parameter estimates are based on the work of Imhoff et al. (2020) to ensure consistency across scales. Remotely sensed soil and land cover data sets are sources for estimating parameters through PTFs (e.g. Brakensiek et al., 1984; Cosby et al., 1984). The PTFs are a collection of predictive functions, socalled super parameters (Tonkin and Doherty, 2005), derived at point scale that estimate soil parameters where underlying data are scarce. For most wflow_sbm model parameters, a priori parameters are available. This is not (yet) the case for the horizontal conductivity fraction (KsatHorFrac) parameter, making it a logical parameter for calibration as it is also one of the more sensitive parameters in the model (Imhoff et al., 2020). The flexible setup of wflow_sbm can be used to assess spatial scaling issues due to quasiscale invariant parameters whilst maintaining similar hydrological process descriptions across scales. This setup includes the recent improvements by Eilander et al. (2021), who developed a scaleinvariant method for upscaling river networks (one of the suggested causes of the inconsistent streamflow estimates across scales as shown in Imhoff et al., 2020).
In this study we quantify the effects of varying spatial resolution on the wflow_sbm streamflow estimates for a large sample of hydrological diverse basins in the CAMELS data set (Newman et al., 2015; Addor et al., 2017). By conducting this research on a large sample of basins, we can assess the results on consistency and locality. The assessment is conducted by creating three model instances at varying spatial resolutions for each basin: a 3 km, 1 km, and 200 m spatial grid resolution. These instances cover a broad range of large and smallscale dynamics, for example, snow accumulation at the mountain range (>1 km) and the mountain ridge scale (<1 km) (e.g. Houze Jr., 2012; Mott et al., 2018; Vionnet et al., 2021), or closing in on the hillslope scale (<100 m) (e.g. Trompvan Meerveld and McDonnell, 2006; Fan et al., 2019). The parameters for the wflow_sbm model instances are estimated at the highest available data resolution and aggregated to the modelling grid using the upscale rules as defined in Imhoff et al. (2020).
Our hypothesis is that the differences in streamflow estimates at various spatial resolutions will be small due to the parameters being quasiscale invariant and hydrological process descriptions in the model remaining the same across spatial scales. We will reject this hypothesis when the results show significantly different streamflow estimates across the studied resolutions. Additionally, we will showcase how the eWaterCycle platform (Hut et al., 2021) can be utilized for computational intensive largesample modelling studies.
2.1 Input data
2.1.1 The CAMELS data set
The CAMELS data set is a collection of hydrologically relevant data on 671 basins located in the contiguous United States (CONUS) (Newman et al., 2015). The basins were selected based on a minimum amount of human influence on the hydrological system, e.g. the absence of large reservoirs. The data set includes 20 years of continuous streamflow records from 1990 to 2009 from the United States Geological Survey (USGS). The CAMELS data set covers a hydrologically and hydroclimatologically diverse selection of basins. The samplesize, hydrological diversity of basins, and common use of this data set in other hydrological modelling studies (e.g. Knoben et al., 2020; Gauch et al., 2021) are the reasons for selecting this case study area.
Of the 671 basins, we ran 567 basins successfully for each of the three model instances (i.e. 3 km, 1 km and 200 m resolution). Reasons for excluding basins in our analysis are missing streamflow observations (7 basins) or errors during parameter estimation (97 basins). Parameter estimation errors occurred mainly during drainage network delineation, when either the basin outlet consisted of a single grid cell that results in a model coding error or inconsistencies occurred in the local drainage direction layer. When a single model instance of the three model instances failed, the basin was excluded from further analysis. Figure 1 shows the locations of the included and excluded basins as well as the reason for exclusion.
2.1.2 Streamflow observations
The USGS streamflow observation records were downloaded to match our model simulation period from 1996 to 2016. The data are resampled to daily data and the units were converted to m^{3} s^{−1}. We ensured consistency in time zones between the observation and model simulations by matching the USGS streamflow data with the UTC time zone. The tooling used for downloading, resampling, unit conversion, and shifting of time zones might be of interest to others in the hydrological community and is available in the GitHub repository (https://github.com/jeromaerts/eWaterCycle_example_notebooks, last access: 8 June 2022).
2.1.3 Meteorological input and preprocessing
The meteorological input requirements of the wflow_sbm model are precipitation, temperature, and potential evapotranspiration. Precipitation data were obtained from the MultiSource WeightedEnsemble Precipitation (MSWEP) Version 2.1 (Beck et al., 2019). The data set was constructed using biascorrected gauge, satellite, and reanalyses data. The data are available at 0.1 degrees spatial (11 km) and 3 h temporal resolution for the period 1979–2017. The temperature variable was obtained from the ERA5 reanalyses data set (Hersbach et al., 2020). The data are available at 0.25^{∘} (31 km) spatial and 1 h temporal resolution. In addition to temperature, we used ERA5 variables to calculate potential evapotranspiration using the De Bruin method (Bruin et al., 2016).
We conducted a preliminary analysis for six basins in which we compared model simulations based on streamflow estimates that use ERA5 precipitation to those that use MSWEP precipitation. Results indicated that the use of ERA5 precipitation did not produce desirable streamflow estimates compared to MSWEP precipitation. Switching to the MSWEP precipitation product improved streamflow estimates throughout the case study area. Figure A1 and Table A1 in Appendix A contain the results of this analysis. Noticeably, some of the basins are very sensitive to changes due to different forcing data sets as shown by the streamflowbased objective functions.
The meteorological input is preprocessed within the eWaterCycle platform using the Earth system model evaluation tool (ESMValTool) Version 2.0 (Righi et al., 2020; Weigel et al., 2021). Before further processing, the data are aggregated to daily values. The precipitation variable is disaggregated to the modelling grid using the secondorder conservative method to ensure consistency of the total volume of precipitation across spatial scales. The temperature variable is disaggregated with the environmental lapse rate and the digital elevation model (DEM) used by the hydrological model. The variables required by the De Bruin method (Bruin et al., 2016) are disaggregated using the (bi)linear method and are subsequently used to calculate potential evapotranspiration. The code used for preprocessing is included in the Jupyter Notebooks made available with this paper (https://doi.org/10.5281/zenodo.5724512).
Yamazaki et al. (2019)Lehner et al. (2011)Messager et al. (2016)Pfeffer et al. (2014)Karger et al. (2017)Kottek et al. (2006)Buchhorn et al. (2020)Myneni et al. (2015)Hengl et al. (2017)2.1.4 Parameter estimation from external data sources
The parameter sets used in this study were derived using the hydroMT software package (Eilander and Boisgontier, 2021). The data sources for deriving parameter sets are opensource global data sets. These include topography, surface water, land cover and land use, soil, meteorology, and river gauge data. The PTF to estimate soil properties is based on Brakensiek et al. (1984). In 25 of the 567 basins, lakes and/or reservoirs were included in the model parameters given a threshold area of 1 and 10 km^{2}, respectively.
An overview of the data and references are provided in Table 1.
2.2 Model experiment setup
2.2.1 The wflow_sbm model (v.2020.1.2)
The wflow_sbm model is available as part of the wflow opensource modelling framework (Schellekens et al., 2020), which is based on PCRaster (Karssenberg et al., 2010) and Python. Figure 2 shows the different processes and fluxes that are part of the wflow_sbm hydrological concept. The soil part of the wflow_sbm model is largely based on the Topog_SBM model (Vertessy and Elsenbeer, 1999), which regards the soil as a “bucket” with a saturated and unsaturated store. This model was developed for simulating smallscale hydrology. For channel, overland and lateral subsurface flow, a kinematic wave approach is used, similar to TOPKAPI (Benning, 1994; Ciarapica and Todini, 2002), G2G (Bell et al., 2007), 1KDHM (Tanaka and Tachikawa, 2015), and Topog_SBM (Vertessy and Elsenbeer, 1999). The wflow_sbm model has a simplified physical basis with parameters that represent physical characteristics, leading to (theoretically) an easy linkage of the parameters to actual physical properties. The Topog_SBM model is mainly used to simulate fast runoff processes during discrete storm events in small basins (<10 km^{2}) (evapotranspiration losses are ignored). Since evapotranspiration losses and capillary rise were added to wflow_sbm, the derived wflow_sbm approach can be applied to a wider variety of basins. The main differences of wflow_sbm with Topog_SBM include

the addition of evapotranspiration and interception losses using the Gash model (Gash, 1979) on daily time steps or a modified Rutter model on subdaily time steps (Rutter et al., 1971, 1975);

the addition of a root water uptake reduction function (Feddes et al., 1978);

the addition of capillary rise;

the addition of glacier and snow buildup and melting processes;

wflow_sbm that routes water over an eight direction (D8) network, instead of the element network based on contour lines and trajectories, used by Topog_SBM;

the option to divide the soil column into any number of different layers;

vertical transfer of water that is controlled by the saturated hydraulic conductivity at the water table or bottom of a layer, the relative saturation of the layer, and a power coefficient depending on the soil texture (Brooks and Corey, 1964).
2.2.2 Model runs and calibration
We derived three model instances at varying spatial model resolutions that cover a 3 km, 1 km, and 200 m grid. While for most parameters of the wflow_sbm model a priori estimates can be derived from external sources, a single nondistributed parameter needs to be calibrated for each basin: the saturated horizontal conductivity often expressed as a fraction (KsatHorFrac) of the vertical conductivity. This parameter cannot be derived from external data sources because it compensates for anisotropy, unrepresentative point measurements of the saturated vertical conductivity, and model resolution (Schellekens et al., 2020). A sensitivity analysis conducted by the model developers concluded that KsatHorFrac is the most effective parameter when it comes to calibration based on streamflow estimates, also briefly discussed in Imhoff et al. (2020). Increasing the value of this parameter leads to an increased base flow component and reduced peak flow and flashiness.
We calibrated the models to match model setups of those used by the users of the hydrological model. The model instances are calibrated using the modified Kling–Gupta efficiency score (KGE) (Gupta et al., 2009; Kling et al., 2012) by comparing the simulated streamflow estimates with the streamflow observations at the basin outlet. Twentyone runs are evaluated based on an interval of KsatHorFrac values ranging between 1 and 1000. The bestperforming model run and its corresponding KsatHorFrac parameter are selected for further analyses during the evaluation period. The model calibration is conducted from 1997 to 2005 and the model evaluation from 2007 to 2016. The years 1996 and 2006 are regarded as spinup years and are not included in the analysis.
2.3 Benchmark selection
To select basins with reasonably good model performance, we applied a statistical benchmark to beat. The use of a benchmark allows for better interpretation of objective functionbased results (Garrick et al., 1978; Pappenberger et al., 2015; Schaefli and Gupta, 2007; Seibert, 2001; Seibert et al., 2018; Knoben et al., 2020). We adopt, in part, the same methodology for statistical benchmark creation as Knoben et al. (2020). The benchmark is created by calculating the mean and median of streamflow observations per calendar day of the 10year evaluation period and the results are compared to streamflow predictions from the model. The KGEs are calculated for the observed streamflow versus these mean and median values. The mean is expected to better represent larger basins with more stable flow regimes whilst the median better covers the flashiness of smaller headwater basins. This benchmark serves as a lower boundary for the model predictions: if the model, for any of the three resolutions has a lower KGE than either the median or mean flow benchmark, it is considered to be unsuitable for this study and removed from further analyses.
2.4 Analysis of results
2.4.1 Objective function and sampling uncertainty
We use the modified KGE metric for the analysis of results. Ideal model performance has a KGE score of 1 and a KGE score of −0.41 is equal to taking the mean flow as a benchmark (Knoben et al., 2019c). The lowest value of the three components determines the KGE score. The KGE score and its components are as follows:
where r is the Pearson correlation, β the mean bias, and γ the variability bias. The mean and standard deviation of the simulations is denoted as μ_{s} and σ_{s}, respectively and μ_{o} and σ_{o} are the mean and standard deviation of the observations.
We quantify the sampling uncertainty of the KGE score for the selected basins based on the statistical benchmark following the methodology of Clark et al. (2021). This applies the nonoverlapping bootstrap method (Efron and Tibshirani, 1986) to calculate tolerance intervals and jackknifeafterbootstrap method (Efron, 1992) for the standard error calculation of those tolerance intervals. This method applies bootstrap and jackknife methods to estimate the standard errors and tolerance interval of KGE uncertainty. The tolerance interval is defined as the difference between the 5th and 95th percentiles. We ran the gumboot analysis package (Clark et al., 2021) with a sample size of 500 during the evaluation period to calculate the results.
2.4.2 Comparison of streamflow estimates
To provide more context for the results in terms of general model performance, we compared the streamflow estimates from wflow_sbm to those of the study by Knoben et al. (2020). The latter ran 36 conceptual models on the CAMELS data set using the Modular Assessment of Rainfall–Runoff Models Toolbox (MARRMoT) Version 1.0 (Knoben et al., 2019a, b) . First, we calculated the mean of the 36 models for each basin. Next, we ensured a match between the basins under investigation by both studies. Due to differences in time period, forcing, and numerical solvers, the results cannot be compared directly to those of this study. It does however provide context for the results.
The intermodel (instance) comparison of the streamflow estimates in this study is assessed using a cumulative distribution function (CDF). We applied the Kolmogorov–Smirnov (KS) test (Kolmogorov, 1933; Smirnov, 1933) to assess whether the differences between the KGE score distributions of the model instances are statistically relevant. This allows the acceptance or rejection of the hypothesis stating that the differences in streamflow estimates at various spatial resolutions will be small.
2.5 eWaterCycle platform
This research was conducted within the eWaterCycle platform (Hut et al., 2021). By design, eWaterCycle follows the FAIR (Findable, Accessible, Interoperable, and Reusable) principles of data science (Wilkinson et al., 2018) and allows highlevel communication with models regardless of programming language through the Basic Model Interface (BMI; Hutton et al., 2020). This study showcases the way eWaterCycle handles the setup of extensive modelling studies. A Jupyter Notebook with the model experiments of this study is provided in the GitHub repository. Since notebooks are not ideal for longrunning experiments on highperformance computing (HPC) machines, we exported the notebooks to regular python code which we ran directly on an HPC. The calibration and evaluation procedures totalled 41.025 model runs on the Dutch national supercomputer Cartesius hosted by Surf.
The results in this section are based on the modified KGE 2012 objective function applied to the streamflow estimates at the basin outlet. The Nash–Sutcliffe efficiency (NSE) (Nash and Sutcliffe, 1970) results are available in the repository (https://doi.org/10.5281/zenodo.5724576).
3.1 Calibration period results
3.1.1 The effect of calibration on streamflow estimates
To illustrate how model calibration affects the streamflow estimates of each model instance, we first show the calibration curves of a single basin (ID:14301000). We selected a basin with moderate performance and only show the last year of calibration to avoid presentation bias. Figure 3a–c shows the calibration curves (yellow to red) for each of the instances that were generated by tuning the horizontal conductivity (i.e. KsatHorFrac) parameter. The parameter values range from 1 to 1000 and are a single value for each basin. For visualization purposes, not all calibration interval values are shown in the figure. The results depict the effect that increasing the KsatHorFrac values has on the hydrograph: base flow increases while peak flow reduces. In addition, large KsatHorFrac values reduce the flashiness of the streamflow estimates as is visible in Fig. 3b and c in the second week of November 2005. It is worth noting that the selection of best calibration parameter values is strongly dependent on the chosen objective function as the NSE score would be more favourable towards flashiness and less towards base flow, for example. As shown in Fig. 3d, the streamflow estimates of the model instance are similar (KGE score 0.58–0.66) while the parameter values deviate (KsatHorFrac 125–1000). No strong apparent trends for KsatHorFrac values in relation to model resolution or geographic location were found after calibration.
Figure 4 shows the CDFs of the KGE score distribution based on the bestperforming calibration model run of the three model instances. Starting with the modified KGE score distributions in Fig. 4a, we find that the model instance distributions are very similar for the entire domain, especially the 1 km and 200 m instances. The 3 km instance has lower scores for 60 % of the distributions between 0.2 and 0.8 CDF. Approximately 18 % of the distributions is lower than a KGE score of −0.41 which corresponds to taking the mean flow. All three model instances show a similar Pearson r correlation CDF (Fig. 4b) and gamma variability ratio (Fig. 4c). The largest differences are visible in the beta bias component (Fig. 4d) that shows similar bias for the 1 km and 200 m instances with 60 % of the distributions being lower than a value of 1 and 40 % being higher. The highest agreement is in the lower and upper 5 % of the distributions. The 3 km instance only agrees with the 1 km and 200 m instances in the upper 10 % of the distribution where 70 % of the 3 km distribution is lower than a value of 1 and 30 % is higher. The bias term of the KGE score has the largest weight in determining the shape of the KGE score CDF as is shown by 20 % of the distributions with bias values larger than 2.0 .
3.2 Evaluation period results
3.2.1 Benchmark selection
The statistical benchmark is applied to determine which basins contain the streamflow estimates of the model instances that are deemed adequate for further analyses. The statistical benchmark is based on the bestperforming type of climatology per calendar day, either mean or median, during the model evaluation period. Figure 5a spatially shows which type of calendarday benchmark is performing best per basin. Of the 567 simulated basins, the results of 454 basins exceed the benchmark for each model instance. This is the case for all basins in the Midwest of the United States. Poor performance in comparison to the benchmark is mainly present in the Southwest. Based on the KGE scores, 83 % of the benchmark is favourable towards using the climatological mean and 27 % towards the median. An overview is provided in Fig. 5c and the distribution of the benchmark KGE scores is shown in Fig. 5d. The distribution ranges from a −0.62 to 0.71 KGE score and is skewed towards values lower than 0.0 with a mean of 0.02 and median of −0.02.
The statistical benchmarks (mean and median during the evaluation period) are plotted as CDFs based on the KGE score and its three individual components in Fig. 6. For most of the KGE distributions in Fig. 6a, the mean benchmark outperforms the median with the upper 10 % being the exception. The Pearson correlation coefficient (Fig. 6b) is slightly higher for the median at 70 % of the distributions. In the case of the gamma variability ratio (Fig. 6c), the bottom 38 % is lower for the median than the mean, and the upper 55 % is slightly higher for the median than the mean. The determining component of the KGE score is the beta component shown in Fig. 6d. As expected, the mean benchmark is, for the most part, close to a value of 1, meaning that it is close to the mean of the observations. Of interest are the points of the median that are not close to 1, i.e. the lower 10 % and the upper 1 % of the distributions. These basins have flow regimes that greatly differ per year compared to the climatology. Considering the mean statistical benchmark of basins with a beta bias score lower than <0.75 and higher than 1.75, we find that the model instances outperform the statistical benchmark in 35 of the 44 basins.
3.2.2 The effect of spatial resolution on streamflow estimates
As with the calibration period, the same three example basins are used to illustrate the differences in streamflow estimates between model instances for the evaluation period. Only the last year of the evaluation period is shown in Fig. 7.
In the case of poor performance (Fig. 7a), it may occur that the model instances are overestimating streamflow during peak flow. The bestperforming model instance has the smallest peakflow estimates which is the coarsest spatial resolution instance (3 km) in many cases. Of the 454 basins, the 3 km instance has the lowest peakflow estimates in 279 occurrences of which the model performs best 148 times. The example in Fig. 7b shows that best performance of the 1 km model instance often occurs in conjunction with a relatively similar performance of the 200 m model instance. In 78.7 % of the cases in which the 1 km instance is best performing, the difference in KGE score with the 200 m instance is smaller than 0.1. The final example in Fig. 7c illustrates when the finest spatial model resolution (200 m) is best performing and the coarsest (3 km) least performing. The 200 m model instance best captures peak flow and the receding limbs of the hydrograph.
3.2.3 Streamflow estimates of model instances
The KGE score results for the evaluation period are shown in Fig. 8a. The KGE score distribution of the mean of 36 hydrological models from Knoben et al. (2020) is included and referred to as “MARRMoT mean results”. It is worth noting that the comparison between studies is not oneonone due to differences in model run periods, forcing, and numerical solvers. We can, however, obtain information about general model performance between both studies.
The mean KGE score distribution of the MARRMoT models (Fig. 8, blue) of Knoben et al. (2020) is close to the mean of the distributions of the three model instances. Differences between study results are mainly present in the tails of the distributions. Below 0.17 of the CDF (worst 17 % of the results), the MARRMoT mean results of KGE score distribution is higher than the 1 km model instance. The MARRMoT mean results for the lower 5 % of the CDF performs better than the distributions of the three model instances. Here, the range of KGE scores is smaller for the MARRMoT mean (−1.55 to 0.09) than for the three model instances (−13.56 to 0.00). Above 0.17 of the CDF (83 % of the results), the distributions of the three model instances are higher in KGE scores than those of the MARRMoT mean.
When we only consider the wflow_sbm instances, approximately 64 % of the results of the model instances are higher than 0.50 KGE score, and of those, 18 % are higher than 0.75 KGE score. The distributions cross at multiple points, for example at the bottom 10 % of the distribution the 3 km instance has the highest and the 1 km the lowest KGE score. At 40 % of the distribution and lower, the 200 m instance is followed by the 1 km and 3 km instances in terms of highest KGE score. The Pearson r component of the KGE score CDF in Fig. 8b and the gamma variability ratio component in Fig. 8c show small differences between model instances. The beta bias component in Fig. 8d shows the largest differences between model instances, especially between the 3 km, 1 km, and 200 m model instances. The bias component is the main factor for the differences in the overall KGE score.
Next, we apply the KS statistic to test whether the CDF of the model instances statistically differ from each other, for a given pvalue of 0.05. The KStest results in Table 2 show that the difference between 3 km and 200 m model instances is statistically relevant at a pvalue of 0.02. On a large sample, this means that increasing the spatial model resolution from 3 km to 1 km or 1 km to 200 m does not lead to significant differences in streamflow performance. When changing resolution from 3 km to 200 m, the distribution of KGE scores is significantly different (p<0.05) according to the KStest.
3.2.4 Objective function uncertainty
In addition to the streamflow evaluation, we conducted a sampling uncertainty assessment of the KGE objective function using bootstrap and jackknifeafterbootstrap methods. The results of this assessment for each of the model instances are shown in Fig. 9. The results for the three model instances are very similar. The toleranceinterval results, denoted with the black lines, show that approximately 100 basins have a KGE sampling uncertainty of 0.1 or lower, and approximately half of the basins have values of 0.2 or lower. Half of the basins show high KGE uncertainty of more than 0.2 with approximately 80 basins surpassing 0.5 KGE for all model instances.
We project the sample uncertainty results on the CDF of the evaluation period (Fig. 8) by calculating the mean of the tolerance interval, jackknife standard error, and bootstrap standard error for each of the three model instances per quarter of the total CDF. The results in Table 3 show slightly higher results for quarters of the CDF for the tolerance interval relative to the jackknife and bootstrap standard errors. The lower tail of the CDF contains the highest average values for the three sample uncertainty statistics, and the upper tail contains the lowest average values, indicating that sample uncertainty is high at low KGE values and vice versa.
3.2.5 Spatial distribution of evaluation period results
The CDF does not provide information at a basin level. To gain insight into the spatial distribution of the KGE scores of the model instances, Fig. 10 shows the KGEs of the streamflow estimates plotted on a map of the CONUS domain. The minimum KGE scores of 0.50 to 0.89, shown in Fig. 10a, are found in the Pacific Northwest, South Atlantic, Appalachia, and northeast of CONUS. The KGE scores lower than −0.41 are found throughout CONUS. The highest KGE scores in Fig. 10b are located in the Northwest, Rocky Mountains, and Appalachia. These regions are characterized as steepsloping headwater basins. Figure 10c shows that there are large local streamflow discrepancies of more than 1.00 KGE score. These are mainly found in the Pacific Southwest, the Southwest, and the Midwest. These regions span a wide range of hydroclimatic diverse basins. The average KGE score difference is 0.22. Figure 10d shows the bestperforming model instance for each of the 454 selected basins. Although regions show clusters in bestperforming model instances, there is no overall geographical trend in results. Best performances for the 1 km model instance are generally close to basins where the 200 m model instance is performs best. The 3 km model instance shows some clusters in the Southwest and Pacific Northwest. The Rocky Mountains contains the bestperforming model instances at 200 m and the Appalachia a shows a mixture between the 1 km and 200 m model instances.
3.3 The effect of spatial scale on terrain characteristics
We illustrate the effect of spatial scaling on the parameter set of the three model instances by showing the difference in topography and drainage density for three basin. To avoid presentation bias, the basins were sampled based on poor streamflow performance (ID:06878000), moderate performance (ID:02231342), and good performance (ID:06043500). Figure 11a shows the PDF of the height distributions of the model instances for each of the three basins, Fig. 11b shows the slope distribution, and Fig. 11c the profile curvature distributions.
The height distribution of the DEM in Fig. 11a shows, most clearly for basin ID 02231342, how the representation of the highest altitudes is underestimated by the 3 km model instance (orange) compared to the 200 m instance (green), and to a lesser extent the 1 km instance (red). Essentially, at coarser spatial resolution the terrain is flattened at high altitudes. An opposite effect is shown in this basin for the lower altitudes where the finerresolution instances better capture gentle slopes that are flattened at coarse resolution. This effect is also detectable in the slope and profile curvature PDFs shown in Fig. 11b and c. As can be expected from the height distributions, the slope of the 200 m instance has more gentle and steep sloping topography than the 1 and 3 km instances. This is shown by the narrower slope distribution for the coarse spatial resolution that broadens with finer resolution. The differences in the mean slope of the basins between model instances is marginal, e.g. 0.00019 m m^{−1} for basin ID 02231342. The profile curvature in Fig. 11c indicates whether a slope is linear (values close to 0), concave (values smaller than 0), and convex (values larger than 0). The 3 and 1 km instances show similar slope geometries with the 3 km instance having slightly more linear slopes. At the finest resolution (200 m) the slope's geometry shifts from linear slopes to either convex or concave curvature profiles.
In addition to topography, we calculated the drainage density for each of the model instances defined as total river length divided by basin area. The results in Table 4 show small differences between the model instances for each of the three basins.
4.1 Benchmarks
We applied an initial statistical benchmark, based on streamflow observations for basin selection, to identify the basins in which streamflow estimates are deemed adequate for further analysis. This does not imply that excluded basins are less relevant. Instead, it implies that an indepth model assessment is required to understand why the model is not able to simulate adequate streamflow estimates in these basins. The CDFs of the KGE score and its components based on the statistical benchmarks in Fig. 6 show that the benchmark is relatively easy to beat given the KGE score distributions. We find, as one might expect, that the beta bias component has values close to 1 for the mean benchmark while this is not the case for the median benchmark. Bias values that are not close to 1 for the mean benchmark indicate that the flow regime changes from year to year. In most cases the hydrological model simulations were able to capture this change better than the benchmark. This shows that the model is able to capture yearly variability and is not overfitting due to extensive calibration. In addition to the benchmark, we added a layer of context by including results from the study by Knoben et al. (2020). This is an imperfect comparison due to differences in model inputs, numerical solvers, and simulation period. However, the results do provide information on general model performance. The results show that the wflow_sbm streamflow estimates are inline with estimates of the mean from the 36 MARRMoT models. The spread of results is smaller for the 36 MARRMoT models which is due to averaging and likely the more extensive calibration routine of the conceptual models. It also implies that when only streamflow at the basin outlet is under consideration, users should carefully consider various model structures before model selection due to the small differences in results at the system scale.
Other studies have conducted large domainmodelling efforts with CONUS as case study area (e.g. Mizukami et al., 2017; Rakovec et al., 2019). However, these are hard to compare with the results from this study as they did not use the same basins. To improve future comparative work, we advocate for the creation of model output storage guidelines that use the CAMELS data set as case study area. These guidelines should encompass the differences between hydrological models, such as distributed and nondistributed modelling grids. A first step is the inclusion of distributed data sources in the CAMELS data set, e.g. meteorological data. This can be extended by including model evaluation products such as snow cover and soil moisture for further benchmarking. We further propose the use of a model experiment environment, such as eWaterCycle (Hut et al., 2021) to generate model results. This allows for similar preprocessing of inputs, standardization of outputs, and reproducible modelling studies. An example of how to apply these steps using the eWaterCycle platform is provided in the Jupyter Notebooks that supplement this publication (https://doi.org/10.5281/zenodo.5724512). The ease of setting up a model experiment and storing output is an incentive for users to store model results while conducting extensive modelling studies, even when results are not deemed not suitable for publication yet but might still benefit the community.
4.2 Streamflow estimates and uncertainty
At the start of the study we hypothesized that differences between model instances would be small due to quasiscale invariant parameter sets and constant hydrological process descriptions within the model. The results of the calibration period in Fig. 4 and the evaluation period in Fig. 8 show that this is the case for the KGE score distributions based on streamflow estimates at the basin outlet. Although the differences are small, the crossing of the distribution lines is a strong indication that there is disagreement about KGE scores between model instances and that there is no single instance outperforming another consistently. The largest difference between distribution for both periods is found in the beta bias component of the KGE score. The benchmark selection mainly excluded the basins that contained large bias values during the calibration period. For the bias component, the 3 km instance deviates from the 1 km and 200 m instances. This shows that the total volume of streamflow differs for the 3 km instance. This difference was confirmed by testing the statistical significance of the differences between distributions of the model instances by applying the KSstatistic on the evaluation period results. Here, a statistical difference between distributions is only found for the modelinstance combination of 3 km and 200 m Table 2.
This study applies a single objective function, the modified KGE (Kling et al., 2012), to determine simulated streamflow adequacy for the model calibration and evaluation time periods. To provide the reader with more context we have included the 2009 KGE (Gupta et al., 2009), modified KGE, nonparametric KGE (Pool et al., 2018), and NSE objective function results in the data repository and in Fig. A2. The hypothesis testing based on conclusions were not affected by the type of KGE objective function. We selected the modified KGE objective function as it is less influenced by extreme combinations of simulated streamflow, observed streamflow, and structural error in the meteorological input. The single objective function for a whole period approach is limited and can be improved by first determining the objective function for individual years and then averaging it for the whole period (Fowler et al., 2018). In addition, as stated in Clark et al. (2021), it is important to determine the sampling uncertainty of objective functions to avoid incorrect conclusions at the system scale. Following their methodology, we investigated the sampling uncertainty by applying bootstrap and jackknifeafterbootstrap methods (Fig. 9). Approximately half of the basins have KGE uncertainties of 0.2 or lower based on the tolerance interval, and half of higher than 0.2 with values surpassing 0.5 KGE uncertainty. The lower 25 % of the CDF (Fig. 8 contains the highest sample uncertainty and the upper 25 % the lowest (Table 3). Given these results, we argue that the differences found using the KSstatistic are easily within the sampling uncertainty margins and therefore not valid to base conclusions on, even more so when considering the model uncertainty and the observation uncertainty. This demonstrates the drawbacks of conducting largesample assessments and the sensitivity to sampling uncertainty of the results.
4.3 Relative model instance differences
We recognize that largesample assessments obscure variations in simulations between instances due to the sample size. On a basin level we find that local variations due to spatial resolution are in effect throughout the domain. This is depicted by the differences between the KGE scores of the model instances (Fig. 10c). On average, there is a 0.22 KGE score difference with extremes of more than a 0.5 KGE score difference at multiple basins. In this case, we are interested in the relative differences between the instances to understand the effect of varying spatial resolution. According to Oreskes et al. (1994), this is the only form of validation that is actually possible. Therefore, we argue that the differences are large locally, even though they might be within the sampling uncertainty range.
We find that the 1 km instance is performs best in basins where the difference between minimum and maximum KGE score of the three instances is small (Fig. 10c). We partly attribute this to the calibration routine finding a better optimal parameter value (KsatHorFrac) since results are often close to those of the 200 m instance. For the bestperforming 3 km or 200 m model instances (Fig. 10d), there are only small geographical trends of best model performance in the South and Appalachia. This information is valuable for future research that will conduct a more indepth assessment of the internal states and fluxes because now we know where the effects of varying gridcell size is small or large and which model performs best.
We conducted a terrain analysis (Fig. 11) to identify changes in terrain characteristics due to spatial resolution that might explain the differences between streamflow estimates of model instances. Minor effects are depicted by the profile curvatures present at 200 m resolution. Slopes are less linear at fine resolution than at coarse resolution. The effect on the hydrological response is, however, expected to be small as stated by Bogaart and Troch (2006). Similarly, small changes are found for the differences in drainage density between model instances (Table 4). This confirms that the drainage network upscaling method of Eilander et al. (2021) is (almost) consistent across spatial scales.
Larger differences between model instances are found for the height distribution of the DEM, which is flattened at course resolution compared to finer resolution. This introduces changes in snow dynamics at high altitudes due to the use of the temperature degreeday method by the hydrological model. The resulting effect on streamflow estimates depends on the relative contribution of snowmelt. Although marginal at a basin level, the difference in slope between instances is expected to effect the partitioning of the lateral fluxes of the wflow_sbm model since the lateral connectivity between grid cells is slopedriven. An increase in slope would lead to larger lateral fluxes and vice versa. Increasing spatial resolution, aggregating the DEM, results in a broader distribution of slopes that affects the volume and timing of streamflow estimates. The effect of terrain smoothing has been reported by Shrestha et al. (2015) and they found this to increase overland flow lateral flux. An indepth assessment of the internal states and fluxes of the model instances is required to determine whether these components are the main cause for the differences in streamflow estimates.
We applied the same meteorological forcing products and preprocessing routine for each model instance. This ensured that the total volume of precipitation remained consistent across scales. A coarse grid cell contains a volume of precipitation that is equally redistributed over the equivalent amount in size of finer grid cells. In reality, this redistribution of water might not be equal across the finer grid cells, and thus scaling behaviour is introduced due to the locality of precipitation. This has an effect on the streamflow estimates as the locality of precipitation directly influences hydrological processes that are dominant at different locations (e.g. hill slope). Additionally, due to the large difference between native data and model instance resolution, it is likely that the effects of disaggregation of precipitation and temperature lapsing are main drivers for differences in streamflow estimates between model instances. However, Shuai et al. (2022) found that by comparing the same forcing product with native implementations at various spatial resolutions the effect on streamflow estimates was relatively small. This was not the case for distributed variables in the basin (e.g. snow water equivalent). It is of interest to investigate this effect on streamflow estimates and to determine the role of native spatial meteorological forcing resolution.
4.3.1 Computational cost
When we consider the increase in streamflowbased model performance as opposed to computational cost, we find that it does not scale linearly with the amount of grid cells in the basin due to lateral connections in the hydrological model. The average nonparallel run time of the 3 km instance is 157 s while that of the 200 m instance 12 120 s with an average gridcell number difference of 28 872 cells. These results point toward the importance of conducting an initial spatial model resolution assessment at the start of largesample assessments as it avoids subpar or computationally expensive model runs. Note that this kind of information can stimulate scientific and/or computational developments, e.g. in the meantime the wflow code was rewritten in Julia (van Verseveld et al., 2021), roughly increasing performance by a factor of 3 while other improvements (threading, mpi) are being implemented. There are alternative approaches to the spatial discretization of basins that are computationally very efficient, such as the vectorbased configurations that have the added benefit to better capture topographic details and are less influenced by native forcing resolution, for example (Gharari et al., 2020).
4.4 Outlook
The results from this study help model developers with model refinement by providing them with an understanding of where and under which circumstances differences due to spatial scaling occur. Based on the aggregated domain and basin level results we can conclude that increasing spatial resolution does not necessarily lead to better streamflow estimates at the basin outlet. The implications of the results for the user are that caution is advised when interpreting highresolution model outputs as this does not directly translate into better model performance. Moreover, the computational cost of increasing model resolution is not always warranted compared to increase in streamflow estimatebased model performance.
We conducted this study as an initial assessment to be followed up with studying scaling effects in distributed hydrological models. As the sampling uncertainty results showed, it is very hard to draw conclusions from a large sample and future research should therefore consider a smaller subset of basins to explore scaling effects in more detail. In this study, we did not investigate individual basins to avoid biased selection of case study areas. We suggest that future work investigates the basins that show large or small differences in model performance, lateral fluxes, and effects of terrain aggregation to be part of this subset. In addition, the evaluation should go beyond streamflow by using multiple evaluation data products (e.g. soil moisture, evaporation, gravitational anomaly, see (Guse et al., 2021) for a recent overview). Conclusions might then be drawn on whether increasing spatial resolution leads to increased model fidelity or not. This should be tested using multiple forcing data sets to evaluate the robustness of the conclusion. With the inclusion of multiple timescales as discussed in Melsen et al. (2016), more information can be obtained about the linearity of hydrological process descriptions in the model.
The aim of this study was to analyse the effects that varying spatial resolution has on the streamflow estimates of the distributed wflow_sbm hydrological model. Distributions of model instance KGE score results were tested for significant differences as well as the sampling uncertainty. A spatial distribution assessment was conducted to derive spatial trends from the results. The main findings of the study are the following:

The difference in the distributions of streamflow estimates of the wflow_sbm model derived at multiple spatial grid resolutions (3 km, 1 km, 200 m) is only statistically significantly different between the 3 km and 200 m model instances (p<0.05). This confirms the hypothesis at an aggregated level that differences between model instances are small due to quasiscale invariant parameter set and process descriptions that remain constant across scales in the hydrological model. However, the sampling uncertainty of the KGE score proved to be large throughout the domain. Therefore, we conclude that the minute differences found between model instances are to small to base conclusions on.

Results show large differences in maximum and minimum KGE scores with an average of 0.22 between model instances throughout CONUS. This provides valuable information for followup research based on the locality of relative model scaling effects.

There is no single bestperforming model resolution across the domain. Finer spatial resolution does not always lead to better streamflow estimates at the outlet.

Changes in terrain characteristics due to varying spatial resolution influence the lateral flux partitioning of the wflow_sbm model and might be an important cause for differences in streamflow estimates between model instances.
This study indicated where locality in results are strong due to varying spatial resolution. Future research should conduct an indepth assessment of basins where differences in streamflow estimates and lateral fluxes are large due to spatial scale. This will lead to a better understanding of why and under which conditions locality in spatial scalingrelated issues occur.
A2 CDFs of multiple objective functions
The software that supplements this study is available at https://github.com/jeromaerts/eWaterCycle_example_notebooks (last access: 5 June 2022.) or https://doi.org/10.5281/zenodo.5724512 (Aerts et al., 2021a). The data that supplement this publication are available at https://doi.org/10.5281/zenodo.5724576 (Aerts et al., 2021b).
JPMA wrote the publication. JPMA, WJvV, AHW, and PH did the conceptualization of the study. JPMA, ND, and PH developed the methodology. JPMA, WJvV, AHW, and PH conducted the analyses. RWH, NCvdG, WJvV, AHW, and PH did an internal review. RWH, ND, and NCvdG are PIs of the eWaterCycle project.
The contact author has declared that none of the authors has any competing interests.
Publisher’s note: Copernicus Publications remains neutral with regard to jurisdictional claims in published maps and institutional affiliations.
We would like to thank the anonymous reviewer and Shervan Gharari for their valuable feedback that helped to improve this manuscript. This work has received funding from the Netherlands eScience Center (NLeSC) under file number 027.017.F0. We would like to thank the research software engineers (RSEs) at NLeSC who cobuilt the eWaterCycle platform and Surf for providing computing infrastructure.
This research has been supported by the Netherlands eScience Center (grant no. 027.017.F0).
This paper was edited by Efrat Morin and reviewed by Shervan Gharari and one anonymous referee.
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