Effects of earthquake and cyclone sequencing on landsliding and fluvial sediment transfer in a mountain catchment

Earth Surface Processes and Landforms 1354 Earth Surf. Process. Landforms 33, 1354 1373 (2008) G-W. Lin et al. Published online in Wiley InterScience (www.interscience.wiley.com).1716 Effects of earthquake and cyclone sequencing on landsliding and fluvial sediment transfer in a mountain catchment Guan-Wei Lin, 1 Hongey Chen, 1 Niels Hovius, 2 * Ming-Jame Horng, 3 Simon Dadson, 4 Patrick Meunier 2 and Max Lines 2 1 Department of Geosciences, National Taiwan University, No.1, Section 4, Roosevelt Road, Taipei, Taiwan, ROC 2 Department of Earth Sciences, University of Cambridge, Downing Street, Cambridge, CB2 3EQ, UK 3 Water Resource Agency, Ministry of Economic Affairs, Taiwan, ROC 4 Centre for Ecology and Hydrology, Maclean Building, Crowmarsh Gifford, Wallingford, OX10 8BB, UK *Correspondence to: Niels Hovius, Department of Earth Sciences, University of Cambridge, Downing Street, Cambridge, CB2 3EQ, UK. E-mail: nhovius@esc.cam.ac.uk Received 16 August 2007; Revised 19 November 2007; Accepted 19 November 2007 Abstract Patterns and rates of landsliding and fluvial sediment transfer in mountain catchments are determined by the strength and location of rain storms and earthquakes, and by the sequence in which they occur. To explore this notion, landslides caused by three tropical cyclones and a very large earthquake have been mapped in the Chenyoulan catchment in the Taiwan Central Range, where water and sediment discharges and rock strengths are well known. Prior to the M W 7 6 Chi-Chi earthquake in 1999, storm-driven landslide rates were modest. Landslides occurred primarily low within the landscape in shallow slopes, reworking older colluvial material. The Chi-Chi earthquake caused wide-spread landsliding in the steepest bedrock slopes high within the catchment due to topographic focusing of incoming seismic waves. After the earthquake landslide rates remained elevated, landslide patterns closely tracking the distribution of coseismic landslides. These patterns have not been strongly affected by rock strength. Sediment loads of the Chenyoulan River have been limited by supply from hillslopes. Prior to the Chi-Chi earthquake, the erosion budget was dominated by one exceptionally large flood, with anomalously high sediment concentrations, caused by typhoon Herb in 1996. Sediment concentrations were much higher than normal in intermediate size floods during the first 5 years after the earthquake, giving high sediment yields. In 2005, sediment concentrations had decreased to values prevalent before 1999. The hillslope response to the Chi-Chi earthquake has been much stronger than the five-fold increase of fluvial sediment loads and concentrations, but since the earthquake, hillslope sediment sources have become increasingly disconnected from the channel system, with 90 per cent of landslides not reaching into channels. Downslope advection of landslide debris associated with the Chi-Chi earthquake is driven by the impact of tropical cyclones, but occurs on a time-scale longer than this study. Copyright 2008 John Wiley & Sons, Ltd. Keywords: Landsliding; fluvial sediment transfer; tropical cyclone; earthquake; event sequencing Introduction Landsliding is the dominant erosion process in uplands where rates of rock uplift and fluvial bedrock incision are high. There, patterns of landsliding reflect substrate properties, hydrological conditions, existing topography and trigger mechanisms (Sitar et al., 1992; Chen and Su, 2001; Khazai and Sitar, 2003; Dadson et al., 2004). In addition, the distribution and intensity of landsliding associated with a given trigger may be controlled by preceding geomorphological, meteorological and seismic events. In Taiwan a rare opportunity exists to evaluate these controls in quantitative detail. In September 1999, the Chi-Chi earthquake with a moment magnitude M W = 7 6 caused widespread landsliding in central west Taiwan (Figure 1). This event was preceded by large tropical cyclones, such as typhoon Herb, in August

Landsliding and sediment transfer caused by typhoon and earthquake sequence 1355 Figure 1. Topography of Taiwan, showing epicentre and focal mechanism of the Chi-Chi earthquake. Dashed lines are peak vertical ground acceleration contours constructed using data from 441 strong ground motion recorders, 60 of which were within 20 km of the seismogenic Chelungpu fault (CPF) (Lee et al., 1999). Shaded region with solid outline indicates area of widespread landsliding (>2 per cent surface area disturbed) measured from SPOT and aerial-photograph landslide inventory. (Modified from Dadson et al., 2004.). This figure is available in colour online at www.interscience.wiley.com/journal/espl

1356 G-W. Lin et al. 1996, and was followed by further typhoons, including Toraji in July 2001 and Mindulle in June 2004. Here we report on landslide patterns and rates caused by this sequence of triggers, and the concomitant fluvial sediment transfer, in a mountain area drained by the Chenyoulan River, close to the epicentre of the Chi-Chi earthquake. Specifically, we have investigated the rate and location of landsliding as a function of topography, substrate properties and the nature of active and preceding triggers. We have also considered how hillslope mass wasting in the Chenyoulan catchment is reflected in the downstream transfer of sediment. Study Area The mountain island of Taiwan has formed from the rapid, oblique collision between the Luzon Arc on the Philippine Sea Plate, and the Eurasian continental margin. Its current mean annual precipitation is 2 5 m year 1, about 80 per cent of which falls between May and October, and the island receives an average of four typhoon hits per year (Shieh, 2000). The combination of strong climatic and tectonic forcing results in rapid rates of geomorphological processes, with average erosion rates of 3 7 mm year 1 (Li, 1975; Dadson et al., 2003). Tectonic forcing is perhaps best illustrated by the 1999 Chi-Chi earthquake. This event was the largest in Taiwan in 50 years and the largest on the Chelungpu thrust fault in 300 620 years (Shin and Teng, 2001; Chen et al., 2001). It occurred more than 80 years after the previous M w > 6 earthquake in central west Taiwan. The earthquake had a reverse fault mechanism, a focal depth of 8 km, and resulted in north south-trending surface rupture over a distance of about 100 km (Figure 1). It produced recorded ground accelerations of up to 1 g (Lin et al., 2000), and vertical surface displacements of up to 3m. The Chenyoulan River flows into the Choshui River 12 km southeast of the Chi-Chi epicentre. Rising at Yushan Mountain, the Chenyoulan River basin is dominated by kilometre-high bedrock ridges, and steep, narrow valleys (Figure 2). The area is sparsely inhabited, with a substantial native forest cover. Floored by bedrock, the Chenyoulan River flows northward, its channel opens up around Tungpu, and reaches a width of 300 m at Neimaopu gauging station (catchment size at the station: 367 km 2 ). This flow direction means that the principal topographic ridges in the catchment trend parallel to the seismogenic Chelungpu fault. The catchment itself is dissected by the east-dipping Chenyoulan thrust fault, with Paleogene slates and metasandstones of the Shueshan Range to the east, and Neogene interbedded shales and sandstones of the Nanchuang and Hoshe formations to the west. Alluvium is distributed along the trunk stream and has accumulated at the confluence of tributaries near Hoshe (Figure 2). The mechanical properties of these substrates are summarized in the next section. Rock Strength One-hundred and twenty-eight sets of rock samples were collected, covering the full range of lithologies within the Chenyoulan catchment. These samples were subjected to an unconfined compression test (Brown, 1981), to estimate the strength of each rock formation. The compressive strength of sandstone beds in the Nanchuang Formation ranged between 29 and 117 MPa (10 6 Nm 2 ), with an average of 69 9 MPa. In contrast, the compressive strength of shale in the Nanchuang Formation was less than 10 MPa. The ratio of sandstone to shale in the Nanchuang Formation is about 1:1, and a straight average of the strengths of both components is 42 MPa. If this is an appropriate measure of the bulk strength of Nanchuang rocks, then they can be classified as medium weak in the classification of the International Society for Rock Mechanics (ISRM) (Brown, 1981). Hoshe rocks were stronger. The compressive strength of sandstone units in this formation was 50 157 MPa, with an average of 96 5 MPa, and Hoshe shale beds had strengths of 37 49 MPa, with an average of 43 MPa. For a 1:1 ratio of sandstone to shale in the Hoshe Formation, the mean formation strength is 70 MPa, which is also regarded as medium weak rock. The compressive strength of metamorphic rocks east of the Chenyoulan River was 45 179 MPa, with an average of 101 8 Mpa and so these rocks are classified as being of medium strength. Metasandstones within the metamorphic series are strong and intercalated slates are much weaker, but their alteration is less regular than in the Nanchuang and Hoshe formations. This may result in a greater variability of rock mass strength within the metamorphic substrate. It should be noted that unconfined compression tests on small rock samples do not permit a complete assessment of rock mass strength on length scales relevant to hillslope failure (cf. Schmidt and Montgomery, 1995). Weighted averaging of local strengths of different components of a geological formation may account for some of the lithological

Landsliding and sediment transfer caused by typhoon and earthquake sequence 1357 Figure 2. Topography of the Chenyoulan catchment, central west Taiwan. (A) Shaded relief map of catchment with fluvial channel network thresholded at 1 2 km 2 upslope area, major geological formations and location of WRA station 1510H049 at Neimaopu. Area upstream of Neimaopu Station: 367 km 2, of which 42 per cent is in metamorphic rocks, 37 per cent is in Nanchuang Formation, 16 per cent is in Hoshe Formation, and 5 per cent is in alluvium. The Nanchuang and Hoshe formations consist of Neogene interbedded sandstones and shales. (B) Map of landslides triggered by Chi-Chi earthquake and subsequent typhoons. Landslides were mapped from time series SPOT images of site 299303, covering the interval 5 January 1999 to 7 July 2004. Also shown: fluvial channel network and topographic ridges mapped by hand from digital elevation model with nominal resolution of 40 m. variability within a formation or landslide site, but does not take into account the effects of weathering, jointing and saturation state. Attitude of beds may be especially important in layered rocks with different strengths. Laterally extensive, weak layers may dominate mechanical properties of the rock mass when the geological dip parallels the topographic slope. In contrast, when beds are horizontal or dip away from the topographic slope then strong layers may dominate rock mass strength. Geological dips in metamorphic rocks east of the Chenyoulan fault are variable. In the layered sediments west of the fault bedding is subhorizontal to gently eastward dipping. From our measurements it is clear that the catchment substrate cannot be treated as uniform. Instead, triggered mass wasting will be evaluated within geological formations as well as catchment-wide. Formation outcrop patterns in the catchment may complicate spatial analyses. In the catchment, rocks of the Hoshe Formation sit stratigraphically and structurally below Nanchuang rocks. Outcrops of the former are concentrated in low topographic positions, whereas outcrops of the latter are mainly in higher topographic positions. Metamorphic rocks crop out over the full height of the east flank of the catchment, and may offer the best opportunity of unbiased analysis.

1358 G-W. Lin et al. Major Triggers of Erosion in the Chenyoulan Catchment Amongst many episodes of erosion, four events have caused widespread landsliding and fluvial sediment transfer in central west Taiwan in the decade preceding 2006. These are: typhoon Herb in late July 1996, the Chi-Chi earthquake of 21 September 1999, and typhoons Toraji and Mindulle in July 2001 and June 2004, respectively. In this section we review these four events, and the landslides and flooding they caused within the Chenyoulan catchment, after a brief summary of relevant methods. We have mapped landslides in the upper Chenyoulan catchment from a sequence of Système Pour l Observation de la Terre (SPOT) satellite images of SPOT reference site 299303. These images have a pixel size of 12 5 m 12 5 m. Landslide locations were checked in the field and on aerial photographs (Lin et al., 2003). No distinction was made between scarps and deposits of landslides. Our maps are comprehensive and accurate for landslides with a surface area greater than about 600 m 2. Omission of smaller landslides has resulted in a minor underestimation of the area affected by slope failure. The SPOT images used in this study were collected on 5 June 1996, less than 2 months before typhoon Herb; 8 November 1996, two months after passage of Herb (additional images from 15 December 1997 and 12 October 1998 were used to map in areas with cloud cover); 5 January 1999, in the winter season preceding the Chi-Chi earthquake; 8 January 2000, four months after the earthquake (supplementary image from 5 March 2001); 12 August 2001, two weeks after typhoon Toraji (supplementary image from 17 December 2003); and 7 July 2004, less than 2 weeks after typhoon Mindulle. The principal images were selected to have an optimal combination of visibility and proximity to the dates of the four events of interest. Most of the mapped landslides can be attributed to these four events with confidence. This has been confirmed qualitatively by regular field inspections and anecdotal evidence. Previous work in similar topography and biomes in east Taiwan (Hovius et al., 2000) has demonstrated that landslide scars and deposits remain visible on remote sensed imagery for approximately 10 years before revegetation precludes systematic detection. Downstream of the area in which we have mapped landslides, the Water Resources Agency (WRA) of Taiwan has collected hydrometric data at Neimaopu station (catchment area 367 km 2 ) since 1972 (WRA, 1973 2001; http:// gweb.wra.gov.tw/wrweb). There, an automatic stage recorder has made continuous measurements of water stage. The water discharge at the station was calculated using stage discharge relations, obtained by surveying the channel crosssection and flow velocity every 6 months. A summary of hydrometric measurements at Neimaopu station is given in Figure 3. We have no measurements of rainfall in the catchment. Typhoon Herb was one of the largest, recent typhoons in Taiwan. The typhoon moved westward over north Taiwan with sustained wind speeds of up to 65 m s 1. Herb produced a total water discharge of 0 32 km 3 over 3 days at Neimaopu station, and flood discharge peaked at 1860 m 3 s 1. Upstream of Neimaopu station, Herb caused at least 1180 landslides with a total area of 4 59 km 2. Only 40 per cent of this area had not previously been affected by landslides in the decade before 1996 (Figure 4). In the 3 years following typhoon Herb typhoons bypassed the Chenyoulan catchment, and peak flows at Neimaopu did not exceed 380 m 3 s 1. Then, in September 1999, at the end of a relatively dry summer, the Chi-Chi earthquake triggered more than 20 000 landslides in central west Taiwan, mostly where peak ground accelerations were greater than 0 2 g (Lee et al., 1999; Dadson et al., 2004). The Chenyoulan catchment is located within this contour (Figure 1). Figure 3. Hydrometric measurements at Neimaopu Station (1510H049) for the decade starting 1 January 1996. Daily water discharge values are indicated in grey. Open circles represent measured suspended sediment concentrations obtained at a frequency of 30 ± 2 per year. Neimaopu Station was destroyed during typhoon Toraji in July 2001, and recommissioned in January 2003. Data are available on http://gweb.wra.gov.tw/wrweb.

Landsliding and sediment transfer caused by typhoon and earthquake sequence 1359 Figure 4. Summary of densities of mapped landslides in (A) the Chenyoulan catchment and (B E) its major geological formations for: approximately 10 years prior to typhoon Herb; typhoon Herb (1996); the Chi-Chi earthquake (1999); and typhoons Toraji (2001) and Mindule (2004). A distinction has been made between new landslides (dark grey) that occurred in locations not previously affected by landslides during the mapping interval, and landslides that reactivated locations previously affected by landslides during the mapping interval (light grey). Shown in black is the ratio of the formation-specific landslide density and the catchment-wide landslide density for a given trigger. It has been assumed that landslides older than approximately 10 years are not recognizable on SPOT images due to revegetation (cf. Hovius et al., 2000). In the catchment 2375 mapped landslides have been attributed to the earthquake. Together, these landslides affected an area of 12 85 km 2, of which 70 per cent had not been affected by landslides in the preceding interval covered by our SPOT images (Figure 4). No typhoons had a major impact on the Chenyoulan catchment in 2000, but in July 2001 typhoon Toraji caused significant mass wasting and flooding, which destroyed the hydrometric station at Neimaopu. Classified as a category 3 tropical cyclone on the Saffir Simpson scale, typhoon Toraji had maximum wind speeds of around 50 m s 1. It produced accumulated rainfall of 400 500 mm in the Chenyoulan catchment, more than elsewhere in the Chi-Chi epicentral area (Galewsky et al., 2006). We have no direct measure of the concomitant flood within our study area, but at a downstream station (1510H071) on the Choshui River peak discharge of the Toraji flood was twice that of the Herb flood whereas the total flood discharge of Herb was 1 5 times that of Toraji. Toraji caused at least 3730 landslides in the upper Chenyoulan catchment, with a total surface area of 24 43 km 2, of which 66 per cent had not been affected by earlier landsliding (Figure 4).

1360 G-W. Lin et al. Measurements were resumed at Neimaopu station in 2003. In the period without record, no major floods occurred downstream in the Choshui catchment, precluding very large discharges at Neimaopu. The next major flood (0 11 km 3 total discharge in 5 days) was in June 2004, during typhoon Mindulle, when discharges peaked at 620 m 3 s 1. Mindulle triggered at least 2051 landslides in the upper Chenyoulan catchment, with a total area of 31 58 km 2, of which 69 per cent had not been previously affected by landslides in the interval covered by our SPOT images (Figure 4). Further typhoon floods occurred in August 2004 and July and August 2005. Landsliding caused by these events has not been documented, but associated sediment transport will be considered briefly in a later section of this paper. Event-specific landslide densities varied significantly between lithologies (Figure 4), but the catchment-wide increase of landslide density during and following the Chi-Chi earthquake is found in all bedrock formations. The catchment-wide pattern is most faithfully reflected in the metamorphic terrain east of the Chenyoulan trunk stream, although slightly higher than average landslide densities were found there for all triggers. West of the river, densities of typhoon-triggered landslides were anomalously high in the Hoshe Formation, and anomalously low in the Nanchuang Formation, and the pattern was reversed after the earthquake. These observations rule out rock strength as a first-order control on the rate of landsliding in the Chenyoulan catchment. Landslide reactivation rates were very high throughout the catchment prior to the earthquake, and significantly lower during and after the earthquake. Landslide reactivation rate is defined here as the proportion of the area affected by landslides in the interval preceding a trigger that was affected by renewed landsliding caused by that trigger. These findings, summarized in Table I, highlight the effect of the Chi-Chi earthquake on mass wasting in the Chenyoulan catchment. Before the earthquake, the rate of landsliding, expressed here as the area affected by landsliding (km 2 year 1 ) per unit area (1 km 2 ), was comparatively low, 0 001 year 1, when taking into account reactivation of older Table I. Landslide statistics. Data used to create Figure 4. The Chenyoulan catchment has been divided according to its constituent geological formations. For each trigger event, the following quantities are tabled: Landslide density, expressed as a percentage of the total catchment or formation area; the percentage of the total landslide area made up of reactivated landslides; the landslide reactivation rate, expressed as the percentage of the total area of prior landslides; and the contribution of individual formations to the catchment landslide total, weighted by the percentage of the catchment area occupied by that formation, where a landslide contribution in exact proportion to the area occupied by the formation gives a proportional contribution value of 1 Landslide Formation density Reactivated landslides (percentage (percentage (percentage of total Reactivation Proportional of catchment area) Year Event of area) landslide area for event) rate (%) contribution Catchment 1995 Pre-Herb 0 66 (367 km 2 = 100%) 1996 Herb 1 02 59 74 93 0 1999 Chichi 2 85 29 75 82 7 2001 Toraji 5 35 34 58 65 0 2004 Mindulle 6 91 30 98 40 0 Nanchuang 36 90% 1995 Pre-Herb 0 37 0 56 1996 Herb 0 69 52 66 99 0 0 68 1999 Chichi 3 69 17 67 94 0 1 30 2001 Toraji 7 94 31 37 67 5 1 48 2004 Mindulle 11 56 27 35 39 8 1 67 Metamorphic 42 40% 1995 Pre-Herb 1 03 1 56 1996 Herb 1 56 61 89 94 0 1 52 1999 Chichi 3 45 37 11 82 0 1 21 2001 Toraji 6 69 39 64 77 0 1 25 2004 Mindulle 7 57 39 51 44 7 1 10 Hoshe 16 10% 1995 Pre-Herb 1 19 1 80 1996 Herb 1 64 63 29 87 7 1 61 1999 Chichi 3 12 39 27 74 5 1 10 2001 Toraji 3 73 22 00 26 3 0 70 2004 Mindulle 5 42 19.04 27 7 0 79 Alluvium 4 60% 1995 Pre-Herb 0 83 1 26 1996 Herb 1 78 40 13 86 0 1 74 1999 Chichi 4 03 34 41 78 0 1 42 2001 Toraji 6 64 30 36 50 0 1 24 2004 Mindulle 5 15 40 81 31 7 0 75

Landsliding and sediment transfer caused by typhoon and earthquake sequence 1361 landslides. Landslides were relatively small, with an average size of A ls = 3 2 10 3 m 2 of mapped landslides. During typhoon Herb, the reactivation rate of existing landslides was extremely high, 93 per cent, and formations with outcrops low in the catchment and close to trunk streams were relatively prone to failure. The Chi-Chi earthquake caused widespread landsliding, and the average size of mapped coseismic landslides was larger, A ls = 5 4 10 3 m 2, but the true impact of the earthquake is reflected in post-seismic landslide statistics. In the 5 years following and including the earthquake, the catchment-wide landslide rate was 0 0133 year 1, 13 times higher than the pre-chi-chi landslide rate. Postseismic landslides were larger than preseismic ones, A ls = 6 5 10 3 m 2, and A ls = 15 4 10 3 m 2, for Toraji and Mindulle, respectively, and reactivation rates were lower, 65 per cent and 40 per cent, respectively. Here it should be noted that landslide size frequency relations often take the form of a power law or Pareto distribution (Hovius et al., 1997; Stark and Hovius, 2001; Malamud et al., 2004), in which the mean landslide size can be strongly influenced by the largest event. In the case of this study, the largest landslides were an order of magnitude smaller than the total area affected by landslides with a shared trigger. Therefore we believe that the mean values presented here are a robust, if incomplete measure of landslide size. A striking comparison is between the impact of typhoons Herb and Mindulle. The former was three times larger in terms of total runoff and peak discharge, but landslides triggered by the latter affected an area that was six times larger. It is clear that landslide rates have remained high in the years since the Chi-Chi earthquake, and that for a given meteorological trigger, propensity to failure of hillslopes in the Chenyoulan catchment has been increased due to the earthquake. Next, we shift attention to the effect of the Chi-Chi earthquake on the location of landslides. Landslide Location Landslide location can be characterized across a range of spatial scales. On a large (10 1 10 2 km) scale, landsliding can be evaluated for example as a function of distance from earthquake epicentre (e.g., Keefer, 1994; Dadson et al., 2004; Meunier et al., 2007), or rainfall pattern. The Chenyoulan catchment is not suitable for work on this scale. Instead, we have documented landslide location at the hillslope scale, asking three questions. 1. How steep were slope segments affected by landslides? 2. Where did landslides occur with respect to the upper and lower limits of slopes? 3. How efficiently did they connect with the channel network? To answer these questions, landslide maps were combined with a digital elevation model (DEM) of the catchment with a node spacing of 40 m. This DEM was constructed by photogrammetric means from aerial photographs taken before 1996. To address the first question, local topographic slopes were measured as the steepest slope within a square of 3 3 DEM grid cells, throughout the Chenyoulan catchment. The frequency plot of all measured topographic slopes in the catchment (Figure 5a) shows a strong modal peak at 37, a rapid decrease of abundance of steeper slopes, and a secondary peak at <10. This distribution is common in mountain landscapes. The high modal peak is thought to reflect the maximum stable gradient of bedrock mountain slopes (Schmidt and Montgomery, 1995; Burbank et al., 1996), set by the bulk strength of the rock mass. The second, low modal peak reflects the presence of active or terraced alluvial flats at the base of larger bedrock valleys, possibly with addition of artificial low slope values returned for squares including opposing sides of a narrower bedrock channel or ridge. Differences exist between the slope frequency plots for individual formations within the catchment (Figure 5b e). Modal slopes are 38 in metamorphic rocks, 34 in the Hoshe Formation, 33 in the Nanchuang Formation and 6 in alluvium. This is in good qualitative agreement with the rock strengths presented in section 3, and implies that stability of slopes in sedimentary rocks of the Hoshe and Nashuang formations is determined by the strength of the competent sandstone beds, rather than the much weaker shale beds. Steeper than modal slopes are relatively common in metamorphic rocks. A secondary, lower modal peak at 12, was found within terrain attributed to the Hoshe Formation. In alluvium, slopes of 15 to 30 occur with similar, intermediate frequencies. Also shown in Figure 5 are slope frequency plots for locations affected by landslides. On the catchment scale (Figure 5a), landslide slope frequency distributions have identical modal values of 40 for all three typhoons as well as the Chi-Chi earthquake. The landslide populations of the two post-chi-chi typhoons have identical slope frequency distributions, tight, symmetric and with very little activity in slopes <25 and >65. The modal peak of slopes affected by coseismic landslides is broader, extending into higher gradients, and the relative frequency of landslides at sites with gradients <25 was significantly higher than during subsequent typhoons. During typhoon Herb, relatively many landslides affected sites with low topographic slopes. High frequencies of steeper than modal

1362 G-W. Lin et al. Figure 5. Probability distribution of all topographic slopes (solid lines) and slopes affected by landsliding (dashed/dotted lines) in (A) the Chenyoulan catchment and (B E) its major geological formations. Local topographic slopes have been calculated as the steepest slope within a square of 3 3 DEM grid cells. Statistics for landslides triggered by typhoon Herb in metamorphic rocks are shown with (grey) and without (black) exceptionally large landslide as discussed in main text.

Landsliding and sediment transfer caused by typhoon and earthquake sequence 1363 slopes for this trigger result primarily from one very large landslide (0 42 km 2 ), high up in steep metamorphic terrain. We suspect that the return time of this landslide is significantly greater than the 10 years covered by our landslide map. Removal of this landslide from the sample results in strong reduction of high slope frequencies, and a downward shift of the modal value to 38. In all three bedrock formations, slope frequency distributions of landslides triggered by typhoons Toraji and Mindulle are similar, with identical modal peaks at 40, irrespective of the slope frequency distribution of the topography in these formations. In metamorphic rocks, coseismic and post-seismic landslides share a similar slope frequency distribution. In the Nanchuang Formation, the coseismic distribution has a modal peak at 47 and is skewed to higher slopes. A similar, very high modal slope value was found for coseismic landslides in the Hoshe Formation, in addition to a second, stronger modal peak at 35. Landslides triggered by typhoon Herb affected sites with significantly lower topographic slopes: modal slopes of landslides are 25 and 34 in the Hoshe and Nanchuang formations, respectively. The very high modal slope value for landslides triggered by Herb in metamorphic rocks is due to the single large failure mentioned earlier. The equivalent modal slope without this landslide is 38. We refrain from a detailed analysis of landslides within alluvium. Insight into the effect of landslide location on topography can be gained by comparing the slope frequency distributions of topography and the landslides within that topography. Where the probability of finding a landslide in a unit area with a given gradient is greater than the probability of finding a unit area with that gradient in the domain of the study, we say that landslides have oversampled the topography, and vice versa for undersampling. In general, co- and postseismic landslides in the Chenyoulan catchment have undersampled locations with low topographic slopes, and oversampled steep terrain. Thus, these landslides have acted to reduce the steepness of the landscape. The switch from undersampling to oversampling is within a few degrees of the modal slope of the topography, lending strength to the hypothesis that this is the maximum stable slope of the landscape. Landslides prior to the Chi-Chi earthquake show a notably different pattern. On the whole, these landslides occurred on shallower slopes than co- and post-seismic landslides, and in the Hoshe Formation, slopes between 15 and 30 were substantially oversampled. These landslides have had a very limited effect on the steepest sites in the catchment. Instead, they have shifted mass in terrain where bedrock can be assumed to have been stable. The second question concerns the position of landslides with respect to ridges and streams. To address this question, ridges were mapped by hand using elevation, slope aspect and curvature of topography obtained from the catchment DEM. Fluvial channels were mapped, using a flow-routing algorithm, thresholded at an upslope area of 1 2 km 2 to reflect the break in slope area scaling found at this scale (Figure 6) (e.g., Montgomery, 2001). For each point, we have calculated the shortest flow path connecting ridge and fluvial channel to represent the hillslope on which the point is located. In order to enable direct comparison between points, the position of any point within the topography can then be expressed in terms of its distance from the limits of the hillslope, normalized by the total length of the slope. Normalized positions vary from 0 for a point in the river channel to 1 for a point on the ridge crest. However, through this normalization an absolute measure of slope and landslide size has been lost. Three types of errors are inherent in this method. Figure 6. Plot of local slope, S, against upslope area, A, for the Chenyoulan catchment. Slope is the mean slope measured in the direction of steepest descent within logarithmically spaced bins of upslope area. The break in scaling at approximately 1 2 km 2 upslope area is interpreted as the transition from hillslopes (A < 1 2 km 2 ) to fluvial channels (A > 1 2 km 2 ). Values of the scaling exponent on power law best-fits to hillslope and channel data are 0 07 and 0 84 respectively.

1364 G-W. Lin et al. I. The resolution of SPOT images limits the precision of landslide maps to 12 5 m. II. The resolution of the DEM limits the precision of ridge and stream location to 40 m, and in our analysis landslides are snapped to the 40 m DEM grid. III. Manually determined ridge tips are subject to a potential error of several DEM pixels. Type I errors are subsumed in type II landslide location errors. Therefore, the maximum quantifiable error is 120 m. In the slope length normalization process this is the error on the denominator. The maximum error on the denominator is 80 m, for ridge and stream location. This gives a maximum quantifiable error of 200 m on a normalized distance plot. The average length of slopes with landslides in the Chenyoulan catchment is 1275 m. The maximum quantifiable error on the normalized distance is 16 per cent of this length. In almost all cases the real error will be considerably smaller than this percentage. In rare cases where landslides occurred directly below the tip of a ridge an undetermined additional error applies. In Figure 7, we have shown the location of landslides as a function of normalized distance to stream, where ridges are a unit distance away from streams. Because the length of the ridge network is less than the length of the stream network, points with a small normalized distance to stream are more common than points closer to ridges. To account for this, we have calculated a probability ratio R p = P ls /P topo, where P ls is the probability of a point within a landslide being at a given normalized distance to the stream network, and P topo is the probability of any point in the catchment topography being at that normalized distance to the stream network. A value of R p = 1 indicates that landslides have affected topography at a given distance to the stream network in exact proportion to the available topography. A value for R p > 1 indicates oversampling by landslides, etc. We are not aware of a precedent for this method. The catchment-wide distance probability plot (Figure 7a) shows two slope domains. In the lower domain, with a normalized distance to stream smaller than about 0 35, probability ratios decrease steadily with increasing distance to stream for landslides triggered by typhoons Herb and Toraji and the Chi-Chi earthquake, and remain approximately constant for landslides triggered by typhoon Mindulle. Beyond a normalized distance of 0 35, in the upper slope domain, landslide probabilities increase for all four triggers, and then decrease again in the uppermost quartile of the topography. This pattern is well developed in the subset of landslides in metamorphic rocks, but less so in the sedimentary rocks of the Hoshe and Nanchuang formations. In the Hoshe Formation, propensity to failure decreased steadily with increasing distance to stream for all triggers, except typhoon Mindulle when landslide probability was distributed relatively uniformly. In the Nanchuang Formation, uniform landslide distributions were found for the Chi-Chi earthquake, and typhoon Mindulle. The other two typhoons caused progressively less mass wasting at greater distance to the stream network. During typhoon Herb, landslides oversampled lower slope segments throughout the catchment, most so in the Hoshe Formation, and to a lesser degree in metamorphic substrates. The upper slope segments were undersampled during Herb, except for the metamorphic domain, where a single landslide distorts the picture. Coseismic landslides also oversampled the lowest slope segments throughout the catchment. Coseismic oversampling further occurred selectively around the broad probability peak two-thirds up within the catchment slopes, especially within the Nanchuang Formation. The coseismic pattern of distance probability is closely tracked by landslides triggered by typhoon Toraji, whereas typhoon Mindulle activated all locations in sedimentary rocks in approximately equal and neutral measures, but oversampled upper slopes in metamorphic substrate. We decline to offer an analysis of the landslide location patterns within alluvium but include the distance probability plot for completeness. Discussion Landsliding Findings on slope stability and failure in the Chenyoulan catchment can be summarized as follows. 1. Prior to the Chi-Chi earthquake, storm-driven landslide rates were modest, landslides were relatively small, and occurred primarily in low positions within the landscape, often with relatively low topographic gradients, mostly reworking older landslide deposits or other colluvial material. 2. The Chi-Chi earthquake caused widespread landsliding, affecting primarily the steepest bedrock slopes within the catchment. Slope segments located within the upper half of the topography were disproportionately prone to coseismic failure, as were locations close to streams. 3. After the earthquake, propensity to slope failure remained elevated throughout the catchment, and patterns of

Landsliding and sediment transfer caused by typhoon and earthquake sequence 1365 Figure 7. Plot of landslide probability against normalized distance from stream for (A) the Chenyoulan catchment and (B E) its major geological formations. For each location affected by a landslide, normalized distance to stream has been measured along a straight line through this location connecting the nearest point with an upslope area A > 1 2 km 2 (our technical definition of a stream) with the nearest point on a mapped ridge, and then normalizing for the length of this line. The probability ratio is the ratio of the probability of being in a landslide at a given normalized distance from stream and the probability of being somewhere in the catchment topography at that distance from stream. The thin solid line indicates equal probability.

1366 G-W. Lin et al. storm-triggered landsliding have closely tracked the distribution of coseismic landslides. Co- and post-seismic landslides were larger on average than pre-seismic landslides. To explain these observations, we consider the ways in which earthquakes affect the stability of slopes. First, seismic shaking may cause loss of cohesion and/or reduction of the frictional strength of the substrate due to shattering of rock mass. Widespread rock mass shattering was observed in many steeper sites within the Chenyoulan catchment, and has also been reported from the epicentral areas of other, large earthquakes (e.g., Harp and Jibson, 1996). Second, addition of seismic acceleration to gravitational acceleration results in short-lived and episodic changes of the normal and shear stresses in hillslopes during earthquakes. Net surface accelerations may be directed upward and/or outward from a hillslope. Peak vertical accelerations of >1 g have been recorded during several earthquakes, not necessarily very large ones (e.g., Geli et al., 1988; Bouchon and Barker, 1996). Thus, earthquakes may be expected to cause failure of the weakest slopes in a landscape. Conforming to this expectation, the steepest bedrock slopes in the Chenyoulan catchment were found to have been especially prone to failure. Similarly, high landslide rates low on Chenyoulan hillslopes may be associated with the low mechanical strength of colluvial material concentrated in these locations, and/or the presence of steep inner gorges (e.g., Kelsey, 1988; Densmore and Hovius, 2000). However, these simple considerations do not explain the observed peak of coand post-seismic failure rates in upper hillslope segments. This may be due to topographic site effects on incoming seismic waves. Unexpectedly large seismic accelerations are often observed at hill tops and ridge crests (e.g., Oldham, 1899). Topographic amplification of seismic energy occurs when seismic waves entering the base of a mountain ridge are partially reflected back into the rock mass upon contact with the Earth s surface. Thus, seismic waves are progressively focused upwards and the constructive interference of their reflections and the associated diffractions increases towards the ridge crest. This often gives rise to a marked amplification of seismic accelerations on topographic highs (Geli et al., 1988; Benites and Haines, 1994; Bouchon et al., 1996). Accordingly, it is expected that propensity to slope failure during an earthquake increases towards mountain ridges, and in convex-up slopes (cf. Harp et al., 1981; Keefer, 2000). Our observations are quantitative confirmation of this expectation. Focusing of landslides in higher slope segments is clear in catchment-wide statistics, and in statistics for landslides in metamorphic rocks that make up the entire east flank of the catchment. It is not immediately evident within the Hoshe and Nanchuang formations, but co- and postseismic landslide rates were lower than the catchment-wide norm in most of the Hoshe terrain, and higher than normal throughout the Nanchuang Formation (Table I). Hoshe and Nanchuang rocks make up the lower and upper west flank of the Chenyoulan catchment, respectively. We tentatively attribute the difference in landsliding between them to effective focusing of landslide-triggering seismic waves in the largest ridges of the catchment only, having eliminated modest differences in rock strength between these formations as a primary control on landslide intensity (see above). Ridges with a subkilometre base width, such as those within the Hoshe Formation, appear not to have suffered significant topographic site effects. This places some useful constraints on the maximum frequency of seismic waves responsible for slope failure in the Chenyoulan catchment, but a full investigation of this aspect is beyond the scope of this paper. Post-seismic typhoons have triggered more and larger landslides in the Chenyoulan catchment, and they have affected a different, higher part of the landscape, when compared with typhoon Herb. Although it is not possible to fully quantify the magnitude of this effect, it is clear that the earthquake has increased the propensity to failure of steep slopes. The majority of post-seismic landslides are new, not seeded within older landslide debris. This implies that the weakening of the substrate due to seismic shaking is a long-lasting effect of the earthquake. Rock mass cracking has been observed in many steep slopes within the catchment, especially around ridges and convex-up breaks in topographic slope, and coalescence of cracks has been inferred from increased post-seismic groundwater flow rates elsewhere in the epicentral area (Wang et al., 2004). This is a likely cause of lasting strength reduction in near-surface rocks. In 2004, five years after the Chi-Chi earthquake, it remained a first-order control on the pattern and rate of mass wasting in the Chenyoulan catchment. There is an indication that the pattern of post-seismic landsliding has evolved in one significant way. Landslides triggered by typhoon Toraji in 2001 have oversampled the lowest hillslope segments, as did pre- and coseismic landslides. Of the total area affected by Toraji landslides, 31 per cent was connected with the channel network. During typhoon Mindulle, 3 years later, landslide rates were normal, or less than normal in the lowest slope segments, but remained elevated in higher sections of the landscape, and only 21 per cent of the total area affected by Mindulle landslides was connected with the channel network. Although it is possible that Mindulle rainfall was focused on higher ridges, we think it more likely that progressive removal of colluvial material from lower hillslopes was the cause of this change. A shift from shallow landsliding in unconsolidated material to bedrock landsliding may also be a cause of the observed increase of the average landslide size over our observation interval. In any case, it is likely

Landsliding and sediment transfer caused by typhoon and earthquake sequence 1367 that the upslope shift of landsliding within the Chenyoulan catchment has resulted in a decrease of the proportion of landslide sediment delivered to rivers. Fluvial sediment transfer Rates of sediment production and hillslope mass wasting in the Chenyoulan catchment have increased due to the Chi- Chi earthquake. This is true throughout the epicentral area of the earthquake, and it has caused an increase of the sediment transport of all rivers in the region. Dadson et al. (2004) have reported that the average suspended sediment load of the Choshui River, to which the Chenyoulan River contributes its discharge, increased by a factor 3 8 following the 1999 Chi-Chi earthquake, and that this increase was larger than average in the Chenyoulan River. Here we look in more detail at suspended sediment transport in the Chenyoulan River, using hydrometric data from Neimaopu station. To measure the suspended sediment load of the Chenyoulan River, the Water Resources Agency (WRA) has collected river water samples by using a DH-48 depth-integrating suspended sediment sampler, on average, 30 ± 2 times per year, at Neimaopu station. Each sample is filtered, dried and weighed, and the concentration of suspended sediment is recorded (Figure 3). Combined with the continuous record of water discharge at the station, sediment concentration measurements yield estimates of daily sediment discharge for sampling days. In turn, these estimates can be used to obtain a total annual suspended sediment discharge using the monthly weighted average method (MWA) (Cohn, 1995): E MWA = 365 12 12 i where E MWA is the calculated amount (tonnes) of sediment discharge (t year 1 ), m i is the measured number in the ith month, Q sij is the jth measured value of sediment discharge in the ith month (t day 1 ). This method has been found to be a good way to estimate sediment transfer in supply-limited rivers from limited sediment concentration data (Fuller et al., 2003; Dadson et al., 2003). Figure 8 gives an overview of annual suspended sediment discharge at Neimaopu station since 1972. In the period 1972 1998, the average annual sediment discharge was 2 7 Mt year 1 ± 1 9 Mt year 1 (standard error), equivalent to a catchment-wide average annual erosion rate of 2 8 mm year 1 (assuming a substrate density of 2 65 t m 3 ). This figure does not include bedload transport. After the Chi-Chi earthquake, the average annual suspended sediment discharge at Neimaopu station was 15 0 Mt year 1 (standard deviation 18 9 Mt year 1 ), taking into account the 2001 2002 interval without data. This is a fivefold increase with respect to pre-earthquake conditions. 1 mi m i j= 1 Q sij Figure 8. Time series of total annual suspended sediment transfer at Neimaopu Station from 1972 to 2005. Total annual values were calculated with the Monthly Weighted Average method of Cohn (1995). No data are available for 2001 2002. Interannual average total annual values for intervals before and after the Chi-Chi earthquake are indicated with dashed lines, dotted lines are standard errors on these averages.

1368 G-W. Lin et al. The monthly weighted average method is not suitable for the estimation of sediment discharge during individual floods, and the effects of this have propagated into our results. Compare, for example, the estimated total annual sediment transfer for 1990 and 1996. The MWA method returns a value of 15 4 Mt year 1 for 1990 and 2 6 Mt year 1 for 1996, using all available sediment concentration data. In 1990, typhoon Yancy was responsible for most sediment transport at Neimaopu station. Peak flood discharge (19 August 1990) was 729 m 3 s 1, and suspended sediment concentration was measured during this peak at 21 000 ppm. This high value was carried into the average sediment load estimate (0 35 Mt day 1 ) for August 1990, giving an artificially high estimate of the sediment discharge during that month of 11 Mt. In contrast, during the flood associated with typhoon Herb in 1996, sediment was first sampled 6 days after water discharge peaked. Maximum daily flood discharge was 1860 m 3 s 1, but at the time of sampling it was only 107 m 3 s 1. Although the suspended sediment concentration was found to be 26 300 ppm, this did not translate into a high sediment discharge value for the day, and this in turn propagated into monthly and annual totals, which remained artificially low. It follows that estimates of sediment transfer should take into account the specifics of individual floods (cf., Milliman and Kao, 2005; Milliman et al., 2007). To do this, a rating curve approach was adopted. Rating curve estimation methods are useful for interpolating between intermittent measurements, provided that a strong relation exists between water discharge and suspended sediment concentration. Figure 9 shows the water discharge for all published sediment concentration measurements at Neimaopu station (http://gweb.wra.gov.tw/wrweb). Data range over three decades of water discharge magnitude, and five decades of sediment concentration magnitude. Prior to the Chi-Chi earthquake, suspended sediment concentrations varied by up to three decades of magnitude for a given water discharge, indicating a strong supply control on fluvial sediment transport (Hovius et al., 2000), but a clear positive correlation existed between suspended sediment concentration and water discharge. The data can be described by a power law of the general form C s = κq α, where C s is the suspended sediment concentration at discharge Q, κ is the expected sediment concentration at unit water discharge (1 m 3 s 1 ), and α is a scaling exponent. Prior to the Chi-Chi earthquake, κ = 137 ± 22 7 ppm, and α = 0 743 ± 0 03 for the best fit power law, with R 2 = 0 311. This fit was obtained with a straight least-squares loss function. Loss functions of this type produce the best overall fit to the data, but by their nature are heavily influenced by the higher values in a data set. In the case of our hydrometric data, the fit procedure deals well with flood events, but less so with more common low and intermediate discharges and sediment concentrations. Given that our focus here is on floods, we have opted for use of a least-squares loss function, rather than another method that would produce a more robust fit to smaller data values. After the Chi-Chi earthquake, suspended sediment concentrations were generally at the high end of the value range, and variability of concentrations for a given water discharge was reduced. For the best-fit power law to postearthquake measurements, κ = 304 ± 117 ppm, and α = 0 827 ± 0 07, with R 2 = 0 553. Fixing the scaling exponent α at the pre-quake value does not significantly reduce the quality of the fit, R 2 = 0 545, but forces an upward shift of the specific sediment concentration to κ = 473 ± 33 ppm. This facilitates a direct comparison of sediment concentration measurements over the entire gauging interval. Our findings indicate that the Chi-Chi earthquake has caused an increase of the quantity of sediment available for uptake by the Chenyoulan River. Across the range of discharges, the suspended sediment concentration was four times higher, on average, after the earthquake. This has been the principal cause of increased post-seismic sediment export from the catchment. During individual typhoon floods, measured sediment concentrations have evolved according to a common pattern (Figure 10). Event measurements are well described by a power law relating water discharge and suspended sediment concentration, with scaling exponents similar to the overall value, α = 0 743. Again, event rating curves tend to overestimate sediment concentrations at low discharges, due to use of the least-squares loss function, but the method gives close fits to high sediment concentrations during flood peaks, when most sediment is transported (Figure 10). Fixing α = 0 743 (as in all later examples), the power law best fitting hydrometric measurements for typhoon Herb has a specific sediment concentration κ = 678 ± 75 ppm (R 2 = 0 88). Using this event rating curve, we estimate that during the first day of the flood 29 ± 3 Mt of suspended sediment passed through Neimaopu station at a concentration of 182 000 ppm, and that the total suspended sediment load of the flood was about 44 ± 5 Mt (Figure 10). This total for the 10 day flood is equivalent to 51 64 per cent of the MWA estimate of all suspended sediment transfer through Neimaopu station between 1972 and 1998. Our estimates of sediment transport during the Herb flood are in good general agreement with values published for stations further downstream (Milliman and Kao, 2005). For comparison, and to underscore our note of caution about results obtained with the MWA method, the best rating curve for three available data points for typhoon Yancy in 1990 gives an estimate of 3 4 ± 0 5 Mt suspended sediment transported during that event. Typhoon Herb had an exceptional impact on the hydropgraph at Neimaopu station. The average discharge of 1860 m 3 s 1 during the first day of the flood (equivalent runoff: 0 44 m day 1 ) was more than 1000 m 3 s 1 greater than