Catastrophic soil erosion in Iceland: Impact of long-term climate change, compounded natural disturbances and human driven land-use changes

Catena 98 (2012) 41–54

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Catastrophic soil erosion in Iceland: Impact of long-term climate change, compounded natural disturbances and human driven land-use changes Sigurdur Greipsson Biology and Physics Department, Kennesaw State University, Kennesaw, GA, 30144, USA

a r t i c l e

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Article history: Received 15 October 2011 Received in revised form 25 April 2012 Accepted 30 May 2012 Keywords: Catastrophic soil erosion Ecosystem degradation Heathlands Katabatic wind Sand encroachment Wind erosion

a b s t r a c t This study examines the interplay between long-term climate changes, compounded natural disturbances and human driven land-use changes on catastrophic soil erosion of the heathland ecosystem of Haukadalsheiði, south Iceland. Soil erosion was catastrophic for three centuries (~1660–1960 AD) and was characterized by almost total loss of vegetation and underlying soil. Soil erosion resulted in a desertified, barren landscape that had no resemblance to the original heathland ecosystem. Soil erosion was spatially reconstructed in a chronological order using information on the average progress of eroding fronts, anecdotal and historical evidence along with tephrochronological information. The progress of the fastest eroding front was rapid (29.7 m yr− 1). Human driven land‐use changes played a role in the heathland degradation: relentless free-range grazing by livestock resulted in decreased resistance of heathland communities to soil erosion. Adverse climate-change during Little Ice Age (LIA: 1550–1850 AD) intensified the effect of grazing. The catastrophic soil erosion was triggered by a massive sand encroachment ~1660 AD from three outwash sand-plains along the glacial River Far. The sand drift was sustained by dry northern glacial (katabatic) winds that drove the soil erosion. Longterm climate change resulted in glacier fluctuation that caused changes in water discharge in the River Far; sand drift was intense during periods of no water discharge (~1660–1708 AD and ~1800–1929 AD) and following glacial river floods (1708, 1884, 1902, 1929 and 1939 AD).Also, sand drift was intense due to unusually frequent volcanic tephra fallouts (1693, 1721 and 1766 AD). Information on factors that increase the risk of soil erosion and trigger and drive soil erosion is critical in understanding catastrophic soil erosion. © 2012 Elsevier B.V. All rights reserved.

1. Introduction Widespread soil erosion has occurred in Iceland and resulted in a complete loss of ~50% of the vegetation and soil that was present at the time of human settlement (874 AD) (Ólafsdóttir et al., 2001). Soil erosion has been particularly devastating in the interior (>300 m elevation) around the major glaciers and along glacial rivers (Arnalds, 1987). The intensity of soil erosion that has occurred in Iceland is unprecedented in northern Europe. Inappropriate land management including clearing of birch woodlands (Betula pubescens var. pumila) and overgrazing by livestock, has been emphasized as the principal cause of catastrophic soil erosion (Arnalds, 1987; Dugmore and Buckland, 1991; Dugmore et al., 2000, 2009; Gathorne-Hardy et al., 2009; Gísladóttir et al., 2010; Simpson et al., 2001). Prior to human settlement dense birch woodlands may have covered ~20% of the island (Jónsson, 2004). Paleoenvironmental reconstructions have shown that woodland species disappeared early after settlement around farms in south Iceland (Erlendsson and Edwards, 2009; Erlendsson et al., 2009; Gathorne-Hardy et al., 2009). Today, ~ 94% of woodland cover has disappeared (Jónsson,

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2004). However, most woodlands were located on the lowland b300 m elevation, but not on the highland plateau where most drastic soil erosion occurred. It has been suggested that overgrazing of livestock breached the cover of vegetation and exposed soils and consequently initiated soil erosion (Gísladóttir et al., 2010). Land management models such as the “tragedy of the commons” (Hardin, 1968) have previously been used to examine probable causes of land degradation in Iceland. This model was not found to adequately explain the drastic soil erosion that occurred in a mountainous range in south Iceland (Simpson et al., 2001) or large scale soil erosion that has occurred in north Iceland (McGovern et al., 2007). Another suggestion is that humans have crossed certain landscape or ecological thresholds and therefore initiated widespread soil erosion (Simpson et al., 2001). Less emphasis has, however, been put on other contributory factors, such as the role of long-term adverse climate change, natural disturbances (including katabatic winds, frequent fallout of volcanic tephra and glacial river floods) in triggering and driving widespread soil erosion. Widespread soil erosion was initiated soon after Settlement and was very intensive during the Little Ice Age (LIA) (Geirsdóttir et al., 2009; Gísladóttir et al., 2010; Ólafsdóttir and Guðmundsson, 2002). However, Gathorne-Hardy et al. (2009) argued that soil erosion was initiated before the LIA. Recent studies have indicated that soil erosion in Iceland precedes Settlement

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Table 1 Information used to reconstruct the advance of soil erosion. Historical data:

Written accounts on farm surveys (Magnússon and Vídalín, 1713) Written documents on farm descriptions (Pálsson, 1884) Map (Danish General Staff, 1908) Aerial photographs (1960, 2008) Locations of ruins (shiels, sheep-shelters) Anecdotal evidence: Interviews with farmers 1972–2010 Tephrochronology: Soil profiles on non-eroded sites surrounding the study area (Jones et al., 2007; Sigbjarnarson, 1969) Geological/geomorphological Location of soil remnants: Andisols and Histosols data (field work): Vegetation remnants Dried river courses

(Geirsdóttir et al., 2009; Haraldsson and Ólafsdóttir, 2003; Ólafsdóttir and Guðmundsson, 2002). This study focuses on the advance and mechanism of catastrophic soil erosion of a well defined area (Haukadalsheiði) on the highland plateau of south Iceland. Soil erosion in the study area has been regarded as being one of the most intensified in Iceland over the 19 and 20th Centuries (Sigbjarnarson, 1969). This paper aims at spatially reconstructing in a chronological order the advance of soil erosion by using empirical, historical and anecdotal evidences along with tephrochronological information (Table 1). 2. Study site Haukadalsheiði (64°20′–64°30′ N, 20°30–20°00′ W), is located in south Iceland (103,000 km 2) ~ 70 km from the south coast and close to Langjökull Glacier (925 km 2) (Fig. 1). The River Far flows from Lake Hagavatn (~5 km 2) to Lake Sandvatn (~10 km 2) and has one tributary: Jarlhettukvísl (Fig. 1). The River Sandá flows east from Lake Sandvatn. The River Ásbrandsá used to flow south of Lake Sandá but was dammed in 1996 AD (Fig. 1). The water level of Lake Sandvatn is at an elevation of 274 m but it is seasonal with the lowest level in early spring (April) and highest level in the fall (October). Most of Lake Sandvatn is covered with ice during winter. Catastrophic soil erosion resulted in almost complete loss of the original heathland vegetation and underlying soil. Today, only ~ 20% of the original vegetation and soil still exist as fragments in Haukadalsheiði. Catastrophic soil erosion shifted the ecosystem to a desertified, barren landscape comprised of gravel flats (glacier moraines), sands, bedrock and lava. Haukadalsheiði was protected from livestock grazing in 1964 and ecological restoration was implemented to halt further erosion and initiate revegetation of barren lands (Greipsson, 2012; Greipsson and El-Mayas, 1999). The study area lies between ~200–360 m a.s.l. elevation. Therefore, it is located between the subarctic and arctic climatic regions (Thórhallsdóttir, 1997). The growing season (>4 °C) is usually from early May to late September. The mean summer (June–August) temperature in 1936–1985 AD was ~ 9 °C (Einarsson, 1991) and annual precipitation exceeded 1200 mm (Flowers et al., 2007). Most precipitation falls as snow during winter. Dry spells are most common during spring. The prevailing wind direction is northeastern (NE) (Fig. 2). Dry northern storms (20–30 m s − 1) are common during the summer and are associated with wind-borne soil erosion. At the time of Settlement probably most of Haukadalsheiði was covered with dwarf-birch-heathland and birch-woodlands. There are still few, very small remnant stands of birch-woodlands in the study area and remnants of dwarf-birch-heathland exist. Ancient local names

of places indicate that such woodlands have previously been more extensive. At lower elevations (~100 m. a.s.l.) Haukadalur farm was surrounded through the centuries by dense birch woodland. The landscape is heterogeneous and harbors several plant communities. The most common communities are barren lands (sand, gravel and lava), dwarf-birch-heathland, grasslands and wetlands. Barren lands have typically low vegetation cover (b5%) of grasses, such as Festuca richardsoonii, Festuca vivipara and Agrostis sp. and herbs such as Rumex acetocella, Silena acaulis, Koenigia islandica, Arenaria norvegica and Spergula arvensis. Recolonization of native plants on barren lands is a very slow and erratic process (Greipsson and El-Mayas, 1999; Marteinsdóttir et al., 2010). Sands in the central part of the study area became colonized by the dune building grass Leymus arenarius. The heathland community is dominated by few dwarf-shrub species. These include Betula nana, Salix lanata and Vaccinium uliginosum. Grasses (Festuca richardsonii, F. vivipara and Agrostis vinealis) and sedges (mainly Carex bigelowii) form dense cover. The grassland community is dominated by several grass species: F. richardsonii, F. vivipara, A. vinealis, and the sedges C. bigelowii and Equisetum pretense. The wetlands are dominated by sedges such as Carex nigra, Carex rariflora and C. bigelowii. Also, Agrostis stolonifera and Salix sp. are found in drier parts of the wetlands. Carex lyngbyei is found in wetlands and this species was critical for hay-making. Heathland soils are classified as Andisols (Arnalds, 2008; Arnalds, et al., 1995). This soil type contains high proportions of sand (74%), low proportions of clay (1.3%) and is highly susceptible to erosion (Arnalds et al., 1995). Soil on eroded barren land is mainly sand and gravel (glacier moraine). Soil particles that have been transported to another site can be classified as Regosol. This soil type is made from unconsolidated material, mixed with sand and it does not have distinct horizons.

3. Anthropogenic disturbances 3.1. Land use in the past The history of Settlement in Iceland is well documented. Haukadalur farm was built early during the human settlement (874 AD) (Greipsson, 1969). The extensive heathlands on the highland plateau provided the farm with vast grazing land (Greipsson, 1969). The wealth of the farm was based on large woodlands at lower elevations and the extensive heathlands. The woodland surrounding the farm was used extensively as fuel and to make brown-coal. This practice has probably enhanced the wealth and status of the farm. Similarly, the status of farms in the north of Iceland was differentiated by their access to such fuel resources (Simpson et al., 2003). Past land use has undoubtedly contributed to the degradation of the heathland ecosystem. The land-use was inferred from historical and anecdotal evidence and derived from farmers in the region through semi-structured interviews conducted in 1972–2010 AD. The heathland was important for livestock grazing for the early settlers of the region and in fact formed the basis of their subsistence (Greipsson, 1969). Livestock grazing was mainly by sheep. Local names (i.e. Lambahraun, Lambavaðsheiði, Héðinssel and Nátthagasel) indicate previous land use and the ancient practice of shieling (Fig. 3). Local names (i.e. Skógarhlíð) also, indicate birch-woodlands that have long vanished (Fig. 3). Most local names on vegetated land of the west, south and east parts of the study area relate to ancient farming practices. The grazing pressure was estimated from the number and location of grazing-huts (Heiðarhús) and shiels (Sel) in the area

Fig. 1. a). Map of Iceland indicating the location of the study area (within the frame). Location of main volcanoes (Mt. Hekla, Mt. Katla and Veiðivötn) and Langjökull glacier (L) are shown. Glaciers are shown in white color within the island. b) Enlarged map of the study area shows Lake Hagavatn and Lake Sandvatn (blue color) as well as the Rivers Far, Ásbrandsá, Sandá and Jarlhettukvisl (blue lines) and glacier in yellow color. Mountains above 500 m a.s.l. are shown in brown color. Contour lines are shown at 100 m intervals.

S. Greipsson / Catena 98 (2012) 41–54

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a

L

Mt. Hekla Veidivotn

N

Mt. Katla

E

W S

b

Hagafellsjokull-Eystri

Mt. Hagafell Lake Hagavatn 400

Mt. Einifell 300

Lake Sandvatn

400

300 400

Mt. Sandfell

N E

W Haukadalur

200

S

1 km

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S. Greipsson / Catena 98 (2012) 41–54

(Fig. 3). Grazing pressure during winter was high on the south and east parts of the area. The intensive land use was influenced by traditional use of the land by settlers. This included the use of siels for housing livestock which were small huts kept some distance from the farm. The siels were inhabited during the summers and their main purpose was to process sheep milk. Some of these siels were also used for the remainder of the year as grazing-huts. Usually the grazing-huts were located nearby a wetland or a meadow where fodder (hay) could be processed. Twigs of shrubs were also collected from the heathland and used as fodder during winter. Twigs were also valuable during the free-range winter grazing period. The grazing-huts usually only housed small numbers of sheep (20–50 animals). Intensive grazing pressure was therefore maintained throughout the year. Usually the sheep housed in grazinghuts were only hay-fed during days of continuous snow cover; otherwise the animals were left free-range grazing. Additional hay was also brought from the farm to the grazing-huts, especially during severe winters and prolonged snow cover during spring. During the 19th century ~6 grazing-huts and two sheep-shelters were located on Haukadalsheiði (Fig. 3). The grazing-huts in this area were gradually abandoned due to the catastrophic soil erosion. In 1929 AD, few grazing-huts were still used in Haukadalsheiði and their ruins can still be located (Fig. 3). The practice of shieling was gradually abandoned in the 19th century and the use of grazing-huts was abandoned in the 20th century. Also, farming practices were changing and sheep were

Fig. 2. Wind-rose derived from data collected from an automatic weather station at Gullfoss (5 km east of the study area). Numbers (1–9) indicate frequency of wind directions. Data collected between Nov. 6, 2001 and Sept. 16, 2009. Data provided by the Icelandic Meteorological Office.

500

400

Lake Sandvatn 300

Stora Grjota

Mt. Sandfell 600

Skersli

400

Storamyri

300

500

Fagraflot Hedinsdael

x x x x Haukadalur

200

x

x

x

N E

W

x 1 km

S

Fig. 3. Location of sites on Haukadalsheiði. Location of shiels is indicated (x). The route Almenningsgötur is indicated with a stippled black line. The dry river course of Stóra Grjótá is indicated with a stippled blue line. Contour lines are shown at 100 m intervals.

S. Greipsson / Catena 98 (2012) 41–54

housed for practical reasons close to farms. Moreover, land cultivation and use of chemical fertilizers resulted in increased and more reliable hay harvests from hay-fields located close to farms. Therefore, the practice of free-range winter grazing was abandoned. In northeast Iceland the effect of historical winter grazing ranged from drastic soil erosion to reduced rates of erosion compared to the regional average (Simpson et al., 2004). 3.2. Degradation of the heathland: general effects of grazing Generally, in the long run intensive grazing practices on heathland ecosystems in the subarctic are unsustainable. Livestock grazing can have profound impact on vegetation composition of the original state of the heathland community. Intensive grazing reduced the dominance of willow shrubs (Salix sp.) and even excluded many broad leaf understory herbs (e.g. Geranium sylvestris), which were typically displaced by low lying herbs (Armeria maritime, R. acetocella, Cardaminopsis petraea and Silena uniflora) and thick covers of mosses (Jónsdóttir, 1984). Lichens are sensitive to grazing and disappeared from the cover as well as cover of plant litter (Jónsdóttir, 1984). Overgrazed heathland represents an alternative state (sensu Suding et al., 2004) that is maintained by high grazing pressure and the thick cover of mosses (Jónsdóttir, 1984). Changes in spatial vegetation patterns of plant communities are often the first visible indication that ecosystem degradation has occurred (Rieterk et al., 2004; Scheffer et al., 2009). Declining cover of willow shrubs most likely reduced the ecosystem's resistance to natural disturbances (sensu Gunderson, 2000; Rieterk et al., 2004; Scheffer et al., 2009). Furthermore, ecosystem degradation is expressed in damage to chemical properties of the soils, such as decline in the soil organic matter and macronutrient content (Jónsdóttir, 1984). Low species richness of the heathland communities results in only a few species in each functional group (Greipsson, 2011). Reduction in functional groups makes the ecosystem vulnerable to natural disturbances. For example, the willow S. lanata is effective in accumulating soil and sand that drifts with wind and can therefore grow vigorously near eroding soil-cliff scarps. However, this species is sensitive to grazing. Reduced cover of willow shrubs reduces the resistance of the ecosystem to soil erosion. 4. Natural disturbances

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Veiðivötn was named the Settlement layer, since soils accumulated above this tephra layer are supposed to be influenced by anthropogenic activities (Grönvold et al., 1995). Soil accumulation beneath this tephra layer took place without human influence on the environment. The next tephra fallout to completely cover the study area was 822 years later (Mt. Hekla, 1693 AD). The tephra fallout of Mt. Hekla in 1104 and 1300 AD covered only the northeastern (NE) part of Haukadalsheiði. The tephra layer of Mt. Hekla in 1104 AD is thick (>9 cm) on NE site of the study area and initiated limited localized soil erosion (Sigbjarnarson, 1969). During the eruption of Mt. Hekla in 1104 AD a thick (5–10 cm) tephra layer was spread over the area east and north of Haukadalsheiði. The next tephra fallout occurred 196 years later (Mt. Hekla 1300 AD). The eruption started on July 10 and most likely the tephra fallout had adverse impacts on vegetation in the study area. The next tephra fallout occurred 392 years later on January 13, 1693 AD and lasted for seven months. The eruption of Mt. Hekla in 1693 AD left behind a 3.5 cm thick layer of tephra on the eastern part of Haukadalsheiði, but the central and western parts of the area only received 1.5 cm tephra layer. This eruption was followed by another one 28 years later (Mt. Katla 1721 AD). The eruption of Mt. Katla in 1721 AD began May 10 and one week later a thick layer of tephra covered the study area. This tephra fallout adversely impacted vegetation in the study area, since it occurred in early spring. Yet another tephra fallout from Mt. Hekla (1766 AD) occurred 45 years later. This eruption started April 5 and lasted for two years. Again the tephra fallout had adverse effects on local vegetation. This tephra fallout (0.24 km3) was the third largest one from Mt. Hekla in historical times (Thórarinsson and Sæmundsson, 1979). This unusually high frequency of tephra fallouts between 1693 and 1766 AD on Haukadalsheiði played a critical role in driving catastrophic soil erosion (Fig. 4). Volcanic tephra that lands on eroded land can easily drift by wind and can therefore drive soil erosion. 4.2. Long term climate change and glacier fluctuation It is generally accepted that soon after the Ice Ages, ~18,000 years BP, glaciers retreated rapidly and ~8000 years BP glaciers were only confined to high mountainous areas in Iceland (Björnsson, 1979). The Settlement of Iceland took place during the Medieval Warm Period (MWP), from 900 to 1300 AD. At that time glaciers were much smaller than today. Glaciers began to advance shortly after 1250 AD as the climate became colder (Bergthórsson, 1969). During the LIA (1550–1850

High frequency of volcanic activity is characteristic for Iceland. Widespread tephra fallout is associated with volcanic activities. Some of the most active volcanoes, Mt. Hekla and Mt. Katla, are located in south Iceland (Fig. 1). Mt. Hekla is located ~50 km southeast from Haukadalsheiði (Fig. 1) and has had immense impacts on the area due to volcanic tephra fallouts. Mt. Hekla has erupted 18 times since Settlement. Most of these eruptions have produced significant amounts of volcanic tephra that have consequently been widespread over the island, depending on wind direction (Hafliðason et al., 2000). Five of these tephra fallouts have been spread over some parts of Haukadalsheiði after the Settlement (i.e. 1104, 1300, 1693, 1766 and 1970 AD) (Thórarinsson and Sæmundsson, 1979). Only one tephra fallout has occurred from Mt. Katla (1721 AD) after the Settlement (Thórarinsson and Sæmundsson, 1979). Many known tephra layers can be identified in the soil profile on Haukadalsheiði and have recently been confirmed (Jones et al., 2007). The frequency of tephra fallout on Haukadalsheiði can therefore be reconstructed. Only one tephra fallout occurred on Haukadalsheiði during the 2740 years between H5 (Mt. Hekla 7060 BP) and H4 (Mt. Hekla 4320 BP) (Sigbjarnarson, 1969). The average frequency of tephra fallout was ~400 years between H3 (Mt. Hekla 4320 BP and Veidivötn (871 AD) (cf. Grönvold et al., 1995). The tephra marker of

Relative intensity of soil erosion (%)

4.1. Frequency of volcanic tephra-fallouts

+ 0 -

Year (AD) Fig. 4. Frequency of natural disturbances: volcanic tephra fallout (red solid vertical lines) and glacial river floods (red broken lines). Relative intensity of soil erosion (blue line) was accessed by tephrochronology using sediment accumulation rate (SeAR) (mm yr− 1) between known tephra markers (Jones et al., 2007; Sigbjarnarson, 1969). Relative intensity of soil erosion was given as a proportion (%) of pre-Settlement (871 AD) (cf. Grönvold et al., 1995) SeAR. Departures from average annual temperatures from the 1961–1990 averages are shown by the black line. The temperature data were derived from Bergthórsson (1969), Hammer et al. (1980), Moberg et al. (2005), Flowers et al. (2007) and Patterson et al. (2010).

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S. Greipsson / Catena 98 (2012) 41–54

AD) glaciers advanced even further, as climate deteriorated. The climate during the LIA was characterized by severe cold inter-decadal periods with slight warming in-between (Bergthórsson, 1969; Flowers et al., 2007). Glaciers advanced and retreated in response to these interdecadal fluctuations in average temperatures during the LIA. Consequently, glacier fluctuation during the LIA most likely caused floods in glacial rivers, along with increased sediment deposits along river courses. Annual temperature rose sharply after 1918 until 1946 AD (Mackintosh et al., 2002). Then annual temperatures lowered in 1950–1980 AD with exceptionally cold averages in 1969 and 1979 AD. Since 1981 AD annual temperatures have been gradually rising. It is possible to simulate glacier fluctuation from the amplitude of climate fluctuations (Fig. 5). The advance and retreat of HagafellsjökullEystri, an outlet glacier from the Langjökull glacier, were reconstructed using data on the effect of climate on glaciers in south Iceland (Mackintosh et al., 2002) and information on the previous location of the glacier through time (Bennett et al., 2000; Green, 1952; Sigbjarnarson, 1967; Theodórsson, 1980; Thórarinsson, 1966). The glaciers in south and central Iceland are sensitive to changes in annual temperature and respond within a decade to climate changes by retreating or advancing (Björnsson and Pálsson, 2008; Björnsson et al., 2003; Kirkbridge and Dugmore, 2006; Sigurðsson and Jónsson, 1995). The reconstruction of the glacier advance is critical in understanding the onset of soil erosion through sediment deposition and sand encroachment. The outlet Hagafellsjökull-Eystri glacier probably advanced ~10 km from the time of the Settlement to its maximum size in 1890 AD (Neoglacial maximum) (Thórarinsson, 1966). The outlet Hagafellsjökull-Eystri glacier retreated ~4500 m between 1890 and 1974 AD (Fig. 5) and this is one of the most rapid retreats recorded in Iceland (Sigbjarnarson, 1976). Retreat is well documented through the use of tephrochronoloy and has been measured annually since 1934 AD (Thórarinsson, 1966). Glacial fluctuations had profound impacts on the size and location of Lake Hagavatn and caused floods (Jökulhlaup) in 1708, 1902, 1929, 1939, 1980 and 1999 AD in the River Far. Glacial retreat was, however, interrupted by short glacier surges. For example, in 1975 AD, the glacier surged forward ~1200 m but retreated until 1980 (Fig. 5). In 1980 AD, the glacier surged forward ~800 m (Theodórsson, 1980). Following these surges the sediment load drastically increased in the River Far. In 1980–1998 AD the glacier retreated ~800 m. In 1999 AD, the glacier surged again about 1200 m. Since 2000 AD, the glacier has retreated on average 70 m yr− 1 (Fig. 5). A

Fig. 5. Fluctuations of the outlet Hagafellsjökull-Eystri glacier between 1660 and 2009. Data was derived from: Green (1952), Thórarinsson (1966), Sigbjarnarson (1967), Theodórsson (1980) and Bennett et al. (2000). Arrows indicate when the advancing glacier blocked the (a) gorge of Nýifoss (1660 AD) and (b) the gorge of Leynifoss (1800 AD). Arrows also indicate when the retreating glacier caused floods in (c) the gorge of Leynifoss (1929 AD) and (d) the gorge of Nýifoss (1939 AD).

thin tephra layer from the Eyjafjallajökull eruption in 2010 AD covered the glacier and is expected to fasten the retreat. Due to ongoing global warming the Langjökull glacier is expected to disappear within the next two centuries (Björnsson and Pálsson, 2008). This will of course have dramatic effects on the landscape. The discharge of glacial rivers is expected to increase during the next century. As the glacier retreats rapidly new ephemeral lagoons will most likely appear and new outwash sand plains will be the source of sand drift further inland. 4.3. Frequency of glacial river floods Fluctuations in glaciers can result in glacial river floods (Jökulhlaup). Such floods erode vegetated embankments and deposit ample sediment and sand bars in river courses and lakes. Therefore, glacial river floods have the potential to initiate sand encroachment. Fluctuations of the outlet Hagafellsjökull-Eystri glacier had dramatic effects on the discharge of the River Far. At ~ 1600 AD the glacier surged into Lake Hagavatn and glacier water flowed out of Lake Hagavatn into the River Far. A short interval (~50 years) without any water-discharge out of Lake Hagavatn occurred ~ 1660 AD when the glacier blocked Nýifoss gorge (Green, 1952). This event initiated massive sand drift from the sand-plains along the River Far. During 1800–1929 AD there was no regular outflow of Lake Hagavatn and the catchment basin of Lake Sandvatn was not continuously inundated with water and large sand-plains developed. Regular water discharge from Lake Hagavatn into the River Far was established again following the 1929 AD flood. Five catastrophic floods have been recorded in the River Far in historical times. The first recorded flood occurred in 1708 AD. The advancing glacier blocked water from flowing through the Nýifoss gorge ~1660 AD; the water found a new outlet through the Leynifoss gorge in 1708 AD. Catastrophic glacial river floods were generally more frequent following glacial recession after 1890 AD (Björnsson, 1979). After 1890 the glacier began to retreat and in 1902 AD a flood occurred as water could temporarily flow through Leynifoss gorge (Fig. 6a). It is estimated that the 1902 flood released ~45 million m 3 of sediment laden glacial-water (Sigbjarnarson, 1967) (Table 2). The water course of the River Far could not carry this amount of flood-water and consequently embankments were severely eroded. In 1929 AD, the glacier had retreated behind Leynifoss gorge and a massive flood occurred (Fig. 5). Following the 1929 AD flood the elevation of Lake Hagavatn was ~447 m (Table 2). The 1929 AD flood eroded a new outlet for Lake Sandvatn. The new river was named Sandá (Sand-river) and this event lowered the level of Lake Sandvatn and therefore increased the size of the sand-plains. In 1939 AD the glacier had retreated behind the Nýifoss gorge and another flood occurred (Fig. 5). Following the 1939 AD flood the elevation of Lake Hagavatn was reduced even further to 439 m (Table 2). During the 1929 and 1939 AD floods about 65 million m 3 of water bursted (as jökulhlaup) in one and three days respectively (Sigbjarnarson, 1967). The average discharge of the River Far running out of Lake Hagavatn is ~10 m 3 s− 1 or 864,000 m 3 of water per 24 h. Therefore, the water-discharge during the glacial bursts in 1929 and 1939 AD was ~75 times the average water volume. The outlet of Nýifoss gorge had eroded further during the 1975 and 1980 AD floods and now the elevation of Lake Hagavatn is ~435 m. The glacial river floods in 1902–1939 AD made up the second cluster of natural disturbances (Fig. 4). The water level of the River Far fluctuates substantially. This in turn leads to exposed sand-plains during low water levels. The annual fluctuation in the water level of the River Sandá is shown in Fig. 7. The River Sandá is today the only outflow of Lake Sandvatn and it contains waters from the rivers Far and Jarlhettukvísl. The water discharge of Sandá was recorded during June 15, 2007 AD and January 28, 2010 AD. The average water discharge of Sandá was 14.3 m3 s − 1 (Fig. 7). The highest water level usually peaks in October and the lowest water level is consistently found in early spring (March and April) (Fig. 7). The water discharge peaked at 55 m 3 s− 1

S. Greipsson / Catena 98 (2012) 41–54

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a

Langjokull

Hagafell

c

Hagavatn

b

a

Einifell

Far

b

Sandfell c

Einifell

b

Far a

Nyifoss gorge

Fig. 6. a) Aerial view of the Langjökull glacier and Hagavatn. The gorges Nýifoss (a) and Leynifoss (b) are indicated with arrows. The maximum extent of the glacier in 1890 is indicated by an arrow (c). The mountains Einifell and Hagafell along with the River Far are shown. The long black arrow indicates the direction of the cross section shown in Fig. 8. b) View from Nýifoss Gorge toward Mt. Sandfell. The maximum extent of the glacier in 1890 is indicated by arrow (a). The dry river course from Leynifoss Gorge is indicated by arrow (b). The maximum extend of the eroding front is indicated by arrow (c). The long white arrow shows the main direction of sand encroachment from the outwash sand plain of the River Far next to Mt. Einifell.

in October 2007 AD and 91 m 3 s− 1 in October 2008 AD (Fig. 7). The greatest flood during this period occurred on December 26, 2007 AD when 106.9 m3 s − 1 of water were recorded (Fig. 7). The lowest water level (6 m 3 s − 1) was recorded in April 2008 AD (Fig. 7). Minor floods were associated with ephemeral lakes that typically form in front of rapidly retreating glaciers (Björnsson, 1976). Ephemeral lakes formed along the Langjökull glacier as it gradually retreated. These

lakes are included in the watershed of River Far and are responsible for a series of floods. In 1884 AD, three small ice-dammed lakes located in Jarlhettudalur east of Lake Hagavatn released ~5 million m3 of water. In 1950–1960 AD three minor floods were recorded in the River Far. These floods were associated with drainage of ephemeral glacier lakes located west of Lake Hagavatn in front of the outlet HagafellsjökullYtri glacier (Fig. 6a).

48

S. Greipsson / Catena 98 (2012) 41–54

Table 2 Physical properties of the flood water during glacial water burst (Jökulhlaup) and changes in Lake Hagavatn. Year

Lake size

Lake elevation

Volume of flood water

Sediment

(AD)

(km2)

(m)

(million m3)

(tons)

1902 1929 1939 1980 1999

13 10 6 5 5

452 447 439 435 435

45 65 65 25 6

270,000 390,000 390,000 150,000 36,000

4.4. Hydrological changes Glacial recession after 1890 AD resulted in critical hydrological changes in the study area. Especially the changes in the location and size of Lake Hagavatn following the floods of 1929 and 1939 AD. The Neoglacial maximum advancement in 1890 resulted in the maximum size of Lake Hagavatn (13 km 2) at an elevation of 457 m. In 1800–1929 AD there was no regular outflow of Lake Hagavatn and drainage was by groundwater flow through porous bedrock (Huddart and Bennett, 2000). Most of this groundwater fed into the River Stóra-Grjótá (Fig. 3). It is estimated that the discharge of this river was ~10 m3 s− 1. Consequently, water-level rose in the study area and several bogs formed. Glacial retreat has resulted in lower groundwater levels on Haukadalsheiði; manifested by drying of the river StóraGrjótá shortly after 1929 AD. In addition, water levels in smaller streams (e.g. Litla-Grjótá, Klofningaskurður and Fljótsbotnaskurður) in the area were drastically reduced (Fig. 3). These streams reduced the rate of soil erosion by acting as barriers for the advancement of the sand drift. The River Stóra-Grjótá acted as a barrier for the advancement of the soil erosion further south in the study area. The glacial flood of the River Far in 1929 AD resulted in a new outlet of Lake Sandvatn through the River Sandá (Fig. 1b). This event lowered the water level in Lake Sandvatn. Consequently, sandplains of Lake Sandvatn increased in size. Also, the groundwater level in the area south of Lake Sandvatn changed; nearby ponds and wetlands dried. Shortly after 1929 AD the largest pond in the study area (Norðlingatjörn) dried. Groundwater lowering also resulted in drying of wetlands, which promoted soil erosion across them. Remnants of wetland soil (Histosol) can be found today where these

wetlands were once located. In particular, the water level was lowered of the largest bogs Stóramýri (~45 ha) and Héðinsdæl (~10 ha) (Fig. 3). 4.5. Glacial (katabatic) winds Persistent katabatic winds blow away from the major glaciers in Iceland (Agustsson et al., 2007). Katabatic winds are generated by heat differences between the glacier and the surrounding barren highland plateau. The average July temperature (1961–1991 AD) was 9 °C for the study area, but only 4 °C for the highest plateau of Langjökull glacier (Flowers et al., 2007). This heat difference generates flow of cold air off the glacier (from an elevation >1300 m) toward the warmer highland plateau of the study area at a much lower (200–350 m) altitude (Fig. 8). Diurnal heat differences between the glacier and the barren highland plateau can be exaggerated further. The basaltic black sands and barren gravel and rock quickly absorb heat during sunny days. Surface temperatures on black basaltic sands can easily be elevated by 40 °C above the ambient air temperature during sunny days. Under such conditions during the summer local heat lows and convections are formed as cold air starts to flow from the glacier (Ashwell, 1986). Katabatic winds were recorded from Langjökull glacier as they blew onto Haukadalsheiði in August and September, 1956 AD (Ashwell and Hannell, 1960). Weather stations were located on a transect from the north outwash plain of the River Far (close to Mt. Einifell) at an elevation of 395 m, on Mt. Hagafell at 959 m and the cap of the glacier at 1145 m and 1330 m (Ashwell and Hannell, 1960). Recordings were made during the calm and sunny days of August 15 and 16, 1956 AD (Ashwell and Hannell, 1960). On the outwash plain the temperature rose from 8.1 °C in the early morning (0600) to 17.8 °C in the afternoon (1730) and then dropped to ~5 °C at 0500 (Ashwell and Hannell, 1960). At the same time temperatures remained low (−3.3–3.9 °C) on the top of the icecap (Ashwell and Hannell, 1960). On the outwash plain wind started to blow from the glacier at noon and gradually increased until 1700 when it reached 6 m s − 1 and then gradually declined until winds reached 0 at 0100 (Ashwell and Hannell, 1960). Winds above the critical speed of 6 m s− 1 were able to cause sand drift (Sandfok) from the outwash plain, even during light rain and drizzle (Ashwell and Hannell, 1960). Silt moved as convectional dust-storms (Mistur) from the outwash plain by katabatic winds (Ashwell, 1966, 1986). Sigbjarnarson (1969) estimated that dry deflated soil (Moldir) can be carried in wind suspension (Moldrok) at 0.2 m s− 1. On the outwash plain under clear skies mirages followed by dust devils formed in the

1400

Water flow (m3 s-1)

1200

Elevation (m)

1000 800 600 400 200 0 0

5

10

15

20

25

30

35

40

Km

Time Fig. 7. Time series of daily fluctuations of water flow (m3 s− 1) of the River Sandá. Data collected between June 15, 2007 and Jan. 28, 2010. Data provided by the Icelandic Meteorological Office.

Fig. 8. Elevation profile across Haukadalsheiði, Mt. Brekknafjöll, Hagavatn and the Langjökull glacier. The transect is 36.2 km long in a straight north–south direction starting at Mt. Sandfell. Solid line indicates surface elevation and elevation of Lake Hagavatn. Broken line indicates elevation of Langjökull glacier. Data on glacier thickness was derived from Flowers et al. (2007).

S. Greipsson / Catena 98 (2012) 41–54

afternoon before a dust-storm of ~100 m height broke out (Ashwell, 1966; Hannell and Ashwell, 1958). A dust-storm in the center of the study area on a sunny day, (29 July 1965 AD) reduced the visibility to only 20–30 m (Sigbjarnarson, 1969). The silt was transported long distances but the coarse sand drifted ~10 m into the heathland vegetation (Sigbjarnarson, 1969). In 1956 AD, katabatic winds extended at least 10 km away from the Langjökull glacier (Ashwell, 1966). Wind eroded soil was redistributed over the land south of the study area and some of the soil was carried in wind suspension (Moldrok) >70 km toward the sea (Sigbjarnarson, 1969). Ashwell (1966) found an overlap between area of high wind erosion and where maximum convection is caused by large glaciers in Iceland. This appears to be the area on the highland plateau where the most intensive soil erosion has taken place. 4.6. Sand encroachment The main direction of sand encroachment in the study area is by north winds that blow away from the glacier (Gísladóttir et al., 2005). Sand transportation by wind was recently measured on a nearby (3 km west) highland plateau (400–500 m elevation) south of the outlet of Hagafellsjökull-Ytri glacier where a glacial lagoon that dried in 1960 AD provides a constant source of sand and silt that drifts ~ 16 km south away from the glacier with devastating effects on vegetated lands (Gísladóttir et al., 2005). The study showed that sand transportation from the dried lake bed was initiated when wind speed >6.2 m s − 1 (Gísladóttir et al., 2005). Such wind speed is not uncommon in the study area (Ashwell, 1966). When wind speed exceeded 10 m s − 1sand transportation increased almost exponentially and at high velocity (17 m s − 1) the sand flux exceeded 1 tonne m − 1 h − 1 (Gísladóttir et al., 2005). Much less wind speed (>4.4 m s-− 1) was, however, needed to maintain sand flux (Gísladóttir et al., 2005). This study demonstrated that huge quantities of sand can be moved by wind over short periods of time during storms. The catchment basin of Lake Sandvatn was certainly supplied regularly with enormous deposits of sediment that formed the sand plains. The gradual infilling of the catchment basin of Lake Sandvatn through deposition of sediment from the River Far started ~1600 AD when the glacier advanced into Lake Hagavatn. Sand deposition into the catchment basin was, however, episodic and enhanced during glacial river floods of 1708, 1884, 1902, 1929 and 1939 AD (Table 2). The River Far brings high loads of sediment (1.6–6.4 kg m− 3) to the catchment basin of Lake Sandvatn (Björnsson, 1979). The sediment varies from fine silt to coarse sand. The River Far flows in braided streams over the sand-plain leaving behind enormous amount of sediment. Following the 1939 AD flood the sand-plain was ~ 2.7 km long (NNE) and ~1.3 km wide. The sand-plain acted as a constant supply of drifting sand and silt during the dry northern glacial (katabatic) winds. The sand drift was especially pronounced when the water level of the River Far was low during spring. Sand drift had been advancing onto vegetated land since ~ 1660 AD and consequently initiated three soil eroding fronts. It was estimated that in 1700 AD the eroding fronts had extended ~1200 m into the vegetated heathland (Fig. 9). The first episode of sand encroachment was most likely initiated when Nýifoss gorge of Lake Hagavatn was blocked by the advancing glacier ~ 1660 AD (Green, 1952). Consequently, the River Far remained dry until the 1708 AD flood from Leynifoss gorge. For ~ 50 years (~ 1660–1708 AD) sand drift occurred onto vegetated land from the outwash plains of the River Far. The oldest written document on the sand drift from the sand-plains of Lake Sandvatn is found in the “Land and livestock register” describing the properties of Haukadalur farm in 1709 AD (Magnússon and Vídalín, 1713). This written evidence strongly indicates that a massive sand encroachment had advanced rapidly onto vegetated land south of Lake Sandvatn and initiated soil erosion. In 1800 AD water in River Far was blocked again by the advancing

49

glacier and it marks the second episode of sand encroachment. The River Far remained blocked for 129 years (1800–1929 AD) with the exception of the 1884 and 1902 AD floods (Fig. 4). The third episode of sand encroachment occurred after the floods of 1929 and 1939 AD (Fig. 4). Following these floods sand drift was further sustained from the three outwash plains (Fig. 9). 5. Reconstruction of soil erosion with time 5.1. Advance of soil eroding fronts The rate of soil erosion was spatially reconstructed on Haukadalsheiði using information on the average progress of eroding fronts. The progress of the three main eroding fronts at different time intervals is shown in Fig. 9. The total length of eroding fronts was measured using a map from the Danish General Staff (1908) and aerial photographs from 1960 to 2008 AD. The location of each eroding front was confirmed in the field. The progress of eroding fronts between 1960 and 2008 AD was negligible due to active restoration efforts (Greipsson, 2011). The eroding front from the outwash plain of the River Far at Mt. Einifell (north) to Mt. Sandfell (south) had progressed 8.2 km in ~300 years (Figs. 6b; 9). Therefore, on average this eroding front had progressed 27.3 m yr− 1. The progress of this eroding front between 1908 and 1960 was 23 m yr− 1. This erosion front was closest to the glacier (only 2.4 km) at an elevation of 336–360 m. The second erosion front progressed 8.9 km from the sand-plains of Lake Sandvatn toward Haukadalur valley at an elevation of 218–275 m (Fig. 9). On average this eroding front progressed 29.7 m yr − 1. The progress of this eroding front between 1908 and 1960 was 25 m yr− 1. This eroding front received a constant supply of drifting sand from the sand plains of Lake Sandvatn. The third erosion front extended 7.1 km from Syðriflói toward Haukadalur valley (Fig. 9). On average this eroding front progressed 23.7 m yr− 1. The progress of this eroding front between 1908 and 1960 was 27 m yr− 1. This eroding front is, however, furthest (8.4 km) away from the glacier at an elevation of 207–275 m. The formation of outwash sand-plains along the River Far in ~1660 marks the first episode of sand encroachment onto vegetated land in the study area. Sand encroachment leads to the formation of actively eroding fronts. Sand encroachment not only buries vegetation but also accelerates detachment of soil particles and forms eroding soil escarpments (Gísladóttir et al., 2005). Soil is exposed on the cliff-scarp and wind erosion is facilitated by sand particles that creep and saltate (Arnalds, 2000). Sand encroachment was sustained by the prevailing dry northern katabatic winds (Gísladóttir et al., 2005). Catastrophic soil erosion in the study area was driven by katabatic winds that formed rapidly moving eroding fronts (Figs. 10; 11). A conceptual model was constructed to simulate the main processes during soil erosion mediated by sand encroachment in the study area (Fig. 10). As the eroding front advances the lowest layer of the soil (Moldir) is typically left behind and few remnants of vegetation (Torfur) (Figs. 10; 11b). Extensive areas of soil remnants can persist for decades and provide enormous source of wind suspension (Moldrok). Furthermore, as erosion advances gravel flats, sands and bedrock (Skersli) are left behind (Fig. 10). Finally, the drifting sand accumulates in river courses and wetlands. The eroding front only slows down or stops if rivers, rivulets or wetlands (acting as sinks for drifting sand) exist in the landscape. Lee sides of hills can also act as sand-sinks where the eroding power of the wind is slowed down or curtailed. 5.2. Progress of soil erosion assessed by tephrochronology Frequent volcanic eruptions in Iceland have produced widespread layers of volcanic tephra in the soil profile. Tephra that falls on vegetated land becomes progressively incorporated in the soil profile as a distinct layer (horizon). Distinctive tephra layers (isochrons) in the

50

S. Greipsson / Catena 98 (2012) 41–54

a

b

Mt. Einifell

Haukadalsheidi

N

Mt. Sandfell

W 1 km

E S

c

d

Fig. 9. Spatial reconstruction of soil erosion on Haukadalsheiði. Vegetated land is indicated in black, eroded land in red color, sand plains in grey and rivers and lakes blue. Direction of sand encroachment is shown by white arrows. The progress of the eroding fronts is shown with time. a) Approximate vegetation cover in 1700. b) Approximate vegetation cover in 1800, c) vegetation cover in 1908 (Danish General Staff, 1908), d) vegetation cover in 1960.

soil profile can be identified over the study area. The date of these tephra layers is known (Jones et al., 2007; Sigbjarnarson, 1969). By measuring the sediment accumulation rate (SeAR) (mm yr − 1) between known tephra markers it is possible to estimate the relative intensity of nearby soil erosion (Thórarinsson, 1979). The relative intensity of catastrophic soil erosion that occurred on Haukadalsheiði was estimated by measuring SeAR between known tephra markers of 36 soil profiles surrounding the eroded land (Sigbjarnarson, 1969). The SeAR was a slow process below the Settlement tephra layer (~871 AD) (i.e. pre-Settlement) (Sigbjarnarson, 1969). The tephra marker of Hekla 1104 AD is only found in the eastern part of the study area. Between 871 and 1104 AD the SeAR (mm yr − 1) increased by 141% compared to pre-Settlement rates (Fig. 4). Between 1104 and 1300 AD the SeAR increased 675% compared to pre-Settlement rates (Fig. 4). This sediment thickening is most likely the result of wind erosion from the highlands (closer to the glacier) north of Haukadalsheiði. Between 1300 and 1693 AD SeAR increased only 182% compared to pre-Settlement rates (Sigbjarnarson, 1969). Between 1693 and 1766 AD (Mt. Katla) SeAR

increased drastically or 929% compared to pre-Settlement rates. The tephra layer of 1693 AD can therefore be used as a marker for the catastrophic soil erosion. From 1766 to1964 AD SeAR increased even further or 1918% compared to pre-Settlement rates (Sigbjarnarson, 1969). 5.3. Rate of soil erosion assessed by historical and anecdotal evidences Historical records of the catastrophic soil erosion are scant, but Magnússon and Vídalín (1713) surveyed the region in 1709 AD and their text strongly indicated that serious soil erosion was already in progress in the north part of Haukadalsheiði. The farm Haukadalur was owned by the State Church for ~500 years (~1300–1800 AD) but was auctioned in 1794 AD when most of the heathland was already devastated by the soil erosion. Pálsson (1884) documented the abandonment of the track (Almenningsgötur) that traversed the study area (Fig. 3). This track was an important commercial route connecting north and south Iceland. In 1646 AD the track was still in use and no mention of soil erosion or sand encroachment in the

S. Greipsson / Catena 98 (2012) 41–54

a

b

51

trace the track on vegetated remnants until 1950 AD (Fig. 2). Today, it is possible to locate few of these vegetated remnants (Fig. 11b). Semi-structured interviews with farmers in the region conducted in 1972–2010 AD provided anecdotal information on the progress of the soil erosion, especially by giving account on previous land use and outlining when particular sites were eroded. For example, the bog (Stóramýri) was used in hay making for nearby shiels until 1870 AD. The bog was infilled by sand and consequently eroded by gullies (Fig. 3). Today, remnant organic soil can be located where the bog once was. Another bog (Héðinsdæl) was used for haymaking until 1929 AD (Fig. 3). Similarly, today remnants of organic soil can be found where this bog was located. Also, a dry meadow named Fagraflöt was located near Mt. Sandfell (Fig. 3). This meadow was used for hay-making until 1920 AD, but is now eroded and only barren land remains. Local names of the landscape reflect events during the soil erosion. Most local names found further north on the barren land relate to erosion phenomenon. A sequence describing the erosion process can be found in the local names: erosional soil-cliff scrap (Bakki), eroding front (Klofningar), remnant of vegetated patches (Torfur), deflated soils (Moldir) and barren lands (Skersli). The density (per km 2) of local names is highest on vegetated lands east and south of the study area, but declines sharply in the northern part of the study area that is today barren. The area south of Lake Sandvatn was first to become barren and this area has by far the lowest density of local names since it was not important for farmers after soil erosion began. 6. Discussion

c

Fig. 10. Conceptual model showing the processes of soil erosion with time across vegetated landscape. a) sand encroachment (red arrow) forming erosion front; green indicates vegetated land, brown indicates deflated soil, b) progress of the eroding front; gray indicates barren land, c) remaining deflated soil, barren land and scattered vegetation remnants.

area. The track was abandoned just before 1884 AD (Pálsson, 1884). The track was mainly abandoned due to sand drift and eroded grasslands (Grasdalir) that was critical for travelers as they began their journey across the highlands (Fig. 3). In 1929 AD the track was only found on few vegetated remnants. It was, however, still possible to

Long term adverse climate change during LIA exacerbated the impact of grazing on heathland communities. Therefore, relentless livestock grazing combined with adverse climate change could have brought the resistance of the heathland community to a critically low level. The long-term effect of the drop in average annual temperature during the LIA (Ogilvie and Jónsson, 2001), on the heathland ecosystem was twofold. First, the growing season was probably reduced by three to four weeks during the peak of the LIA (Ogilvie, 2001). This undoubtedly resulted in less vigorous plant cover. Secondly, the timberline of the birch-woodlands probably moved just below the study area (Ogilvie, 2001). Indirect effects include the increase in grazing pressure on the heathlands as a result of reduced harvest on meadows located on Haukadalsheiði. Most hay-production took place on non-fertilized meadows. Hay production was very sensitive to changes in annual summer temperatures. Therefore, decline in hay-production resulted in greater reliance on free-range grazing the year around, intensifying grazing pressure on heathlands (Ogilvie, 2001). Simulation of vegetation cover in Iceland in response to long-term climatic variations has shown oscillating vegetation cover before the Settlement and about 25% reduction in vegetation cover after the Settlement (Haraldsson and Ólafsdóttir, 2003). Catastrophic soil erosion on Haukadalsheiði was triggered in ~1660 AD by sand drift from outwash sand plains. This event was followed by two clusters of unusually frequent natural disturbances (Fig. 4). The first cluster included volcanic tephra fallouts in 1693 AD (Mt. Hekla), 1721 AD (Mt. Katla), 1766 AD (Mt. Hekla) and the glacial river flood in 1708 AD. The second cluster of unusually frequent natural disturbances included glacial river floods of 1884, 1902, 1929 and 1939 AD. It is also noteworthy that catastrophic soil erosion began during coldest period of the LIA (1590–1690 AD) (Ogilvie and Jónsson, 2001). The climate during the LIA was cold and exceptionally windy (Jackson et al., 2005); strong northerly katabatic winds drove the soil erosion. The rate of soil erosion has been measured previously in Iceland. Fridriksson (1995) measured retreat of an eroding soil scarp to be on average 0.16 m yr − 1. Higher erosion rates have, however, been

52

S. Greipsson / Catena 98 (2012) 41–54

a

Asbrandsa c

b

a

d a

b

Fig. 11. Features of soil erosion on Haukadalsheiði: a) eroding front on eastern Haukadalsheiði; vegetated land (a), soil escarpment (b), barren land (c) and deflated soil (d) are indicated. The white arrow shows the direction of the sand encroachment in a north–south direction. The aerial photo was taken in July 1982, b) soil and vegetation remnant. The surrounding area was vegetated in 1940. The soil mantle is ~ 120 cm tall and the remnant was ~ 3 m across. Mt. Einifell is on the left. The photo was taken in 1984.

recorded on eroding soil escarpments; 0.20 m yr − 1 (Dugmore et al., 2009) and 0.40 m yr − 1 (Arnalds, 2000). These studies were conducted on retreating soil-cliff scarps by measuring annual disappearance of vegetation and soil. It is, however, noted that these real-time erosion rates are too low to account for the catastrophic soil erosion that has occurred in Iceland after the Settlement. The rate of advancing eroding fronts (29.7, 24.7 and 23.7 m yr − 1) reported here appears to be exceptionally rapid. However, the erosion rates reported here reflect more realistically on the catastrophic erosion that has taken place over the past centuries on the highland plateau. Tephrochronological study was used to assess the relative intensity of soil erosion and the tephra fallout of the volcanic eruption of Mt. Hekla in 1693 AD was used as a marker for the catastrophic soil erosion on Haukadalsheiði. In a comprehensive study on the effect

of human settlement on nearby environments in the highlands surrounding Lake Mývatn in north Iceland, the relative intensity of soil erosion was documented using tephra markers in the soil profile (McGovern et al., 2007). Following human settlement in north Iceland, limited soil erosion was recorded, but slow erosion was recorded throughout the Middle Ages. A massive volcanic eruption in 1717 AD was used as a marker for the initiation of catastrophic soil erosion in north Iceland. The eruption in 1717 came shortly after the coldest period of LIA (1590–1690 AD) (Ogilvie and Jónsson, 2001). The SeAR after 1717 AD was found in selected sites to be 2000% greater compared to SeAR between 1477 and 1717 AD (McGovern et al., 2007). A conceptual model dealing with catastrophic soil erosion in Iceland has recently been proposed by Dugmore et al. (2009). The model attempts to explain landscape erosion of heathlands through

S. Greipsson / Catena 98 (2012) 41–54

degradation of vegetation cover by livestock grazing. This in turn initiates localized soil erosion that develops into eroding cliff-scarp that have slow erosion rate (max 0.2 m yr − 1). However, total erosion is supposed to be higher, since high density of eroding cliff scarps can develop throughout the landscape (Dugmore et al., 2009). An important component of this model is that soil erosion is initiated at high altitude and then spreads to lower altitudes (Dugmore et al., 2009). This study provides integrated approach on catastrophic soil erosion where resistance of heathland community to soil erosion was lowered by livestock grazing and adverse long-term climatic change. Soil erosion was triggered by sand encroachment and further enhanced by an unusually high frequency of natural disturbances (volcanic tephra fallout and glacial river floods) where strong northerly katabatic dry winds drove the soil erosion. Landscape model (i.e. “erosion cell”) has been used to describe soil erosion processes (Westoby, 1987). Soil erosion is generally viewed as a multi-stage process involving detachment of soil-particles, transfer of particles over the landscape and eventual deposition of soilparticles into sinks (Lal, 2001). An “erosion cell” includes production zone, transfer zone and sink (Westoby, 1987). In this study, the production zone is the outwash sand-plains along the River Far. The transfer zone is the barren land along the north wind direction. Sinks are rivers, rivulets and wetlands that transect the landscape. The power of the sand encroachment is most likely reduced with distance from the sand source and with increased number of sinks in the landscape. The eroded soil is lost via wind (silt and clay) and water erosion (sand) out of the area. Sand and eroded soil is only temporarily deposited along the transfer zone, but it is eventually lost in sinks (rivers and rivulets). Sigbjarnarson (1969) emphasized such combined activity of water and wind in soil erosion on Haukadalsheiði. The results of this study have implications for restoration ecology, especially in prioritizing restoration efforts. First, increasing resilience of heatland communities by restricting livestock grazing will most likely decrease the risk of soil erosion. Secondly, it is critical to be able to identify “triggers” of soil erosion and through active restoration efforts curtail their impacts. 7. Conclusion Catastrophic soil erosion on Haukadalsheiði, south Iceland was triggered ~ 1660 by sand drift from outwash sand plains along the glacial River Far. Relentless, grazing by livestock and adverse climate change (LIA) undoubtedly lowered the resistance of heathland communities to soil erosion. In addition, Andisol are sandy and have low resistance to erosion. Unusually high frequency of natural disturbances (volcanic tephra fallout and glacial river floods) sustained soil erosion that resulted in total destruction of heathland ecosystem. The progress of eroding fronts was exceptionally rapid (max. 29.7 m yr − 1) and was driven by dry northern katabatic winds. The widespread soil erosion that has taken place on the highland plateau in Iceland could in part be explained by similar processes as described in this study. Acknowledgments This work was partly supported by the Soil Conservation of Iceland (Landgræðsla ríkisins). I would like to thank Dr. N. Pullen-Holst and Dr. M. Fullen for their help in improving this paper. References Agustsson, H., Cuxart, J., Mira, J., Olafsson, H., 2007. Observations and simulation of katabatic flows during a heatwave in Iceland. Meteorologische Zeitscrift 16, 99–110. Arnalds, A., 1987. Ecosystem disturbance in Iceland. Arctic and Alpine Research 19, 508–513.

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Arnalds, O., 2000. The Icelandic “rofabard” soil erosion features. Earth Surface Processes and Landforms 25, 17–28. Arnalds, O., 2008. Andisols. In: Chesworth, W. (Ed.), Encyclopedia of Soil Science. Springer, Dordrect, pp. 39–46. Arnalds, O., Hallmark, C.T., Wilding, L.P., 1995. Andisols from four different regions of Iceland. Soil Science Society of America Journal 59, 161–169. Ashwell, I.Y., 1966. Glacial control of wind and of soil erosion in Iceland. Annals of the Association of American Geographers 56, 529–540. Ashwell, I.Y., 1986. Meteorology and duststorms in central Iceland. Arctic and Alpine Research 18, 223–234. Ashwell, I.Y., Hannell, F.G., 1960. Wind and temperature variations at the edge of an ice-cap. Meteorological Magazine 89, 17–24. Bennett, M.R., Huddart, D., McCormick, T., 2000. An integrated approach to the study of glaciolacustrine landforms and sediments: a case study from Hagavatn, Iceland. Quaternary Science Reviews 19, 633–665. Bergthórsson, P., 1969. 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